Heading 1 1

Professional Paper 1750

U.S. Department of the Interior

U.S. Geological Survey

A Volcano Rekindled: 

the Renewed Eruption of Mount St. Helens, 2004–2006
2 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

COVER

Mount St. Helens crater and new dome, September 12, 2006. Oblique aerial view from the north. 

Crater rim is about 2 km across. Spines 4, 5, and 7 of the new dome (see labeled sketch below) 

form conical masses that have nearly or completely buried other spines of the eruptive sequence. 

At slightly lower altitude, Crater Glacier’s east and west arms nestle the craggy 1980–86 dome as 

they plunge northward from the crater floor. USGS photo by W.E. Scott.
A Volcano Rekindled: The Renewed 

Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

Professional Paper 1750

U.S. Department of the Interior

U.S. Geological Survey
U.S. Department of the Interior

DIRK KEMPTHORNE, Secretary

U.S. Geological Survey

Mark D. Myers, Director

U.S. Geological Survey, Reston, Virginia: 2008

This report and any updates to it are available online at:

http://pubs.usgs.gov/pp/1750/

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Telephone: 1-888-ASK-USGS (1-888-275-8747)

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Telephone: 1-888-ASK-USGS (1-888-275-8747)

Any use of trade, product, or firm names is for descriptive purposes only and does not imply 

endorsement by the U.S. Government.

Although this report is in the public domain, permission must be secured from the individual copyright owners to 

reproduce any copyrighted material contained within this report.

Suggested citation:

Sherrod, D.R., Scott, W.E., and Stauffer, P.H.,eds., 2008, A volcano rekindled; the renewed eruption of Mount St. Helens, 

2004-2006: U.S. Geological Survey Professional Paper 1750, 856 p. and DVD-ROM.

Cataloging-in-publication data are on file with the Library of Congress (http://www.loc.gov/).

Produced in the Western Region, Menlo Park, California

Manuscript approved for publication, August 27, 2008

Text edited by Peter H. Stauffer, Tracey L. Suzuki, 

James W. Hendley II, and George A. Havach

Layout by Judy Weathers and Susan Mayfield 
iii

Foreword

Late September and early October 2004 saw intense seismic unrest, rapid ground deformation, and 

explosions heralding the renewal of lava-dome growth in the crater of Mount St. Helens, Washington. Thus ended the relative calm of two decades since the renowned 1980–86 eruptions. Memories of the May 18, 1980, tragedy and its aftermath prompted a high level of concern among the 

public, public officials, and the media. Response to this unrest and eruption was led by scientists 

at the U.S. Geological Survey’s (USGS) David A. Johnston Cascades Volcano Observatory (CVO). 

Established following the 1980 eruption, CVO was named for the spirited young USGS volcanologist who lost his life during the initial minutes of the May 18 cataclysm. USGS scientists worked 

in close cooperation with CVO’s long-time partner in monitoring the volcanoes of Washington and 

Oregon, the Pacific Northwest Seismic Network at the University of Washington. This cooperation ensured that local land and emergency managers, public officials, key government agencies, 

media representatives, and the public received the information needed to protect life and property, 

as well as to appreciate the spectacle of an erupting volcano. USGS scientists from sister volcano 

observatories—Alaska, Hawaiian, Long Valley, and Yellowstone—and the broader USGS community contributed significantly to the effort. USGS scientists relied on lessons learned from responding to previous volcanic crises in the United States and abroad. Scientists from more than a dozen 

academic institutions undertook valuable scientific studies of the eruption. New instrumentation 

from the National Science Foundation-sponsored Plate Boundary Observatory aided in geodetic 

monitoring. The response was bolstered by vast improvements in monitoring technology, information management, and understanding of volcanic processes that were spawned by the 1980 eruption and other eruptions during the late 20th century.

The USGS is proud to have cooperated closely with the Gifford Pinchot National Forest, home of 

the Mount St. Helens National Volcanic Monument, the Washington State Emergency Management Division, county emergency-management agencies, the National Weather Service, the 

Federal Aviation Administration, and the Federal Emergency Management Agency in fulfilling its 

mandated role under the Disaster Relief Act of 1974 (P.L.93–288) “to provide technical assistance 

to State and local governments to ensure timely and effective disaster warning is provided.” This 

multichaptered Professional Paper, released as the most recent period of the Mount St. Helens 

lava-dome building ends after nearly three and one-half years of continuous activity, adds to 

our understanding of how volcanic systems work, and how scientists must work together with a 

broad array of agencies and public officials to mitigate risk from volcanic activity.

Occurring more than a quarter century ago, the catastrophic 1980 eruption of Mount St. Helens 

may seem to some as just a historical footnote. This Professional Paper documenting both the 

volcano’s recent activity and the need for coordinated warning and disaster response should 

serve the public well as a reminder of the challenges and consequences of living near an active 

volcano. Previous USGS Director Dallas Peck said it best: “The lessons of Mount St. Helens 

must not—and will not—be forgotten.”

Mark D. Myers

Director,

U.S. Geological Survey
iv

Anthony I. Qamar, 1943–2005

Research Assistant Professor, University of Washington

Co-Principal Investigator, Pacific Northwest Seismic Network

Washington State Seismologist, 1988–2005

Tony Qamar poses near seismic equipment 

he helped to install on the flanks of 

Mount St. Helens in the summer of 2005

Anthony “Tony” Qamar was a highly 

regarded, generous scientist whose career 

encompassed the Mount St. Helens eruptions of 1980–86 and 2004 (ongoing at this 

writing, June 2007). Tony and his colleague, 

Dan Johnson, were killed in a tragic auto 

accident October 4, 2005, while traveling to 

Washington’s Olympic Peninsula to retrieve 

a GPS receiver.

Tony loved mountains, particularly 

volcanoes. Unbeknownst to many colleagues, 

Tony was an accomplished climber with 

major first ascents to his credit. He often 

combined his professional world with his 

love of the outdoors by accepting the task of 

repairing remote and inaccessible seismographs, such as those high on Mount Rainier 

and Glacier Peak. In 1980 he was one of the 

first scientists to appear at Mount St. Helens 

in response to precursory seismicity. Camping in the snow with a portable seismograph 

and thermal infrared imager, he witnessed and 

recorded many of the early phreatic eruptions.

In September 2004, Tony again became 

immersed in volcano seismology as Mount 

St. Helens reawakened. Moving from his 

regular office, he set up “camp” in a small, 

cramped room across from the seismology 

lab, where he was a resource for the staff 

and seismology students at the University of 

Washington’s Department of Earth and Space 

Sciences and also for news media who used 

the lab during the early chaotic months of the 

eruption. He developed several automated 

computer codes (in FORTRAN) that enabled 

the staff of the Cascades Volcano Observatory and the Pacific Northwest Seismic 

Network to make sense of the overwhelming 

number of earthquakes and to track changes 

in real time. The paper on which he is first 

author in this volume (chap. 3) is based on 

these computational efforts and is, we hope, 

close to the paper he intended to write.

While fundamentally a seismologist interested in the stresses that trigger 

earthquakes, Tony let his curiosity drive 

his research in diverse directions. He published papers on the seismic structure of the 

Earth’s inner core, seismic signals generated by 

meteors or the space shuttle, ice quakes from 

glaciers recorded by high-frequency seismographs, and the slowly straining tectonic plates 

as detected by GPS receivers.

Tony was an excellent classroom teacher 

and even better mentor. As a research collaborator he was generous with his time and 

ideas, interested more in working with others 

to achieve a better understanding of a scientific 

problem than in claiming his piece of the intellectual pie through first-author publications. 

Those around him were enriched by his presence, knowing he was a friend first and professor second. Students frequently approached him 

with questions, both technical and personal, and 

they always left with their answers in hand and 

their egos intact. The hole left by his passing is 

both wide and deep.
v

Contributors to This Professional Paper

Cascades Volcano Observatory

Beeler, Nick M.

Couchman, Marvin R.

Denlinger, Roger P.

Doukas, Michael P.

Driedger, Caroyn L.

Dzurisin, Daniel

Endo, Elliot T.

Gardner, Cynthia A.

Gerlach, Terrence M.

Gooding, Daniel J.

Iverson, Richard M.

Iwatsubo, Eugene Y.

Janda, Christine G.

Kingsbury, Cole G.

LaHusen, Richard G.

Lisowski, Michael

Lockhart, Andrew B.

Logan, Matthew

Lu, Zhong

Major, Jon J.

Mastin, Larry G.

McGee, Kenneth A.

Moran, Seth C.

Pallister, John S.

Roeloffs, Evelyn

Schilling, Steve P.

Scott, William E.

Sherrod, David R.

Spicer, Kurt R.

Swinford, Kelly J.

Thornber, Carl R.

Vallance, James W.

Walder, Joseph S.

Wolfe, Edward W.

U.S. Geological Survey

1300 SE Cardinal Court

Vancouver, WA 98683

Menlo Park

Bergfeld, Deborah

Calvert, Andrew T.

Clynne, Michael A.

Evans, William C.

Evarts, Russell C.

Fleck, Robert J.

Lanphere, Marvin A.

U.S. Geological Survey

345 Middlefield Road

Menlo Park, CA 94025

Denver

Brownfield, Isabelle K.

Lowers, Heather A.

Meeker, Gregory P.

Messerich, James A.

Thompson, Ren A.

U.S. Geological Survey

Denver Federal Center

Denver, CO 80225

Alaska Volcano Observatory

Neal, Christina A.

Schneider, David J.

Wessels, Rick L.

U.S. Geological Survey

4200 University Drive

Anchorage, AK 99508

Hawaiian Volcano Observatory

Poland, Michael P.

U.S. Geological Survey

PO Box 51

Hawaii National Park, HI 96718
vi

University of Washington

Creager, Kenneth C.

Crosson, Robert S.

Donnelly, Carrie T.

Malone, Stephen D.

McChesney, Patrick J.

Norris, Robert D., USGS

Qamar, Anthony I. *

Steele, William P.

Thelen, Weston A.

Wright, Amy K.

Earth and Space Sciences

Box 351310

Seattle, WA 98195

U.S. Department of Agriculture,

Forest Service

Frenzen, Peter M.

Knappenberger, Thomas H.

Lapcewich, Dennis

Matarrese, Michael T.

Gifford Pinchot National Forest

10600 N.E. 51st Circle

Vancouver, WA 98682

Brown University

Devine, Joseph D., III

Rutherford, Malcolm J.

Department of Geological Sciences

324 Brook Street

Providence, RI 02912

Iowa State University

Iverson, Neal R.

Moore, Peter L.

Department of Geological and 

Atmospheric Sciences

Ames, IA 50011

McGill University

Berlo, Kim

Earth & Planetary Sciences,

McGill University

3450 University St.

Montreal, Quebec, Canada

H3A 2A7

Oregon State University

Kent, Adam J.R.

Department of Geosciences, OSU

Wilkinson Hall

Corvallis, OR 97331

Portland State University

Broderick, Cindy A.

Streck, Martin J.

Department of Geology, PSU

PO Box 751

Portland, OR 97207

Southern Methodist University

Quick, James E.

Office of Research and Graduate 

Studies

P.O. Box 750240 

Dallas, TX 75275-0240

University of Bristol

Blundy, Jon

Department of Earth Sciences

Wills Memorial Building

Bristol BS8 1RJ, UK

*

 Deceased
vii

University of California

Cooper, Kari M.

Geology Department, UC Davis

One Shields Avenue

Davis, CA 95616

Herriott, Trystan M.

Department of Earth Science,

UC Santa Barbara

Webb Hall, Building 526

Santa Barbara, CA 93106

University of Cambridge

Edmonds, Marie

Downing Street

Cambridge CB2 3EQ, UK

University of Iowa

Reagan, Mark K.

Rowe, Michael C.

Wortel, Matthew

Department of Geoscience

121 Trowbridge Hall

Iowa City, IA 52242

University of Memphis

Horton, Stephen P.

Center for Earthquake Research 

and Information

3890 Central Ave

Memphis, TN 38152

University of Pittsburgh

Ramsey, Michael S.

Department of Geology & 

Planetary Science

200 SRCC Building

Pittsburgh, PA 15260

University of Oregon

Cashman, Katharine V.

Department of Geological Sciences, 

UO

1272 University of Oregon

Eugene, OR 97403

Western Washington University

Caplan-Auerbach, Jacqueline

Geology Department, WWU

516 High Street

Bellingham, WA 98225

Clark County

Needham, Deborah H.

Clark Regional Emergency Services 

Agency

710 W 13th St.

Vancouver, WA 98660

American Museum of Natural 

History

Mandeville, Charles W.

American Museum of 

Natural History

Central Park West at 79th Street

New York, NY 10024

Washington Military Department

Harper, Robert B.

Emergency Management Division

Building 20

Camp Murray, WA 98430
viii

 Overview

 1 Overview of the 2004 to 2006, and continuing, eruption of Mount 

St. Helens, Washington .................................................................................................................... 3

 By W.E. Scott, D.R. Sherrod, and C.A. Gardner

 Seismicity of the eruption

 2 Seismicity associated with renewed dome building at Mount St. Helens, 2004–2005........ 27

 By S.C. Moran, S.D. Malone, A.I. Qamar, W.A.Thelen, A.K. Wright, 

 and J. Caplan-Auerbach

 3 Near-real-time information products for Mount St. Helens—tracking 

the ongoing eruption ........................................................................................................................61

 By A.I. Qamar, S.D. Malone, S.C. Moran, W.P. Steele, and W.A. Thelen

 4 Absolute and relative locations of earthquakes at Mount St. Helens, Washington, using 

continuous data: implications for magmatic processes........................................................... 71

 By W.A. Thelen, R.S. Crosson, and K.C.Creager

 5 Broadband characteristics of earthquakes recorded during a dome-building eruption at 

Mount St. Helens, Washington, between October 2004 and May 2005 ................................. 97

 By S.P. Horton, R.D. Norris, and S.C. Moran

 6 Seismicity and infrasound associated with explosions at Mount St. Helens, 2004–2005 .. 111

 By S.C. Moran, P.J. McChesney, and A.B. Lockhart

 7 Seismic-monitoring changes and the remote deployment of seismic stations (seismic 

spider) at Mount St. Helens, 2004–2005 .................................................................................... 129

 By P.J. McChesney, M.R. Couchman, S.C. Moran, A.B. Lockhart, K.J. Swinford,

 and R.G. LaHusen

 Geological observations of lava-dome growth

 8 Use of digital aerophotogrammetry to determine rates of lava dome growth, Mount St. 

Helens, Washington, 2004–2005................................................................................................. 145

 By S.P. Schilling, R.A. Thompson, J.A. Messerich, and E.Y. Iwatsubo 

 9 Growth of the 2004–2006 lava-dome complex at Mount St. Helens, Washington.............. 169

 By J.W. Vallance, D.J. Schneider, and S.P. Schilling

10 Photogeologic maps of the 2004–2005 Mount St. Helens eruption ...................................... 209

 By T.M. Herriott, D.R. Sherrod, J.S. Pallister, and J.W. Vallance

Contents

Chapters

Foreword ...................................................................................................................................................... iii

Anthony I. Qamar, 1943–2005 .................................................................................................................... iv

Contributors to This Professional Paper.................................................................................................. v
ix

11 Remote camera observations of lava dome growth at Mount St. Helens, Washington, 

October 2004 to February 2006.................................................................................................... 225

 By M.P. Poland, D. Dzurisin, R.G. LaHusen, J.J. Major, D. Lapcewich, 

 E.T. Endo, D.J. Gooding, S.P. Schilling, and C.G. Janda

12 Extrusion rate of the Mount St. Helens lava dome estimated from terrestrial imagery, 

November 2004–December 2005 ................................................................................................ 237

 By J.J. Major, C.G. Kingsbury, M.P. Poland, and R.G. LaHusen

13 Effects of lava-dome growth on the Crater Glacier of Mount St. Helens, Washington.... 257

 By J.S. Walder, S.P. Schilling, J.W. Vallance, and R.G. LaHusen

 Geodesy and remote sensing

14 Constraints and conundrums resulting from ground-deformation measurements made 

during the 2004–2005 dome-building eruption of Mount St. Helens, Washington ............. 281

 By D. Dzurisin, M. Lisowski, M.P. Poland, D.R. Sherrod, and R.G. LaHusen.

15 Analysis of GPS-measured deformation associated with the 2004–2006 dome-building 

eruption of Mount St. Helens, Washington............................................................................... 301

 By M. Lisowski, D. Dzurisin, R.P. Denlinger, and E.Y. Iwatsubo

16 Instrumentation in remote and dangerous settings; examples using data from GPS 

“spider” deployments during the 2004–2005 eruption of Mount St. Helens, Washington..335

 By R.G. LaHusen, K.J. Swinford, M. Logan, and M. Lisowski

17 Use of thermal infrared imaging for monitoring renewed dome growth at Mount St. 

Helens, 2004 .....................................................................................................................................347

 By D.J. Schneider, J.W. Vallance, R.L. Wessels, M. Logan, and M.S. Ramsey

18 Radar interferometry observations of surface displacements during pre- and coeruptive 

periods at Mount St. Helens, Washington, 1992–2005............................................................ 361

 By M.P. Poland and Z. Lu

 Models and mechanics of eruptive processes

19 From dome to dust: shallow crystallization and fragmentation of conduit magma during 

the 2004–2006 dome extrusion of Mount St. Helens, Washington ....................................... 387

 By K.V. Cashman, C.R. Thornber, and J.S. Pallister

20 Frictional properties of the Mount St. Helens gouge ............................................................. 415

 By P.L. Moore, N.R. Iverson, and R.M. Iverson

21 Dynamics of seismogenic volcanic extrusion resisted by a solid surface plug, Mount St. 

Helens, 2004–2005........................................................................................................................ 425

 By R.M. Iverson

22 Constraints on the size, overpressure, and volatile content of the Mount St. Helens 

magma system from geodetic and dome-growth measurements during the 2004–2006+ 

eruption .......................................................................................................................................... 461

 By L.G. Mastin, E. Roeloffs, N.M. Beeler, and J.E. Quick
x

 Crisis management

23 Managing public and media response to a reawakening volcano: lessons from the 2004 

eruptive activity of Mount St. Helens......................................................................................... 493

 By P.M. Frenzen and M.T. Matarrese 

24 Hazard information management during the autumn 2004 reawakening

 of Mount St. Helens volcano, Washington............................................................................... 505

 By C.L. Driedger, C.A. Neal, T.H. Knappenberger, D.H. Needham, R.B. Harper, 

 and W.P. Steele

Volcanic emissions

25 Pre- and post-eruptive investigations of gas and water samples from Mount St. Helens, 

Washington, 2002 to 2005 ............................................................................................................. 523

 By D. Bergfeld, W.C. Evans, K.A. McGee, and K.R. Spicer

26 Emission rates of CO2, SO2, and H2S, scrubbing, and preeruption excess volatiles at Mount 

St. Helens, 2004–2005 .................................................................................................................... 543

 By T.M. Gerlach, K.A. McGee, and M.P. Doukas

27 Chlorine degassing during the lava dome-building eruption of 

Mount St. Helens, 2004–2005 ....................................................................................................... 573

 By M. Edmonds, K.A. McGee, and M.P. Doukas

Petrologic and geochemical investigations of eruptive products

28 The Pleistocene eruptive history of Mount St. Helens, Washington, from 300,000 to 12,800 

years before present..................................................................................................................... 593

 By M.A. Clynne, A.T. Calvert, E.W. Wolfe, R.C. Evarts, R.J. Fleck, and M.A. Lanphere

29 Identification and evolution of the juvenile component in 2004–2005 Mount

 St. Helens ash................................................................................................................................. 629

 By M.C. Rowe, C.R. Thornber, and A.J.R. Kent

30 Petrology of the 2004–2006 Mount St. Helens lava dome—implications for magmatic 

plumbing and eruption triggering................................................................................................ 647

 By J.S. Pallister, C.R. Thornber, K.V. Cashman, M.A. Clynne, H.A. Lowers, 

 C.W. Mandeville, I.K. Brownfield, and G.P. Meeker, 

31 Magmatic conditions and processes in the storage zone of the 2004–2006 

Mount St. Helens dacite .............................................................................................................. 703

 By M.J. Rutherford and J.D. Devine, III 

32 Chemistry, mineralogy, and petrology of amphibole in Mount St. Helens

 2004–2006 dacite............................................................................................................................ 727

 By C.R. Thornber, J.S. Pallister, H.A. Lowers, M.C. Rowe, C.W. Mandeville,

 and G.P. Meeker

33 Evolving magma storage conditions beneath Mount St. Helens inferred from chemical 

variations in melt inclusions from the 1980–1986 and 

current (2004–2006) eruptions...................................................................................................... 755

 By J. Blundy, K.V. Cashman, and K. Berlo
xi

34 Plagioclase populations and zoning in dacite of the 2004–2005 Mount St. Helens eruption: 

constraints for magma origin and dynamics............................................................................. 791

 By M.J. Streck, C.A. Broderick, C.R. Thornber, M.A. Clynne, and J.S. Pallister

35 Trace element and Pb isotope composition of plagioclase from dome samples from the 

2004–2005 eruption of Mount St Helens, Washington ............................................................. 809 

 By A.J.R. Kent, M.C. Rowe, C.R. Thornber, and J.S. Pallister

36 238U-230Th-226Ra disequilibria in dacite and plagioclase from the 2004–2005 eruption of Mount 

St. Helens ........................................................................................................................................ 827

 By K.M. Cooper and C.T. Donnelly

37 Timing of degassing and plagioclase growth in lavas erupted from Mount St. Helens, 

2004–2005, from 210Po–210Pb–226Ra disequilibria......................................................................... 847

 By M.K. Reagan, K.M. Cooper, J.S. Pallister, C.R. Thornber, and M. Wortel

In pocket at back:

DVD-ROM containing the entire book, including digital appendixes to various chapters.
xii
Heading 1 1

Overview

Basic facts of the 2004–2006, and continuing, eruption of Mount St. Helens, Washington

Early unrest

Seismic swarm onset September 23, 2004, 0200 hr Pacific daylight time (PDT)

First explosion October 1, 2004, 1202 hr PDT

Lava extruded to surface October 11, 2004 (not observed because of poor weather 

but deduced subsequently)

Explosions

Number of explosions 6

Volcano explosivity index (VEI) 1–2 for early October 2004 vent-forming explosions

Largest VEI 2 (Oct. 5, 2004; Mar. 8, 2005)

Maximum known height of eruption clouds 11 km according to pilot reports

6 km minimum by Doppler radar (NEXRAD), March 8, 2005

Gases

Sulfur dioxide, maximum 240 tons per day (Oct 27, 2004)

Carbon dioxide, maximum 2,415 tons per day (Oct 7, 2004)

Lava dome

Composition Dacite, 65% SiO2

Phenocrysts Plagioclase, amphibole, hypersthene

Total volume extruded About 95×106 m3 (as of September 11, 2007)

Maximum thickness of new dome 450 m

Highest altitude measured for new dome 2,368 m above sea level (spine 5, July 2005)

Volcano’s summit (crater rim highest point) 2,540 m above sea level

Most notable rock avalanche May 30, 2006, starting volume 50,000–100,000 m3

Crater Glacier

Volume before eruption 80×106

 m3

Volume after eruption 74×106

 m3

Maximum thickness of glacier 2004: about 150 m (datum is 1986 glacier surface)

April 2005–July 2007: about 240 m, east arm, 

(deformed upper surface)

Monitoring sites within 10 km of volcano on 

September 22, 2004 (before onset)

Seismometers 8

GPS receivers 1

Fixed transmitting cameras 1 (Gifford Pinchot National Forest)

Borehole tiltmeters 0

Number of monitoring sites installed within 10 km, 

September 2004–January 2008

Seismometers (includes accelerometers) 20 (some destroyed or removed)

GPS receivers 20 (includes 14 installed by PBO *)

Fixed cameras 8

Tiltmeters 9 (includes 4 installed by PBO)

Borehole strainmeters 2 (installed by PBO)

Spiders

Number deployed to crater for GPS or seismic monitoring 31

Number lost to natural causes 9

* Plate Boundary Observatory, a research component of EarthScope, which is sponsored by the National Science Foundation.
2 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 1

Overview of the 2004 to 2006, and Continuing, Eruption of 

Mount St. Helens, Washington

By William E. Scott1, David R. Sherrod1, and Cynthia A. Gardner1

Abstract

Rapid onset of unrest at Mount St. Helens on September 

23, 2004, initiated an uninterrupted lava-dome-building eruption that continues to the time of writing this overview (spring 

2006) for a volume of papers focused on this eruption. About 

three weeks of intense seismic unrest and localized surface 

uplift, punctuated by four brief explosions, constituted a ventclearing phase, during which there was a frenzy of media 

attention and considerable uncertainty regarding the likely 

course of the eruption. The third week exhibited lessened 

seismicity and only minor venting of steam and ash, but rapid 

growth of the uplift, or welt, south of the 1980–86 lava dome 

proceeded as magma continued to push upward. Crystalrich dacite (~65 weight percent SiO2

) lava first appeared at 

the surface on October 11, 2004, beginning the growth of a 

complex lava dome of uniform chemical composition accompanied by persistent but low levels of seismicity, rare explosions, low gas emissions, and frequent rockfalls. Petrologic 

studies suggest that the dome lava is chiefly of 1980s vintage, 

but with an admixed portion of new dacite. Alternatively, it 

may derive from a part of the magma chamber not tapped by 

1980s eruptions. Regardless, detailed investigations of crystal 

chemistry, melt inclusions, and isotopes reveal a complex 

magmatic history.

Largely episodic extrusion between 1980 and 1986 produced a relatively symmetrical lava dome composed of stubby 

lobes. In contrast, continuous extrusion at mean rates of about 

5 m3

/s in autumn 2004 to <1 m3/s in early 2006 has produced 

an east-west ridge of three mounds with total volume about 

equal to that of the old dome. During much of late 2004 to 

summer 2005, a succession of spines, two recumbent and one 

steeply sloping and each mantled by striated gouge, grew to 

nearly 500 m in length in the southeastern sector of the 1980 

crater and later disintegrated into two mounds. Since then, 

growth has been concentrated in the southwestern sector, 

producing a relatively symmetrical mound with steep gougecovered slabs on its east flank. Throughout the eruption, the 

position of the extrusive vent has remained more or less fixed. 

Lack of geodetic evidence for either volume increase or pressure increase in the deep magmatic system since about 1990 

and geodetic modeling that can account for only 20 to 30 percent of the 2004-to-present dome volume puzzles geodesists. 

Better constraints on parameters such as magma-chamber 

volume, crustal properties, and magma compressibility are 

needed to improve the models.

Development of the welt and the new dome bisected 

horseshoe-shaped Crater Glacier, which formerly wrapped 

around three sides of the 1980s dome, and fractured, compressed, and thickened the glacier’s surviving east and west 

arms. Doubling of ice thickness resulted in increased flow rate 

and advance of termini, although rapid infiltration of water 

into the highly porous glacier bed prevented substantial basal 

sliding. Overall, dome growth and disintegration has removed 

surprisingly little ice.

The outcome of the ongoing eruption remains uncertain, 

but Mount St. Helens’ varied eruptive history suggests multiple possibilities. One dynamical model and several petrologic 

investigations regard the current eruption as an extension of 

1980s dome building that may persist continuously or episodically for years to come.

Introduction

A commonly asked question in the Pacific Northwest 

during the past 20 years, “When will Mount St. Helens erupt 

again?,” was answered in late September to early October 

2004, when a typical days-long swarm of small earthquakes 

escalated into intense unrest and eruption (Dzurisin and others, 

2005). Many of the widely acknowledged principles of successful volcano-risk mitigation were reinforced: 1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683
4 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

•	 Understand a volcano’s eruptive history and have an 

up-to-date hazard assessment.

•	 Have adequate monitoring systems installed and a good 

record of background behavior.

•	 Have a strong team of scientists and technicians on site, 

be able to draw on other experienced personnel, and 

have replacement monitoring equipment available.

•	 Have an interagency coordination or response plan in 

place and a working relationship with local emergencyresponse and land-management agencies.

•	 Have a well-coordinated joint information center to deliver 

updates to the media, public officials, and the public.

Fortunately, these conditions were largely met, owing 

to the volcano’s memorable eruption in 1980 (Lipman and 

Mullineaux, 1981), to long-established monitoring systems 

operated by the U.S. Geological Survey (USGS) at its David 

A. Johnston Cascades Volcano Observatory (CVO) and the 

Pacific Northwest Seismic Network (PNSN) at the University 

of Washington, and to long-term cooperation in volcano-hazard and risk-mitigation issues by the Gifford Pinchot National 

Forest (GPNF; in which Mount St. Helens is located), Washington State Military Department–Emergency Management 

Division, and USGS–CVO. Nonetheless, scientific uncertainty, 

strong media interest and scrutiny, and the excitement of a 

volcanic crisis created an intense and, at times, chaotic scene.

As described in the contributions to this volume, more 

than 18 months of continuous activity has given scientists 

from the USGS and numerous academic institutions an 

opportunity to closely study an eruption that has progressed 

much differently from that of 1980–86. Rapid onset of unrest 

and eruption, apparently continuous extrusion of gas-poor, 

mostly crystallized lava into a glacier-filled crater, and month 

after month of monotonous seismicity, among numerous 

other observations and surprises, have led to new insights 

and models of eruptive processes. The eventual duration and 

outcome of the current eruption remain unknown, but scientists look forward to exploiting the rich research environment 

that Mount St. Helens continues to provide. Furthermore, the 

eruption has stimulated the seemingly insatiable interest of the 

public and media in volcanoes and reinvigorated visitation at 

the GPNF’s Mount St. Helens National Volcanic Monument 

(herein, the Monument). The result is a more volcano-savvy 

citizenry in the Pacific Northwest, which is essential for 

maintaining awareness of potential volcano hazards posed by 

the other 12 Cascade volcanic centers in the U.S. and the three 

major Canadian volcanic centers in southern British Columbia.

This overview briefly summarizes the recent eruptive history of Mount St. Helens, including the 1980–86 eruption, the 

1986–2004 period of quiescence punctuated by several notable 

periods of unrest and minor explosions, and the current eruption. Although not addressed in this overview, comparisons of 

the current eruption to historical dome-building eruptions at 

other volcanoes are provided in several papers in this volume 

(Schilling and others, chap. 8; Vallance and others, chap. 9; 

Cashman and others, chap. 19).

Previous Eruptive History

Mount St. Helens (fig. 1) has been by far the most frequently active volcano in the Cascade Range during the past 

few thousand years, producing eruptions of a wide variety of 

types and scales (Hoblitt and others, 1980; Crandell, 1987; 

Scott, 1989; Mullineaux, 1996; Clynne and others, this volume, chap. 28). Recent work shows that an eruptive center 

has existed in the Mount St. Helens area for at least 300,000 

years. Early dacitic lava domes created a broad volcano surrounded by aprons of pyroclastic and volcaniclastic deposits. 

Only during the past few thousand years did the volcano grow 

into the high graceful cone of early 1980. Late Holocene 

cone building, which followed an apparent dormant period of 

about 7,000 years, began with several periods of lava-dome 

and plinian eruptions of dacite, not unlike earlier events in 

the volcano’s history. However, starting about 2,500 years 

ago, substantial amounts of basalt and andesite began to erupt 

between dacitic eruptions. Such lava flows buried large parts 

of a central cluster of dacite domes and flanking fans and 

started cone building in earnest. Eruptions during the 17th 

and early 18th centuries raised the cone about 300 m, with 

emplacement of the summit dacitic lava dome, and added 

fans of debris on all flanks of the volcano. Eruptions during 

the early to middle 19th century did not alter the shape of the 

volcano greatly but provided an opportunity for early settlers 

to witness several small eruptions and to realize that Mount St. 

Helens is an active volcano.

1980–1986 Eruption

Although a detailed understanding of Mount St. Helens’ 

eruptive history and a hazard-zonation map were available 

in spring 1980 (Crandell and Mullineaux, 1978), monitoring 

of the volcano was minimal (one telemetered seismometer 

on the west flank and one surveyed electronic-distancemeasurement line on the east flank). With no recent eruptions having occurred, local officials and land managers had 

no experience with volcanic crises. The onset of seismicity 

in mid-March 1980, which accelerated greatly on March 25, 

and steam explosions starting on March 27 generated intense 

interest by public officials and the media. This interest 

required rapid creation of an emergency coordination center 

to bring together representatives of key agencies. GPNF’s 

experience in fire operations helped greatly in organizing a 

response. Scientists learned numerous lessons, especially the 

need to speak with one voice and to quickly address rumors 

and conflicts (Miller and others, 1981). During April and 

early May 1980, a decrease in seismicity and frequency of 

steam explosions reduced public concern and intensified 

calls to open closed areas. However, the rapidly growing 
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 5

north-flank bulge, driven by intrusion of a cryptodome into 

the volcano, kept scientists wary.

Hopes that a distinctive change in seismicity or rapid 

acceleration of ground deformation would permit a shortterm hazard forecast were dashed in minutes on May 18, 

when the north flank failed in a great debris avalanche 

(Glicken, 1998). This avalanche was the largest subaerial 

landslide on Earth in historical time. The resulting rapid 

decompression of the cryptodome generated explosions that 

produced a lateral blast, a rapidly moving pyroclastic density 

current that leveled ~600 km2 of forest (Hoblitt, 2000) and 

killed most of the 57 victims of the eruption. The debris 

avalanche and blast were of a scale unknown in Mount St. 

Helens’ history and proved milestones in the recognition of 

such events as an important hazard at composite volcanoes 

worldwide (Siebert, 1984). The hours-long eruption of pumice and ash that followed the blast was of expectable scale, 

although many affected communities in eastern WashingColdwater Ridge

Visitor Center

Johnston Ridge

Observatory

Sugar Bowl

dome

Plains of

Abraham

Shoestring notch

1980–86

lava dome

2004–6

lava dome

Crater

Glacier

Spirit Lake

x

North Fork

Toutle River valley

Pumice Plain

x Summit

2540 m

46o15’

46o11’15”

122o15’ 122o7’30”

0 1 2

KILOMETERS

Figure 1. Shaded-relief index map of Mount St. Helens and adjacent areas, showing locations discussed 

in text. Base is chiefly from digital elevation model (DEM) of 2003 lidar survey (Schilling and others, chap. 

8, this volume). Crater features are from DEM made from February 2006 aerial photographs, and northwest 

corner of map is from the National Elevation Dataset.
6 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ton, northern Idaho, and western Montana were unaware of 

potential hazards and poorly prepared to respond to tephra 

fallout and resulting deposits on the order of millimeters to 

as much as 8 cm thick (Warrick and others, 1981).

Five smaller subplinian eruptions during summer 1980 

were preceded by recognizable precursors, as were the 20 

lava-dome-building eruptions that ended in October 1986 

(Malone and others, 1981; Swanson and others, 1983). 

Dome-building eruptions were chiefly several-day episodes 

of extrusion of 1–6×106 m3 of lava at average rates as high 

as 25 m3/s. Three episodes lasted longer, including one of 

largely endogenous growth in 1982–83 that lasted 368 days 

and had an average eruption rate of 0.7 m3/s. By October 

1986, the lava dome stood about 270 m above the 1980 crater 

floor and had a total volume in the range of 77 to 91×106 m3

(as estimated by Swanson and Holcomb, 1990, and Mills, 

1992, respectively).

1986 to 2004

During 1980–86 dome building, seismicity was largely 

confined to shallow depth (<3 km) and related temporally to 

periods of extrusion. Starting in 1987, deeper earthquakes 

became more frequent. Stress-field modeling using focal 

mechanisms of deep earthquakes supported a hypothesis of 

pressurization of the magmatic system, which was thought 

to result from sealing of the shallow conduit (Moran, 1994). 

By late 1989, deep (3–10 km) seismicity was dominant and, 

over the next 22 months, 28 shallow, explosion-like seismic 

signals were recorded, at least 6 of which were accompanied 

by confirmed explosions that ejected blocks and produced ash 

clouds. Gas to drive these events was inferred to come from a 

deep magmatic source, likely from crystallization of magma in 

the conduit (Mastin, 1994).

Several other months-long periods of increased deep seismicity occurred after 1992; the most persistent and energetic 

was in summer 1998. A detectable efflux of CO2 during this 

1998 seismicity (Gerlach and others, this volume, chap. 26) 

suggested that magmatic intrusion was involved. Although 

scientists infer that some intrusion happened between 1987 

and 2004 (Moran and others, this volume, chap. 2), geodetic 

surveys offer no support for a substantial increase in volume 

or pressure in the magmatic system after about 1990 (this volume: Dzurisin and others, chap. 14; Lisowski and others, chap. 

15; Poland and Lu, chap. 18).

Throughout 1986–2004, winter snowfall and avalanche 

snow, alternating with summer rockfall debris from the crater 

walls, was accumulating in the moat between the 1980–86 

lava dome and crater wall. By 1996 the mass was thick enough 

to initiate glacier flow and to form steep advancing snouts at 

a time when most of Earth’s glaciers were shrinking (Schilling and others, 2004). As the volcano reawakened in 2004, 

the glacier, formally named Crater Glacier in 2006, formed a 

collar around three sides of the dome and was gaining mass 

yearly. South of the dome, the glacier’s surface stood as much 

as 150 m above the 1986 crater floor.

During this time of relative quiescence, media attention 

to periods of increased seismicity and to anniversaries of the 

1980 eruption, as well as other Mount St. Helens’ issues, kept 

the volcano in the news and maintained awareness among 

the public that the volcano was active and could erupt again. 

Attention was periodically focused on the volcano as a result 

of policy discussions related to continuing high sediment production in the Toutle River basin (Major and others, 2000), to 

longevity and effectiveness of the sediment-retention structure 

built on the North Fork Toutle River, to the environmental 

impact of a proposed extension of State Route 504 across 

the Pumice Plain north of the crater, and to concerns regarding long-term funding issues at the Monument. Educational 

programs at the Monument and in local school districts also 

helped to maintain awareness of potential hazards.

Chronology of Events: 2004 to Early 2006

Beginning on September 23, 2004, three weeks of intense 

seismicity, punctuated by four brief explosions along with 

localized intense ground deformation, signaled a vent-clearing 

phase. During this phase there was great uncertainty as to how 

the unrest would progress and what style of eruption would 

likely result. The final week of vent clearing was characterized by lessened seismicity, but continued localized intense 

ground deformation. A phase of continuous lava-dome building, which is subdivided further by Vallance and others (this 

volume, chap. 9), began on October 11, 2004, and continues at 

the time of this writing (spring 2006). The highlights of each 

phase are summarized below, taken largely from contributions 

in this volume.

Vent Clearing—September 23 to October 5, 2004

A two-day swarm of tiny (mostly Md <1), shallow 

volcano-tectonic earthquakes beginning early on September 

23, 2004, resembled a swarm in November 2001 (Moran 

and others, this volume, chap. 2) and prompted release of an 

Information Statement by USGS–CVO and PNSN at 1800 

PDT on September 23 (table 1, fig. 2). The statement discussed the swarm’s similarity to previous ones and surmised 

that this swarm might reflect heavier-than-normal precipitation in the preceding four weeks. A rain gage at the mouth 

of the Mount St. Helens crater had recorded an abnormally 

wet last month of summer—more than 30 cm of rainfall. An 

update on the morning of September 24 noted the difference 

between the shallow seismicity of the past day and activity during summer 1998 that consisted of larger and deeper 

events and was considered evidence of deep magmatic intrusion. Both statements considered eruption unlikely without 

significant further precursors.

Beginning on the afternoon of Saturday, September 25, 

earthquakes increased in magnitude (Md ≤2.8), and by the 

following morning a total of ten Md >2.0 earthquakes had 
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 7

1,600

1,200

800

400

0

RSAM, IN COUNTS

09/24 10/02 10/10 10/18 10/26 11/03

DATE IN 2004

I

Start of

earthquake

swarm

Lahar Lahar

Growth of lava dome

1 2 3 2

minor wet dry

CHARACTER OF VOLCANIC DEGASSING

ALERT LEVEL NUMBER AND AVIATION COLOR CODE

Growth

of welt

Explosion

Tremor

EXPLANATION

Table 1. Systems used by U.S. Geological Survey–Cascades Volcano Observatory for event notification at Mount St. 

Helens during 2004–2006 eruption for hazards on the ground and for ash hazard to aircraft in flight.

[Terms in use when eruption began. On October 1, 2006, all U.S. Geological Survey volcano observatories adopted a new notification 

system for volcano hazards that uses some different terms and changes definitions of some levels. See http://volcanoes.usgs.gov/2006/

warnschemes.html or http://pubs.usgs.gov/fs/2006/3139/ for details.]

Figure 2. Real-time seismic amplitude measurement (RSAM) during first 6 weeks of unrest and eruption at Mount St. Helens 

served as proxy for rate of seismicity (modified from Moran and others, this volume, chap. 2). Times of explosions and periods of 

tremor are shown by symbols. Character of degassing from Dzurisin and others (2005). For meaning of numerals and colored bars 

indicating alert levels and aviation color codes, see table 1; I, release of first information statement.

Notification for Ground Hazard

Information Statement Describes short-lived events that may or may not be hazardous or gives commentary 

on status of volcano. May also be issued to provide commentary about notable events 

occurring within any staged alert level during volcanic unrest.

Staged alert levels

1 Notice of Volcanic Unrest First recognition of conditions that could lead to a hazardous event

2 Volcano Advisory Hazardous volcanic event likely but not expected immediately, or ongoing eruption 

with localized hazards1

3 Volcano Alert Hazardous volcanic event underway or expected within a few hours or days

Notification of Ash Hazards to Aircraft

Aviation Color Code

Green Volcano is quiet; no eruption is anticipated

Yellow Volcano is restless; eruption is possible but not known to be imminent

Orange Small explosive eruption(s) either imminent or occurring; ash plume(s) not expected 

to reach 25,000 feet above sea level

Red Major explosive eruption imminent or occurring; large ash plumes expected to reach 

at least 25,000 feet above sea level

1

 The second part of the definition was added during the current eruption.
8 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

been recorded—the most in a 24-hr period since lava-dome 

building ended in 1986. This increase in seismicity, along 

with the appearance of low-frequency and hybrid earthquakes, 

prompted reassessment of the probability of hazardous activity 

(Moran and others, this volume, chap. 2). At 1500 PDT on 

Sunday, September 26, after notifying the Washington State 

Emergency Management Division and GPNF, USGS–CVO 

and PNSN released a Notice of Volcanic Unrest (table 1; fig. 

2), indicating that seismicity had surpassed background level 

and that the volcano was in a state that could evolve toward 

eruption. Greatest concern was for explosions such as those 

between 1989 and 1991 that could shower the crater and upper 

flanks with ballistic fragments and create ash clouds that could 

affect aircraft in flight and downwind communities.

The increase in alert level spurred several actions. The 

GPNF activated their Emergency Coordination Center and 

closed the south-flank climbers’ trail and other trails near the 

crater (Frenzen and Matarrese, this volume, chap. 23). The 

number of daily media inquiries, chiefly to USGS–CVO, 

PNSN, and the Monument, rose rapidly, requiring a considerable effort to be directed toward interviews and briefings 

(Driedger and others, this volume, chap. 24). High rates of 

both seismicity and public interest in earthquake information 

compelled PNSN to develop new visualization tools for tracking seismicity and to establish independent Web sites for the 

public and for sharing information with USGS–CVO and other 

scientific partners (Qamar and others, this volume, chap. 3).

Field crews took advantage of uncommonly clear autumn 

weather to begin installation of additional seismometers and 

global positioning system (GPS) receivers around the volcano. 

Fortunately, Earthscope’s Plate Boundary Observatory had 

been planning installation of a network of continuous GPS 

instruments at Mount St. Helens, and they were able to accelerate their program. Aerial observations of the crater at this 

time showed no obvious changes, but, in retrospect, a series 

of new cracks had appeared in the glacier immediately south 

of the 1980–86 lava dome between September 25 and 26. An 

uplifted area coincident with the cracks was confirmed by 

USGS observers on September 29 (Dzurisin and others, this 

volume, chap. 14).

By September 29 seismicity had intensified to about three 

events per minute with maximum magnitudes of Md 2.4–2.8, 

rising RSAM (real-time seismic amplitude measurement) 

values (Moran and others, this volume, chap 2), and hours-long 

periods during which repetitive earthquakes of similar waveform dominated the records (Thelen and others, this volume, 

chap. 4). The well-known association of such seismicity with 

lava-dome-building eruptions of the past few decades at Mount 

St. Helens and elsewhere raised eruption concerns further and 

prompted issuance of a Volcano Advisory (table 1, fig. 2) midmorning on September 29. The aviation color code (see table 

1) was raised to Orange because of increasing concern that 

explosions could send ash to altitudes where air traffic would 

be affected. The Volcano Advisory signified that processes 

were underway that could lead to hazardous eruptive events, 

but not imminently. Initial airborne missions to detect volcanic 

gases (carbon dioxide, sulfur dioxide, and hydrogen sulfide) 

found little, if any (Gerlach and others, this volume, chap. 26). 

Interpretation of the apparent lack of volcanic gas was tempered by the possibility that such gas had been scrubbed at the 

high water-to-gas mass ratios likely to be present at shallow 

levels beneath the nearly glacier-covered crater floor.

The alert-level rise to Advisory triggered the GPNF 

to activate a local Incident Management Team and to begin 

discussions with other agencies regarding a Joint Operation 

Center, which would be necessary if unrest continued to escalate or eruptions began (Frenzen and Matarrese, this volume, 

chap. 23). Intense media interest, directed chiefly toward 

USGS–CVO and the Monument, was beginning to adversely 

affect operations, so local and State emergency management 

agencies joined with USGS–CVO and the GPNF to begin 

organizing a Joint Information Center (Driedger and others, 

this volume, chap. 24).

Real-time seismic amplitude values increased again on the 

evening of September 29, reflecting an increase in maximum 

earthquake magnitudes (Md

 2.8–3.3) and rate of earthquakes of 

Md >2 to about 1 per minute. The RSAM values fluctuated over 

the next 1.5 days. Flights during this time period again failed 

to detect significant amounts of volcanic gas. However, the 

uplifted and cracked area of the glacier south of the 1980–86 

lava dome, along with a sliver of the dome, was continuing to 

fracture and rise several meters per day, creating what became 

known as the welt (Vallance and others, this volume, chap. 9; 

Dzurisin and others, this volume, chap. 14). Data from a GPS 

receiver on the west side of the 1980s lava dome showed northward movement of several centimeters per day, consistent with 

a shove from a rising mass south of the dome (LaHusen and 

others, this volume, chap. 16).

The inaugural flight at Mount St. Helens of a helicoptermounted forward-looking infrared radiometer (FLIR) midday 

on October 1 was rewarded with close observation of the 

first explosion of the current eruption, which bored an icewalled crater through the western part of the welt (figs. 3A, 

B; Schneider and others, this volume, chap. 17). The FLIR 

recorded a maximum temperature at the base of the explosion column of ~160°C, well below magmatic temperature, 

which suggested that the explosion was driven largely by 

steam. The explosion was also witnessed by several thousand 

cheering visitors at the Monument and became clearly visible from the Portland metropolitan area, 80 km southwest, 

as an ash and vapor cloud rose about 2 km above the crater 

rim and drifted southwestward. Neither this nor other explosions over the next few days had any recognizable precursors (Moran and others, this volume, chap. 6). Earthquakes 

stopped about 1 minute after the explosion started, and when 

the explosion signal also ended abruptly 19 minutes later, 

helicorders were quiet for about 3 hours. The rate of earthquakes rapidly increased again during the evening, and by late 

evening RSAM values exceeded those prior to the explosion, 

suggesting that the explosion had relieved elevated pressure 

in the conduit but that the system was repressurizing quickly 

(Moran and others, this volume, chap. 2).
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 9

For the next two days (October 2–3), seismicity remained 

at a high level of several thousand events per day, punctuated by earthquakes of Md

 3.5–3.9. Midday on October 2, a 

50-minute episode of energetic, relatively broadband seismic 

tremor (Moran and others, this volume, chap. 2), coupled with 

hints of volcanic gas during the previous few days (Gerlach 

and others, this volume, chap. 26) and a continued high rate 

of growth of the welt, prompted USGS–CVO and PNSN to 

issue a Volcano Alert at 1400 PDT (table 1, fig. 2). Because 

of concern over an imminent and potentially hazardous event, 

the GPNF evacuated the Johnston Ridge Observatory (JRO), 

8 km north of the crater, the Washington State Department of 

Transportation closed State Route 504 at the Coldwater Ridge 

Visitor Center, and land and air space was closed within 8 km 

(5 mi) of the crater (Frenzen and Matarrese, this volume, chap. 

23). A key concern was that an explosive magmatic eruption 

could produce pyroclastic flows, swiftly melt large amounts 

of snow and ice surrounding the vent, and generate lahars that 

could sweep into the upper North Fork Toutle River valley. 

The increased possibility of ash clouds reaching high altitudes 

where they could affect air traffic warranted raising the aviation color code to Red.

Memories of the catastrophic events of May 18, 1980, 

preoccupied many media and citizens, so a major publicinformation effort was required to maintain a realistic hazard 

perspective (Driedger and others, this volume, chap. 24). 

An explanation that seemed to reduce anxiety was that the 

1980 eruption had so eviscerated Mount St. Helens, creating a deep crater open to the north, that the volcano could no 

longer support a scenario of intrusion of a cryptodome, large 

flank failure, and catastrophic lateral blast. Consensus among 

scientists at USGS–CVO and PNSN was converging on two 

possible scenarios for volcanic activity over the coming weeks 

to months. These were transmitted to land and emergency 

managers. One focused on resumption of lava-dome growth 

accompanied by explosive events of volcanic explosivity index 

(VEI) 2–3. The other included the possibility of eruptions of 

VEI 4 to low 5, similar to the hours-long plinian eruption on 

the late morning and afternoon of May 18, 1980.

Interest on the part of the media and the public peaked 

during the first few days of October, resulting in high rates of 

inquiry and visits to Web sites (Driedger and others, this volume, chap. 24). A Joint Information Center opened on October 

3, and a Joint Operations Center under unified command 

A

B

C

Figure 3. Crater features seen during first two weeks of the eruption of Mount St. Helens. A, Start of explosion of October 1, 2004, 

looking north across glacier toward 1980–86 lava dome. USGS photograph by J.S. Pallister. B, Looking west about 30 minutes after 

explosion of October 1, 2004, ended. USGS photograph by J.S. Pallister. Note tephra-covered snow and crater (white arrow) bored at 

west end of welt (dashed line), which is formed of uplifted glacier ice and southern part of 1980–86 lava dome. C, Noticeably larger welt 

on October 4, 2004, looking west. Some tephra of October 1 and 4 explosions is drying to light color. USGS photograph by S.P. Schilling.
10 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

was established the following day with representatives from 

key land-management and emergency-management agencies 

(Frenzen and Matarrese, this volume, chap. 23).

Additional explosions late on October 3 and midmorning on October 4 were followed by decreases and rebounds 

in seismicity much less dramatic than those on October 1; the 

last explosion on October 5 initiated a drop in RSAM that has 

never recovered (fig. 2; Moran and others, this volume, chap. 

2). This last explosion, rating a VEI of 2, was the most vigorous, generating an ash plume that reached about 2 km above 

the crater rim. The plume drifted to the north-northeast, lightly 

dusting several communities and the eastern part of Mount 

Rainier National Park, about 100 km away.

The relatively low level of seismicity that followed the 

October 5 explosion, including multiplets of hybrid events 

with similar wave forms (Thelen and others, this volume, 

chap. 4), combined with moderate gas emissions (Gerlach and 

others, this volume, chap. 26) and continued growth of the 

welt (fig. 3C; Dzurisin and others, this volume, chap. 14), suggested that the eruption would probably progress to lava-dome 

extrusion. Under such conditions, the probability of hazardous 

events affecting areas beyond closures was small, except for 

ash clouds that could pose hazards to aircraft. Therefore, at 

0915 PDT on October 6, USGS–CVO and PNSN lowered the 

alert level to Volcano Advisory, where it has remained since 

(table 1, fig. 2). Concurrently, the aviation color code was lowered to Orange. The GPNF reverted to a local Incident Management Team (Frenzen and Matarrese, this volume, chap. 

23), and inquiries to the Joint Information Center began to 

decline, eventually leading to its being disbanded on October 

13 (Driedger and others, this volume, chap. 24). USGS–CVO 

continued to update agencies through daily conference calls.

Vent Clearing—Waiting for Lava—October 5 to 

October 11, 2004

Several days of stormy weather beginning on October 

6 drove visitors and media from the Monument, increased 

RSAM with storm noise, and generated a small lahar from the 

crater—a common occurrence with the onset of autumn rain. 

Brief views through the clouds showed that the welt continued 

to grow upward and outward, reaching more than 100 m above 

the former glacier surface. On October 10 a fixed camera on 

Sugar Bowl dome, at the northeast mouth of the crater, began 

transmitting images that tracked growth of part of the welt 

(Poland and others, this volume, chap. 11). Later analysis of 

digital elevation models (DEMs) showed that by October 11 

the deformed area had a volume of about 10×106

 m3 (Schilling and others, this volume, chap. 8). FLIR measurements 

confirmed that by October 10 an area on the northwest part 

of the welt had reached temperatures >270°C, suggesting that 

the crater floor was being heated and pushed upward by rising 

magma (Schneider and others, this volume, chap. 17). An 

airborne survey on October 7 measured gas-emission rates of 

~2,400 metric tons per day (t/d) CO2

, ~100 t/d SO2, and ~10 

t/d H2S, which, although relatively modest for volcanoes such 

as Mount St. Helens, suggested that a dry pathway had been 

established and scrubbing of SO2 reduced (Gerlach and others, 

this volume, chap. 26). The question on everyone’s mind and 

asked repeatedly by the media was, “Has new lava appeared 

on the surface yet?”

Lava-Dome Growth—October 11, 2004, to 

Spring 2006

On October 11, a FLIR survey revealed a newly extruded 

fin-shaped rock spine (in retrospect designated spine 1) about 

30 m high and 60 m long and having a maximum temperature 

of about 580°C (Schneider and others, this volume, chap. 17); 

this spine is considered to have been the initial appearance of 

new lava (fig. 4) and the start of growth of what we call the 

new lava dome (this volume: Schilling and others, chap. 8; 

Vallance and others, chap. 9; Herriott and others, chap. 10). 

The spine occupied the approximate location of the vent for 

early October explosions and stood >200 m above the crater 

floor. In the following days the spine grew, its base and cracks 

showed temperatures as high as 700°C, and additional areas of 

hot rock appeared to the south of spine 1 (designated spine 2; 

Vallance and others, this volume, chap. 9). The first samples of 

the new lava dome were dredged from spine 1 on October 20 

using a weighted bucket on the end of a 30-m line slung from 

a helicopter. The samples looked like typical Mount St. Helens 

dacite (Pallister and others, this volume, chap. 30). During midOctober, RSAM fluctuated broadly, multiplets of hybrid earthquakes appeared and faded, and seismicity settled into a pattern 

of repetitive small (Md

 <1) earthquakes, dubbed “drumbeats,” 

occurring at a rate of one to two per minute, with larger events 

at longer intervals (Moran and others, this volume, chap. 2).

The potential for unpredictable explosions induced decisions to minimize the exposure of field crews; beginning in 

early October 2004, most new and replacement instruments 

in the crater were deployed by helicopter sling. Singlefrequency GPS receivers and accelerometers (seismometers) 

mounted alone or together on tripods, dubbed “spiders,” 

proved invaluable for maintaining close-in monitoring by 

telemetered seismic instruments (McChesney and others, 

this volume, chap. 7) and for allowing real-time geodetic 

measurements of rates and directions of movement of the 

welt and new dome, as well as movement of the glacier and 

deformation of the old lava dome (LaHusen and others, this 

volume, chap. 16; Dzurisin and others, this volume, chap. 14; 

Walder and others, this volume, chap. 13).

Growth of the new lava dome at Mount St. Helens has 

progressed steadily since mid-October 2004, apparently without pause (this volume: Schilling and others, chap. 8; Vallance 

and others, chap. 9; Herriott and others, chap 10; Poland and 

others, chap. 11; Major and others, chap. 12). In contrast, the 

old (1980–86) lava dome grew in episodic spurts on either 

side of a one-year period of continuous growth (Swanson and 

Holcomb, 1990). Rather than building a single dome-shaped 
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 11

structure similar to the old dome, the new dome grew initially 

as a series of recumbent, smoothly surfaced spines or “whalebacks” that extruded to lengths of almost 500 m. As is typical 

for such spines (Williams, 1932), their surfaces were striated 

and grooved and formed of powdery, crushed rock (gouge) 

and cataclasite that result from the solid extrusion grinding 

against the conduit walls during its last few hundred meters 

of ascent (Cashman and others, this volume, chap. 19; Moore 

and others, this volume, chap. 20). The gouge on early spines 

contained both dacite from the extrusion and more mafic constituents interpreted to have come from the conduit walls, but 

during most of the eruption, gouge has been formed entirely of 

dome dacite (Rowe and others, this volume, chap. 29; Pallister 

and others, this volume, chap. 30).

The dacite building the new dome is similar in chemical 

composition (~65 weight percent SiO2) and mineralogy 

(plagioclase>hornblende≈hypersthene) to the dacite erupted 

on May 18, 1980, and is among the most silica-rich and 

incompatible-element-depleted magmas of the past 500 years 

at Mount St. Helens (Pallister and others, this volume, chap. 

30). The dome lava contains 40–50 percent phenocrysts, and 

the groundmass is largely crystalline for most of its last 500 m 

of ascent (Cashman and others, this volume, chap. 19). In several contributions in this volume, petrologists and geochemists 

explore evidence for the origin of the lava and its recent history. Some of it is thought to represent residual magma from 

the 1980s, but several lines of evidence suggest that a component of a different dacite is mixed with the 1980s dacite—one 

newly arrived in the magma chamber or one from a part of the 

magma chamber that was not tapped in the 1980s (this volume: Pallister and others, chap. 30; Thornber and others, chap. 

32; Blundy and others, chap. 33; Streck and others, chap. 34; 

Kent and others, chap. 35; Cooper and Donnelly, chap. 36; 

Reagan and others, chap. 37). Thin reaction rims on amphibole 

phenocrysts suggest storage for a prolonged period at depths 

of 4–5 km or deeper, followed by rapid ascent through the 

conduit (Rutherford and Devine, this volume, chap. 31).

Late Autumn 2004

The first of the whaleback-shaped spines (spine 3) started 

growing in late October 2004 from a spot just southeast of 

spine 1 and east of spine 2. A GPS spider riding on the spine 

moved at an average rate of about 10 m/d for 8 days (LaHusen 

Deformed

glacier ice

1980–86 of welt

lava dome

Spine 1

Crater Glacier

Crater Glacier

Figure 4. First view (toward southeast) of hot, new lava at surface in crater of Mount St. Helens, October 11, 2004. Spine 1 

is composed of pinkish-gray dacite that is disintegrating and forming light-colored debris fans between welt, old dome, and 

west arm of glacier. USGS photograph by J.J. Major. 
12 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

and others; this volume, chap. 16). By the end of November 

2004, spine 3 was about 475 m long and had reached the 

base of the southeast crater wall; by mid-December, the total 

volume increase represented by the three spines of the new 

dome and welt, using the 1986 crater floor as a base datum, 

was about 30×106 m3 (Schilling and others, this volume, chap. 

8). Whether because of the increasing difficulty of extruding an 

ever-lengthening mass or because of stress imposed by impinging on the crater wall, longitudinal fractures appeared and 

widened and transverse fractures severed the spine from the 

vent in mid- to late December (fig. 5). The breakup of the spine 

was accompanied by sporadic earthquakes of Md ~2.5–3.5, suggesting a causal relation (Moran and others, this volume, chap. 

2). The remaining stump of spine 3 grew to form a new spine 

(spine 4) that pushed most of the old spine 3 aside to the east, 

except for a small amount that was stranded on the west side of 

spine 4 (Vallance and others, this volume, chap. 9; Herriott and 

others, this volume, chap. 10). As spine 3 was pushed eastward, it became increasingly fractured, and little of its original 

smooth, gouge-covered surface remained intact.

January to July 2005

Growth and disintegration of spines continued through 

the first 8 months of 2005. Spine 4 formed a prominent 

whaleback whose growth ended similarly to that of spine 

3, whereas spine 5 grew at a steeper angle and its south end 

trended more southwesterly than did spines 3 and 4 (Vallance 

and others, this volume, chap. 9). Growth periods of spines 

4 and 5 each lasted about 13 to 14 weeks. Extrusion rates 

during the growth of spines 4 and 5 were lower than during 

growth of spine 3 (Schilling and others, this volume, chap. 

8), as were linear rates of motion measured by GPS spiders 

(LaHusen and others, this volume, chap. 16), by analysis of 

repeat photographs from the crater mouth (Major and others, 

this volume, chap. 12), and by tracking of points through 

sequential DEMs (Vallance and others, this volume, chap. 9). 

By midsummer 2005, the top of spine 5 reached the highest altitude attained by the new dome, 2,368 m. At that time 

the top stood only a few meters below the lowest point on 

the crater rim, Shoestring notch, and about 180 m below the 

present summit of Mount St. Helens; the dome then towered about 450 m above the 1986 crater floor. Spine 5 was 

a prominent feature in the view of the crater from JRO in 

midsummer 2005, with the crumbled remains of spine 4 and 

its large raft of intact gouge-covered surface lying off to the 

east (fig. 6A).

August 2005 to Spring 2006

A key change in the pattern of dome growth began in 

late July 2005, as towering spine 5 began to crumble rapidly. 

Rather than follow the previous pattern of a new spine growing southward and bulldozing the remains of older spines 

chiefly eastward, the growth of spine 6 progressed westward. 

The likeliest explanation is that spine 6 faced less resistance 

by plowing into the west arm of the glacier than by continuing to push into the previously extruded dome (Vallance and 

others, this volume, chap. 9). Although spine 6 had some 

areas displaying the smooth, gouge-covered surface typical of 

earlier whalebacks, it formed a more domical mass than had 

previous spines. Photogrammetric techniques estimated a rate 

of movement of several meters per day westward and upward 

(this volume: Major and others, chap. 12; Dzurizin and others, 

chap. 14). As spine 6 moved westward, the west side of spine 

5, which was by then highly fractured, slumped westward into 

a widening sag forming between the two spines.

By mid-October 2005, a new spine was detected in 

FLIR imagery to be rising between spines 5 and 6 (Vallance 

and others, this volume, chap. 9). Spine 7 rode up the east 

side of spine 6, and both continued to move westward (fig. 

B

A

f

a

Spine 1

Spine 4

Spine 3

Crater wall

Figure 5. Views of spine 3. A, On December 16, 2004, spine was 

at its maximum extent just as fractures (f, right of the white vapor 

plume) were beginning to develop that would eventually sever 

the spine from the vent; view to west. Note thin deposit of ash 

formed by density current(s) of ash (a) originating as slough(s) of 

fine-grained gouge carapace. USGS photograph by W.E. Scott. 

B, One week later, fractures had expanded and severed spine 3 

from vent; spine 4 was riding up on and pushing the disintegrating 

remains of spine 3 chiefly southeastward; view to southwest. 

USGS photograph by M.P. Poland.
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 13

6B). By spring 2006, much of spine 6 had been covered by 

spine 7 or by an apron of rockfall debris being shed westward from it. The combined mass of spines 6 and 7 continued to move west-northwestward and slightly downward at 

about 1 m/d during spring 2006 (LaHusen and others, this 

volume, chap. 16). Even though spines have formed and 

crumbled repeatedly throughout the eruption, the vent area 

(the area at which the extrusion breaches the ground surface) 

has remained in the same approximate position of the early 

October 2004 explosion pit and initial spine (this volume: 

Schilling and others, chap. 8; Vallance and others, chap. 9; 

Herriott and others, chap. 10).

Modest rates of gas emission during initial stages of 

dome growth in late 2004 diminished to lower rates throughout most of 2005 and early 2006. Typical daily average rates 

were a few hundred metric tons of CO2 and a few tens of 

metric tons of SO2

 (Gerlach and others, this volume, chap. 26). 

In addition, during late August 2005, a survey from the crater 

rim using open-path Fourier-transform infrared spectroscopy 

yielded the first direct measurements of HCl in the plume at 

Mount St. Helens. The average daily emission rate of HCl 

was about 60 percent that of SO2 (Edmonds and others, this 

volume, chap. 27). Intriguingly, studies of crater-floor and 

flank thermal and cold springs have shown no effect from the 

current eruption on either the content of dissolved magmatic 

volatiles or the temperatures of ground water (Bergfeld and 

others, this volume, chap. 25).

The future of dome growth at Mount St. Helens remains 

uncertain. By the end of 2005, the new dome had a total 

volume of about 73×106 m3, nearly equaling the volume of 

the old lava dome in a time span about one-fifth as long (fig. 

7). The rate of extrusion had declined from about 5 m3/s in 

late 2004 to less than 1 m3/s at the end of 2005. Overall, the 

rate of extrusion appeared to be either slowing exponentially 

or becoming relatively linear and fluctuating between <1 and 

2 m3

/s (this volume: Schilling and others, chap. 8; Mastin 

and others, chap. 22). Will the eruption come slowly to a 

halt over the coming months or years, or will it stabilize at a 

steady long-term rate? Alternatively, will the rate of extrusion become variable and episodic—more reminiscent of 

dome growth during the 1980s? Textural similarity of current 

lava to that of the pre-1980 summit lava dome, which is 

known to have been emplaced over about 150 years during 

the 17th and 18th centuries, suggests that the current eruption may last for many more years (Pallister and others, 

1992; Pallister and others, this volume, chap. 30). Finally, 

appreciating the varied eruptive history of Mount St. Helens 

A

B

3

4 5

3

4

5

7

6

Figure 6. Views south to lava dome in Mount St. Helens’ 

crater from fixed camera point at Johnston Ridge Observatory, 

9 km to the north. Numerals denote spines of new lava dome. 

USGS photographs by E.T. Endo. A, June 24, 2005; actively 

extruding spine 5 is approaching its highest altitude. Note highly 

crevassed east (left) arm of Crater Glacier, which had been 

compressed by eastwardly migrating lava dome. B, December 

5, 2005; spine 5 has disintegrated, and spines 6 and 7 are 

migrating westward (to right). Yellow dashed line marks profile 

of 1980s lava dome. 

0

20

0

40

60

80

500 1,000 1,500 2,000

2004–2005

1980–1986

Total

Dome

DAYS SINCE START OF DOME GROWTH

VOLUME, IN MILLIONS OF CUBIC METERS

Figure 7. Comparison of cumulative extruded lava volumes 

for domes of 1980–86 (Swanson and Holcomb, 1990) and 2004–5 

(Schilling and others, this volume, chap. 8) at Mount St. Helens. 

Total for 1980–86 includes tephra and rockfall debris from 

growing dome.
14 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

(Hoblitt and others, 1980), does the volcano have something 

entirely different in store?

Other Events Related to Dome Growth

Small Explosions Accompany Dome Growth

An explosion early on the stormy morning of January 16, 

2005, was the first of two notable explosions during that year. 

There were no recognizable precursors to the explosion, but 

any subtle seismic precursors would probably have been lost 

in storm noise. Transmission from several crater instruments 

stopped for seconds-long intervals during the explosion’s 

seismic signal, presumably as a result of ash in the air blocking radio signals (Moran and others, this volume, chap. 6). 

Reception from one near-vent seismometer stopped for several 

hours. Several other instruments on and close to the actively 

growing part of the lava dome were destroyed. Poor weather 

precluded field observations for several days, but a reconnaissance flight eventually revealed evidence of the explosion. 

Craters from ballistic blocks pockmarked the glacier surface, 

chiefly east of the new dome, for hundreds of meters. Ash 

covered the crater east and west of the new dome and formed a 

narrow deposit on the outer east flank of the volcano, consistent with strong westerly winds. About 5 mm of ash fell on the 

lower east flank; the extent of ashfall farther east is unknown 

because of a lack of snow-covered surface to preserve the 

ash and the heavy rainfall that accompanied and followed the 

explosion. Effects of the explosion were similar to those of the 

October 2004 events, but no visible vent crater was formed. A 

shallow trough along the northeast margin of spine 4 may have 

marked the vent. The similarity in texture and composition 

of the ash with previously collected samples of that spine’s 

surficial gouge suggested that the explosion vented along the 

margin of the spine, entraining powdery gouge and blocks 

(Rowe and others, this volume, chap. 29).

The second explosion occurred under good viewing 

conditions late on the afternoon of March 8, 2005 (fig. 8A). 

Seismicity had increased slightly for several hours before the 

explosion and had initiated a close watch in the operations 

room at USGS–CVO, but it was not recognized as a precursor to an imminent explosion (Moran and others, this volume, 

chap. 6). The Sugar Bowl camera at the northeast mouth of 

the crater captured images of a dense ash cloud rising from 

near the new dome (Moran and others, this volume, chap. 6, 

fig. 10) and ballistic-impact craters in snow on the north flank 

of the old dome. Ballistic fragments destroyed one seismometer located between the old and new lava domes and one 

seismometer and two microphones located on the old lava 

A

B

1980–86

lava dome

Crater Glacier

C

a

Figure 8. Explosion of March 8, 2005, at Mount St. Helens. A, Ash 

and vapor cloud as viewed toward north-northeast from Cascades 

Volcano Observatory, 70 km away. USGS photograph by M. Logan. 

B, View to southeast showing northwest base of 1980–86 lava 

dome and west arm of Crater Glacier pockmarked by craters 

(as large as 2 m across) formed by rain of ballistic fragments. 

USGS photograph by D. Dzurisin. C, Tephra deposit (a) on 

east-northeast flank of the volcano as seen from the Plains of 

Abraham. USGS photograph by D.R. Sherrod. Note the meager 

snowpack of the winter of 2004–5. D, Map showing extent of 

ballistic fragments and tephra fall from the explosion.
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 15

dome. A white vapor cloud billowed high above the crater 

rim and drifted east-northeastward. Pilots reported the top of 

the cloud reached an altitude of 11 km; the National Weather 

Service’s NEXRAD detected it up to 6 km. The whiteness of 

the upper parts of the cloud suggested that it contained little 

ash. Application of a new one-dimensional steady-state model 

for wet volcanic plumes using that day’s atmospheric conditions indicates that relatively high humidity may have boosted 

the plume height several additional kilometers (Mastin, 2007). 

The explosion continued vigorously for about 10 minutes 

and then waned over several tens of minutes. Investigations 

the following day revealed a field of ballistic craters (fig. 8B) 

extending about 1 km north-northwest of a poorly defined 

and vapor-shrouded possible vent area at the north end of the 

new lava dome. A narrow deposit of coarse ash and fine lapilli 

extended east-northeastward, discernible on snow for about 7 

km from the vent (figs. 8C, 8D). This fallout deposit was about 

20 mm thick on the crater rim and about 2 mm thick on the 

lower east flank of the volcano. Dustings of ash were reported 

in Ellensburg, Yakima, and Toppenish, Washington, as far as 

150 km from Mount St. Helens. Lithic lapilli as large as 4 cm 

fell near the crater rim and fragments as large as 1 cm fell on 

the lower east flank.

Rockfalls and Rock Avalanches

Rockfalls and rock avalanches have both generated 

small pyroclastic density currents, with reaches of <1 km, 

and numerous ash clouds that rose above the crater rim. At 

most, such clouds only weakly dusted the outer flanks of 

the volcano and traveled only a short distance downwind. 

Signals of rockfalls and small rock avalanches were common on seismic records, and such events have been directly 

observed by scientists during hours-long occupations of 

sites on the old lava dome and crater rim, as well as during 

fortuitously timed flights (fig. 9). Many such events have 

also been recorded by fixed cameras (Poland and others, this 

volume, chap. 11). The largest avalanches observed have 

been on the order of tens of thousands of cubic meters, but 

most have been much smaller. Periods of rapid disintegration of spines and accompanying large earthquakes favored 

large avalanches and frequent rockfalls. The high degree of 

crystallinity and low gas content of the lava (this volume: 

Pallister and others, chap. 30; Gerlach and others, chap. 26; 

Edmonds and others, chap. 27) are probably responsible for 

the relatively small amount of ash generated by avalanches 

and rockfalls and the restricted distribution of deposits of 

Figure 8—Continued.

KILOMETERS

0 0.5 1

46°

12.5'

122o10'

D

TRACE

5

TRACE

20

100

1

EXPLANATION

Isopach, showing thickness in millimeters

Limit of ballistic fallout

Ash thickness, in millimeters

Approximate location of vent

1

20

2-3

1

1

<1
16 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

resulting density currents. Sloughing of ashy gouge from the 

outer parts of whaleback-shaped spines, especially spine 3, 

has also created weak density currents of ash and accompanying 

ash clouds (fig. 5A).

Dome Growth Perturbs Glacier

Continued eastward movement of spines 3, 4, and 5 

through late summer 2005 greatly compressed, thickened, 

and fractured the east arm of Crater Glacier (fig. 6A). Such 

conditions created an unprecedented glaciological experiment 

(Walder and others, this volume, chap. 13). The resulting doubling of thickness accelerated surface movement to more than 

1 m/d, which, although high for such a small glacier, is less 

than would be expected from such a dramatic thickening. The 

explanation for lower-than-expected flow rate is that the bed 

of the glacier, which comprises talus and coarse pyroclastic 

debris, is so permeable that water pressure at the bed cannot rise sufficiently to induce rapid basal sliding (Walder and 

others, 2005). In response to this perturbation, the snout of the 

glacier’s east arm thickened noticeably and advanced markedly during winter of 2005.

Beginning in late summer 2005, westward movement of 

spines 6 and 7 started compressing the west arm of Crater Glacier (fig. 10). As a result, a similar thickening and accelerated 

flow rate began to occur on the west arm (Walder and others, 

this volume, chap. 13).

A remarkable aspect of the interaction of the growing lava dome and glacier is the apparent lack of significant 

glacier melting (Walder and others, 2005; this volume, chap. 

13). As the glacier grew in the crater between 1986 and 2004, 

USGS–CVO noted the potential for increasingly larger lahars 

if future explosive eruptions or lava-dome collapses generated 

pyroclastic flows. Such concerns dominated the hazard outlook for the first few weeks of unrest in 2004. Once lava-dome 

growth commenced, the lahar hazard assessment focused on 

the possibility of explosions from the dome swiftly melting 

snow in the crater, as had happened in 1982 and 1984 (Pierson, 1999), as well as on collapses from the new lava dome 

incorporating and melting snow and ice. Neither scenario has 

transpired to date (summer 2006). Explosions in January and 

March 2005 did not generate pyroclastic flows or surges of 

note and apparently melted little snow and ice. Rockfalls and 

rock avalanches from the dome have been of modest size, and 

neither they nor pyroclastic density currents derived from them 

have induced much melting (fig. 9B).

New Insights from the Ongoing 

Eruption

Mount St. Helens’ ongoing extrusion of gas-poor dacitic 

lava has presented an opportunity to develop and improve a 

variety of models. These models can illuminate controls on the 

volcano’s eruptive behavior.

Scientists were tantalized by the persistent “drumbeat” earthquakes, at times amazingly periodic, that have 

accompanied nearly steady lava extrusion since early in the 

eruption (Moran and others, this volume, chap. 2). Were the 

drumbeats an indication of repetitive stick-slip motion along 

the margins of the extruding plug, producing the coating of 

striated fault gouge? Strength tests of gouge samples support this hypothesis and also show evidence for rate-weakening friction (Moore and others, this volume, chap. 20). 

A dynamical model demonstrates that repetitive stick-slip 

events are an almost inevitable consequence of magma influx 

at a near-equilibrium rate occurring in conjunction with rateweakening frictional slip along the margins of an extruding 

solid plug (Iverson and others, 2006; Iverson, this volume, 

Figure 9. Rock avalanche from spine 5 and associated ash 

cloud and small pyroclastic density current of May 12, 2005; 

spine 5 and older spines designated by numerals. A, View from 

east crater rim of density current flowing northwestward and 

ash cloud rising and obscuring spine 5. USGS photograph by M. 

Logan. B, Aerial view from northwest following event, showing 

scar at rockfall source (arrow) and thin ash deposit (a) on snow. 

USGS photograph by D. Dzurisin.

A

4

1

2

5

3

B

West arm of

Crater Glacier

a

East arm

of Crater Glacier
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 17

chap. 21). Because such a condition was attained very early 

in this eruption, the model implies that the magma-plug 

system was probably close to equilibrium at the onset and 

perhaps needed only a small perturbation to be triggered into 

action. It remains unknown whether the trigger was related to 

a small increase in magmatic pressure, to pressure increase 

caused by accumulation of an excess-volatile phase at the 

top of the magma chamber (Kent and others, 2007; Rowe 

and others, this volume, chap. 29), or to a shallow process 

such as weakening of the conduit cap rock through fracturing 

induced by late summer rain and glacier melt (Iverson, this 

volume, chap. 21). In any case, from the perspective of the 

model, the current eruption is a continuation of 1980s dome 

building, in which conditions of the magma-plug system differ little between eruptive and noneruptive periods.

Geodetic modeling of GPS data, primarily the long-term 

record from the receiver at Johnston Ridge Observatory, 9 km 

from the new lava dome, suggests that the source responsible 

for deformation is an ellipsoidal body that extends from about 5 

km to 10–20 km deep, but that the volume lost from the source 

is only 20 to 30 percent of the erupted volume of lava (Lisowski 

and others, this volume, chap. 15). A model that incorporates 

geodetic and dome-growth data explores possible reasons for 

this apparent inconsistency, which include ongoing partial 

recharge, magma compressibility controlled by gas content, and 

size of the Mount St. Helens magma chamber (Mastin and others, this volume, chap. 22).

Summary

The contributions that follow in this volume represent the 

work of more than 100 scientists, emergency managers, and 

information specialists from numerous academic institutions, 

the USGS, the Gifford Pinchot National Forest, and State and 

local government. Together they provide a broad perspective 

of the ongoing eruption of Mount St. Helens and highlight the 

following events, findings, and lessons learned.

Figure 10. View of Mount St. Helens crater looking east-northeastward taken April 4, 2006. Spines 4 to 7 are identified by 

numerals; a steeply east-dipping gouge-covered slab of spine 7 (high point) is riding up over earlier parts of spine 7 and spine 6 

and pushing them toward viewer’s left. At time of photograph, a GPS spider riding to left of numeral 7 was moving about 1 m/d 

west-northwestward (to left) and slightly downward. Movement of spines 6 and 7 westward was compressing and fracturing 

west arm of Crater Glacier, but greater-than-normal snowfall during winter of 2005–6 was filling in crevasses and burying much 

of older part of new lava dome. Shoestring notch, lowest point on the east crater rim, lies just above spine 4.

USGS photograph by J.W. Vallance. 

6

5

7

4

1980–86

lava dome

West arm of

Crater Glacier
18 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The key requirements for successful volcano-crisis 

response were in place at Mount St. Helens when unrest began 

on September 23, 2004. These included geophysical monitoring systems, detailed knowledge of the volcano’s eruptive 

history and hazards, a well-trained and highly experienced 

scientific staff, many public officials and citizens knowledgeable about volcano hazards, and years of close coordination 

among scientists, land managers, and emergency managers in 

planning for possible eruptions.

Once volcanic unrest began, intense interest from both 

media and the public required rapid establishment of a joint 

information center staffed by scientists and representatives of 

land-management, emergency-management, and other relevant 

State and local agencies. The center coordinated dissemination of clear, consistent messages and expanded and shrank as 

demand dictated.

The 2004-to-present eruption progressed rapidly from 

unrest to long-term lava extrusion. Between September 23 and 

October 11, 2004, intense seismic unrest and localized surface 

deformation, punctuated by four short-lived explosions, constituted a vent-clearing phase. During the final week of this 

phase, lessened seismicity, minor venting of steam and ash, 

and rapid localized surface deformation signaled that magma 

was continuing to push upward through crater-floor debris and 

glacier ice. Between October 11, 2004, and the present (spring 

2006), persistent low levels of small, regular drumbeat earthquakes and sporadic larger seismic events, rare explosions, low 

gas emissions, and frequent rockfalls accompanied continuous 

lava-dome extrusion.

Ongoing continuous lava-dome growth has contrasted 

markedly with the largely episodic growth of the 1980s dome 

and has resulted in a strikingly different-looking dome. A 

succession of spines, some recumbent, smoothly gougecoated, and nearly 500 m long, has built a dome about equal 

to the volume of the 1980s dome in about one-quarter of the 

time required to build that earlier dome. Rather than a single 

mound, the new dome currently comprises three main rock 

masses arrayed east-west across the crater between the 1980s 

dome and south crater wall. The vent for extruding lava has 

remained relatively fixed, and successive spines have been 

able to push older masses aside across the glacier-covered 

crater floor.

A dynamical model demonstrates that repetitive stick-slip 

events, such as drumbeat earthquakes might represent, are an 

almost inevitable consequence of magma influx at a nearequilibrium rate occurring in conjunction with rate weakening 

frictional slip along the margins of an extruding solid plug. 

Because such a condition was attained very early in the eruption, the model implies that the magma-plug system was close 

to equilibrium at the onset and perhaps needed only a small 

perturbation to be triggered into action. From this perspective, 

the current eruption is a continuation of 1980s dome building wherein conditions of the magma-plug system differ little 

between eruptive and noneruptive periods.

Crystal-rich dacite with a bulk chemical composition 

similar to that erupted explosively on May 18, 1980, has been 

building the new lava dome. Some of this is thought to be 

residual dacite from the 1980s, but a new component is likely 

admixed. The lava of the current eruption has texture similar 

to that of the pre-1980 summit dome, which was emplaced 

over about 150 years. This suggests that the current eruption 

may last for many more years.

The outcome of the ongoing eruption is uncertain. During 

the past few thousand years, Mount St. Helens has at times 

sustained dome growth for more than a century, as during 

construction of the pre-1980 summit dome; has alternated 

between explosive and effusive eruptions of dacite; and has 

also quickly switched from eruptions of dacite to eruptions of 

lava flows of more mafic composition.

Lava-dome extrusion in a glacier-covered crater has created an unprecedented glaciological experiment. The lava dome 

bisected the formerly horseshoe-shaped Crater Glacier into two 

arms that have been successively squeezed against the crater 

wall and thickened as lava spines bulldozed them outward. The 

arms were approximately doubled in thickness over a period of 

months. Flow rates increased, and both termini are advancing 

vigorously. However, rapid dewatering of glacier beds through 

the highly permeable crater-floor material has discouraged basal 

sliding, and most glacier movement consists of internal flow. 

In addition, extrusion of the lava dome through the glacier has 

resulted in amazingly little melting.

Localized surface deformation and dome growth, 

potential hazards to scientists and instruments in proximal 

areas, and a relatively slow but persistent pace of lava extrusion required development and adaptation of monitoring 

systems to study the eruption. Low-cost, portable alternatives 

to traditional GPS and seismic installations, dubbed “spiders,” 

allowed helicopter deployment and retrieval of telemetered 

instruments in areas on and near the growing lava dome and 

highly crevassed glacier. Ingenuity combined with the skill of 

pilots ensured collection of a suite of lava samples from helicopter-borne dredges. Fixed cameras that telemetered images 

permitted visual observations and repeat images, from which 

rates of movement could be estimated. DEMs made from a 

succession of vertical aerial photographs tracked volumetric 

rates of lava-dome growth and provided detailed bases for a 

variety of investigations. Throughout the eruption, helicoptermounted FLIR surveys, a relatively new technique for USGS 

scientists, provided key information regarding extrusion temperature, vent location, and dome structure.

Update of Recent Activity

Since preparation of this overview and the accompanying volume of reports in 2006, lava-dome growth continued 

at a slowing rate through 2007 and paused in late January 

2008, after more than 39 months of continuous growth. This 

pause has been characterized by a very low level of seismicity, cessation of ground-tilt events that were discernible 

on shallow-borehole tiltmeters installed in late 2005 and 
1. Overview of the 2004 to 2006, and Continuing, Eruption of Mount St. Helens, Washington 19

summer 2006, and barely detectable efflux of sulfur dioxide. 

Analysis of repeat photographs from fixed cameras shows 

no evidence of extrusion, but rather only movement indicative of gravitational settling of spine 7, which is confirmed 

by a GPS spider atop the spine. As a result of these observations, USGS–CVO reduced the alert level to Volcano Advisory/Aviation Color Code Yellow on February 21, 2008, 

to signify that volcanic activity has decreased significantly 

but continues to be closely monitored for possible renewed 

increase. (As noted in table 1, the unified notification 

system adopted by all USGS volcano observatories in 2006 

uses some new terms and redefines others from the system 

in use in 2004. Advisory/Yellow under the new system is 

roughly equivalent to Notice of Volcanic Unrest/Yellow in 

the old system).

From 2006 to early 2008, spine 7 continued to broaden. 

By July 2007, the new lava dome was slightly more than 1 

km long (axis west-southwest to east-southeast) and about 

0.6 km wide. Extruded volume had reached about 93×106

m3, and average extrusion rate had dropped from about 0.5 

m3/s during 2006 to 0.1–0.2 m3/s (S.P. Schilling, oral commun., 2007). Seismicity gradually diminished through 2006–7 

and remained shallow (S.C. Moran, oral commun., 2007), 

the craterward motion of GPS instruments slowed greatly 

and probably ceased by mid-2007 (M. Lisowski, oral commun., 2007), gas emissions have remained at low levels (K.A. 

McGee, oral commun. 2007), and the composition of the 

erupting dacite remains unchanged from that reported in the 

first 2004 samples (Thornber and others, 2008b). No explosions have been detected since March 2005. Rockfalls and 

rock avalanches have continued, but the lack of high, steep 

fins and crags since autumn 2006 has kept their volume small. 

The arms of Crater Glacier continue to advance, the west arm 

most rapidly. As of February 2008, the snouts of the arms had 

touched, thereby enveloping the 1980–86 lava dome. Whether 

this pause in extrusion in early 2008 signals the end of dome 

growth or signals the start of a period of episodic lava extrusion, as happened in the 1980s, will become evident during the 

coming months to years.

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24 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 25

Seismicity of the Eruption

 Seismicity has long been the key monitoring tool for recognizing 

precursors to volcanic eruptions. The unrest at Mount St. Helens in 2004 

was no different, beginning on September 23 as a seismic swarm that 

ultimately produced tens of thousands of earthquakes in the days before lava 

reached the surface on October 11. The sheer number of earthquakes forced 

adoption of new methods of processing and analyzing the data, as well as 

new techniques to ensure the rapid dissemination of seismic information 

over the Internet.

In this eruption, small earthquakes struck a cadence so regular that they 

became known as drumbeats—a phenomenon that persisted longer at Mount 

St. Helens than previously described elsewhere. Models to explain drumbeat 

earthquakes typically invoke stick-slip as the source mechanism. At Mount 

St. Helens, changes in drumbeat character were unrelated to variations in 

magma flux at the conduit; thus, drumbeat size and spacing are more likely 

a function of extrusion mechanics than of extrusion rate.

Also known from other eruptions and seen at Mount St. Helens were 

multiplets—families of earthquakes—so similar in their waveform that they 

must have originated from the same source in the shallow crust and traveled 

along similar paths to the seismometers. Their persistence indicated stationary sources that were stable over long periods of time.

Explosion detection capability was enhanced by the use of microphones 

to record infrasound. The infrasound signals were subtle, so multiple microphones were required at several sites within 1–2 km of the vent. The six 

explosions during this eruption lacked seismic precursors, although seismicity dropped after each in response to depressurization of the conduit. Seismic-amplitude values peaked an hour or two before at least one of the explosions, but such instances were too sporadic to serve as a predictive tool.
26 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

View west across 1980s dome and deformed glacial ice to explosion of 

October 1, 2004, the first eruptive event of renewed activity at Mount St. Helens. 

USGS photo by J.S. Pallister.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 2

Seismicity Associated with Renewed Dome Building at 

Mount St. Helens, 2004–2005

By Seth C. Moran1, Stephen D. Malone2, Anthony I. Qamar2 *, Weston A. Thelen2, Amy K. Wright2, and 

Jacqueline Caplan-Auerbach3

Abstract

The reawakening of Mount St. Helens after 17 years and 

11 months of slumber was heralded by a swarm of shallow 

(depth <2 km) volcano-tectonic earthquakes on September 

23, 2004. After an initial decline on September 25, seismicity 

rapidly intensified; by September 29, Md >2 earthquakes were 

occurring at a rate of ~1 per minute. A gradual transition from 

volcano-tectonic to hybrid and low-frequency events occurred 

along with this intensification, a characteristic of many precursory swarms at Mount St. Helens before dome-building 

eruptions in the 1980s. The first explosion occurred October 

1, 2004, 8.5 days after the first earthquakes, and was followed 

by three other explosions over the next four days. Seismicity 

declined after each explosion and after two energetic noneruptive tremor episodes on October 2 and 3. Following the last 

explosion of this series, on October 5, seismicity declined significantly. Over the next ten days seismicity was dominated by 

several event families; by October 16, spacing between events 

had become so regular that we dubbed the earthquakes “drumbeats.” Through the end of 2005 seismicity was dominated by 

these drumbeats, although occasional larger earthquakes (Md

2.0–3.4) dominated seismic energy release. Over time there 

were significant variations in drumbeat size, spacing, and 

spectra that correlated with changes in the style of extrusion 

at the surface. Changes in drumbeat character did not correspond to variations in magma flux at the conduit, indicating 

that drumbeat size and spacing may be more a function of the 

mechanics of extrusion than of the extrusion rate.

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 Pacific Northwest Seismic Network, Department of Earth and Space Sciences, University of Washington, Box 351310, Seattle, WA 98195

3

 Geology Department, Western Washington University, 516 High Street, 

Bellingham, WA 98225

* Deceased

Introduction

As of the writing of this paper (2006), more than 26 

years of recorded seismic history have accumulated at Mount 

St. Helens since the Pacific Northwest Seismic Network’s 

(PNSN) digital recording system began operation in March 

1980. This 26-year period includes the precursory buildup to 

the May 18, 1980, eruption (Endo and others, 1981; Malone 

and others, 1981), post-May 18, 1980, seismicity associated with the 20 eruptions occurring between 1980 and 1986 

(Malone, 1983; Malone and others, 1983; Swanson and others, 

1985; Swanson and Holcomb, 1990), quiescence between 

1987 and 2004 (Moran, 1994; Musumeci and others, 2002), 

and the onset of the eruption in 2004. This history is made all 

the richer by its being recorded on a relatively dense seismic 

network of short-period vertical-component seismometers in 

place since shortly after the first earthquakes began in late 

March of 1980 (fig. 1). Seismic data recorded by this network 

have been used in many studies, including (1) forecasting 

eruptions and detecting explosions (Endo and others, 1981; 

Malone and others, 1981; Malone and others, 1983; Swanson 

and others, 1983; Endo and others, 1990; Jonientz-Trisler and 

others, 1994); (2) eruption dynamics (Qamar and others, 1983; 

Weaver and others, 1983; Shemeta and Weaver, 1986); (3) 

modeling the Mount St. Helens magmatic system (Weaver and 

others, 1981; Scandone and Malone, 1985; Endo and others, 

1990; Barker and Malone, 1991; Moran, 1994; Musumeci and 

others, 2002); (4) interpreting the structure of the magmatic 

system through tomography (Lees and Crosson, 1989; Lees, 

1992; Moran and others, 1999), earthquake locations (Fehler, 

1985; Weaver and others, 1987; Musumeci and others, 2002), 

and b values (Endo and others, 1981; Wiemer and McNutt, 

1997); (5) determining the source process responsible for 

various type of seismic signals recorded at Mount St. Helens 

(Fehler and Chouet, 1982; Malone, 1983; Fehler, 1985; Norris, 1994); (6) describing the stress field associated with the 

magmatic system (Barker and Malone, 1991; Moran, 1994; 
28 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Giampiccolo and others, 1999; Musumeci and others, 2002); 

and (7) defining repetitive events (or earthquake families), 

including the first published study of repetitive events in a 

volcanic setting (Frémont and Malone, 1987).

As a result of these and other studies, by 2004 we had a 

fairly good working model of the magmatic system beneath 

Mount St. Helens. We also had a very good understanding 

of the character of typical eruptive and noneruptive seismicity. These perspectives proved vital to our ability to correctly 

assess the significance of seismicity when Mount St. Helens 

reawakened in 2004.

The reawakening of Mount St. Helens was heralded by 

a swarm of shallow (depth <2 km) volcano-tectonic (VT) 

earthquakes beginning at roughly 0200 Pacific daylight time 

(PDT) on September 23, 2004. More than one million earthFigure 1. Location of stations in Mount 

St. Helens seismic network. Black triangles 

correspond to Pacific Northwest Seismic 

Network (PNSN) stations in place when 

swarm began on September 23, 2004, and red 

triangles to stations installed after eruption 

began. This network had similar geometry 

and density as the Mount St. Helens network 

of 1983. Inset map shows crater stations 

superimposed on digital elevation model 

from April 19, 2005. Most crater stations were 

added after start of 2004 eruption. Many newly 

added stations operated for short periods; at 

most, six crater stations were operational at 

any one time. See McChesney and others (this 

volume, chap. 7) for station installation details.

quakes were recorded by the end of 2005, with events occurring every 1 to 5 minutes. Of these, roughly 8,000 events were 

located by the PNSN through the end of 2005. Many will 

likely never be located, even with sophisticated crosscorrelation and relative relocation techniques (for example, Thelen 

and others, this volume, chap. 4), because many are small, 

occur in the coda of previous events, and (or) have nonimpulsive arrivals. With such a sizable dataset we are still very 

much in discovery mode, and it is likely that at least some of 

what we present in this paper will be superseded by subsequent research. Our principal purpose is to give an overview 

of our understanding of seismicity that occurred before and 

during the eruption in 2004–5, as well as to put the eruption’s 

seismicity in the broader context of the 26 years of seismic 

history of Mount St. Helens.

JRO

SHW

HSR

JUN

STD

SUG

46°15'

46°00'

122°22'30" 122°15' 122°7'30" 122°00'

46°7'30"

46°22'30"

0 5 10 km

WASHINGTON

AREA OF MAP

SEND

RAFT

BLIS

NED

MIBL

MIDE

AHAB

SEP

YEL

WESG

500 m
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 29

Seismicity During Quiescence, 1987–

2004

Although short-term precursory seismicity began on 

September 23, 2004, in a longer term view the first precursory 

seismicity likely started as early as late 1987 (fig. 2), when 

the first of several years-long swarms of “deep” (>3 km) 

VT events began (Moran, 1994). The depth of these swarms 

contrasted sharply with the exclusively shallow (<3 km) 

precursory swarms associated with 19 of the 20 post-May 18, 

1980, dome-building eruptions (Malone and others, 1983). In 

map view, the deeper events were concentrated to the north 

and south of the 1980–86 lava dome (Moran, 1994; Musumeci and others, 2002), in contrast to the east-west orientation of “deep” VT events that extended to 20 km immediately 

after the May 18, 1980, eruption (Weaver and others, 1981; 

Shemeta and Weaver, 1986; Barker and Malone, 1991; Moran, 

1994). Fault-plane solutions for the post-May 18, 1980, 

“deep” events had tangentially oriented P axes with respect 

to the 1980 crater, in contrast to radially oriented P axes for 

post-1987 “deep” events (Barker and Malone, 1991; Moran, 

1994; Musumeci and others, 2002). The changes in stress-field 

orientation and epicentral distribution are best explained by 

a pressure decrease in the magmatic plumbing system after 

1980 and a pressure increase after 1986 (Barker and Malone, DEPTH, IN KILOMETERS

0

3

6

9

12

15

1980 1984 1988 1992 1996 2000 2004

YEAR

Figure 2. Time versus depth plot showing all earthquakes (circles) with well-constrained locations (gap <135°, nearest station <2 

km, and arrival times on at least seven stations) occurring at Mount St. Helens between 1980 and 2005. Depths are relative to a datum 

corresponding to average altitude of stations in Mount St. Helens network, roughly 1.1 km above sea level. Horizontal-axis tick marks 

indicate start of a year, with gaps of 2 years between tick marks.

1991; Moran, 1994; Musumeci and others, 2002). Moran 

(1994) argued that sizable swarms preceding the last two 

dome-building eruptions in 1986 were an indication of the 

formation of a plug in the conduit to the surface, a plug that 

subsequent batches of magma were unable to penetrate. The 

occurrence of six phreatic explosions in 1989–91 (Mastin, 

1994) and detection of magmatic CO2 in association with 

the last significant months-long swarm of deeper seismicity 

in 1998 (Gerlach and others, this volume, chap. 26) provide 

further evidence that intrusion of magma may have occurred 

between 1987 and 2004.

Most seismicity after 1998 was concentrated at depths 

~2 km below the crater floor, although occasional deeper and 

shallower VT events also occurred (fig. 2). The 2-km-deep 

seismicity band could represent the development of a fracture 

system in the base of the plug inferred by Pallister and others 

(1992) and Moran (1994) to extend from the surface down 

to 2 km. The b values for earthquakes in this depth range 

increased marginally from 0.8–1.0 before 1992 to 1.0–1.4 

after 1992 (fig. 3). The b values for events between 3 and 

10 km showed no such change over the same time period, 

indicating that the shallow b-value change is not a result of a 

systematic bias in the PNSN catalog. The b-value increase, 

if real, could indicate an increase in the number of small 

fractures (Mogi, 1962) between 1991 and 1993 and is the 

type of change that would be expected in association with the 
30 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

fracturing of the base of a plug (Vinciguerra and others, 2001). 

However, b values remained relatively constant in the subsequent 12 years leading up to the 2004 eruption (fig. 3), despite 

the evidence for a possible intrusion of magma in 1998.

The last significant earthquake swarm of the 1987–2004 

quiescent period occurred on November 2–4, 2001. The swarm 

began relatively abruptly at around 1730 Pacific standard time 

(PST) on November 2 and consisted of many small (most 

with coda-duration magnitudes Md <0.0, with the largest Md

1.8) and shallow (<2 km) VT events occurring at a rate of 2–3 

per minute. There were also a number of 10–30-minute-long 

spasmodic bursts of 8–10 VT events per minute. The Cascades 

Volcano Observatory (CVO) issued an Information Statement 

at 1800 PST on November 3, 2001, stating that a swarm was 

underway with no evidence that an eruption was imminent. 

Following the last spasmodic burst of VTs at 0400 PST on 

November 4, the number and size of earthquakes declined 

steadily, reaching background levels by November 5, 2001.

September 23–October 5, 2004: VentClearing Phase

September 23–September 25, 2004

The first short-term signs of unrest came in the form of a 

swarm of VT events starting at roughly 0200 PDT on September 23, 2004 (fig. 4), with the onset of unrest defined as the 

time when hourly event counts reached sustained levels above 

1–2 events per hour. The number of earthquakes increased 

throughout September 23, peaked midday on September 

24, and then declined to a minimum early on September 25, 

roughly 48 hours after the swarm began. By September 24, 

earthquakes had become so numerous that the PNSN began 

locating only a subset of them, and real-time seismic amplitude measurement (or RSAM; Endo and Murray, 1991), 

“webicorder,” and spectrogram plots became the principal 

tools for tracking seismicity (Qamar and others, this volume, 

chap. 3). All located events were shallow (<2 km) and small, 

with the largest a Md

 2.2 and most having magnitudes <1.0. 

The b values for catalog locations were ~1, consistent with 

b values before 2004 (fig. 3). Fault-plane solutions from the 

PNSN catalog show a variety of failure mechanisms and 

stress-field orientations, with no clear dominant fracture pattern (appendix 1). To this point, seismicity closely resembled 

the swarm of November 2–4, 2001. This similarity governed 

initial interpretations by the PNSN and CVO. Although somewhat more energetic than the 2001 swarm, we judged that this 

swarm was within the bounds of noneruptive seismicity seen 

during the previous 18 years at Mount St. Helens. An information statement was released on September 24, stating that 

a seismic swarm was occurring at Mount St. Helens, with no 

indications that an eruption was imminent (Dzurisin and others, 2005; Scott and others, this volume, chap. 1).

1987 1991 1995 1999 2003

0

0.4

0.8

1.2

1.6

b VALUE

YEAR

Figure 3. Plot of b values for shallow (depths <3 km) earthquakes 

occurring between 1987 and 09/23/2004, the last day for which the 

PNSN catalog is complete. Horizontal-axis tick marks indicate start 

of a year, with gaps of 2 years between tick marks. The b values 

are computed using maximum-likelihood method (Bender, 1983) 

and are determined for nonoverlapping groups of 150 events with 

Md 0.4 (magnitude of completeness for entire interval) selected 

with same criteria used for earthquakes shown in figure 2.

September 25–October 1, 2004

Seismicity departed from the 2001 script midday on 

September 25, 2004. Instead of continuing to decline, the 

earthquake rate stabilized and then began to increase (fig. 5). 

Lower-frequency events, including hybrid and low-frequency 

(LF) events (previously identified as type “m” and “l” events 

at Mount St. Helens in the 1980s by Malone, 1983, and 

Malone and others, 1983) also began occurring alongside 

the VT events (fig. 6). We elected to refer to these as lowfrequency, or “LF,” events rather than the commonly used 

long-period, or “LP,” event label, because the definition of 

an LP event includes a specific source mechanism (that is, 

the vibration of fluid- or gas-filled cracks; Lahr and others, 

1994; Chouet, 1996). Experience with low-frequency events 

at Mount St. Helens in the 1980s indicated that other source 

mechanisms might be involved, and so we felt it best to use the 

purely descriptive “LF event” terminology without presuming 

a source mechanism. It should be noted that LF events commonly have extended codas at Mount St. Helens, resulting in 

overly large coda-duration magnitudes relative to amplitudebased magnitudes (Qamar and others, this volume, chap. 3).

The transition from VT to hybrid and LF events took 

place over several days, as illustrated by the earthquake 

spectral amplitude (or “ESAM”) plot in figure 5 (see appendix 

2 for details), with LF events becoming progressively more 

dominant through October 5. Similar transitions in earthquake 

types were seen in precursory swarms preceding many domebuilding eruptions at Mount St. Helens in the 1980s (Malone, 
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 31

09/23

09/24

09/25

09/26

12:00

0:00

12:00

0:00

12:00

0:00

12:00

0:00

TIME, IN MINUTES

0 10 20 30 40 50

Start of swarm

60

09/23 09/27 10/01 10/05 10/09 10/13/2004

PEAK AMPLITUDE, IN COUNTS SPECTRAL FREQUENCY,

 IN HERTZ

0

5

10

15

1,000

0

2,000

3,000

DATE (PACIFIC STANDARD TIME)

4,000

Figure 4. Seismic record from station SEP (~500 m north of the vent; see fig. 1) from 1900 PDT September 22 through 1900 

PDT September 26, 2004, showing start of seismic swarm at 0200 PDT September 23 and buildup over next three days.

Figure 5. Plot of 

earthquake spectral 

amplitudes, or ESAM 

(top), and peak amplitudes 

(bottom) for detected 

events at station HSR (fig. 

1) between September 

23 and October 15, 2004 

(see appendix 2 for 

event detection details). 

Horizontal-axis tick marks 

indicate start of a day, with 

gaps of 2 days between 

tick marks. Note that peak 

amplitudes clip at ~3,400 

counts for station HSR.
32 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0 4 8 12 16 20

TIME, IN SECONDS 

09/26/04 1412 PDT

Hybrid

09/25/04 1848 PDT

Low-frequency (LF)

09/24/04 0555 PDT

Volcano-tectonic (VT)

Figure 6. Waveforms for typical volcano-tectonic (VT), hybrid, 

and low-frequency (LF) events recorded at station HSR (fig. 1) 

during September 23–26, 2004.

1983; Malone and others, 1983), as well as in the first week 

of seismicity prior to the May 18, 1980, eruption (Malone and 

others, 1981; Endo and others, 1981). Those transitions were 

also gradual, taking place over periods of hours to days. The 

transitions before the 1980–86 dome-building eruptions were 

particularly noticeable several hours before eruptions commenced, and they therefore became one of the diagnostic tools 

that allowed for increasingly precise predictions of eruption 

onset times (Malone and others, 1983; Swanson and others, 

1983; Swanson and others, 1985). Such changes in event character were inferred by Malone and others (1983) to represent a 

shallowing of the deformation front (as suggested by slightly 

shallower depths for type “l” events) and to be caused either 

by the fracturing of different rock types or by a change in 

source mechanism. In this context, the appearance of lowerfrequency events on September 25, 2004, increased the likelihood in our minds that the seismic swarm would ultimately 

lead to an eruption.

By the morning of September 26 the earthquake rate had 

further intensified. By midday, maximum event sizes had also 

started increasing (fig. 5). The increase in event rate and event 

size and the change in event character caused CVO to issue a 

Notice of Volcanic Unrest (or Alert Level 1) in the afternoon 

of September 26. Earthquake locations showed no obvious 

temporal changes between September 25 and 26, in large 

part because station density within the crater and the velocity model were insufficient to locate shallow earthquakes 

in a geologically complex medium with a precision of less 

than several hundred meters. Between September 25 and 27, 

however, P-wave arrival-time differences increased between 

station SEP (fig. 1), located on the 1980–86 lava dome essentially on top of most earthquakes, and other stations located 

within 4 km of the vent (fig. 7). Daily average arrival-time 

differences between stations SEP and YEL increased the 

most (0.13 s), with differences at other stations increasing by 

0.07–0.1 s. Arrival-time differences between station pairs not 

including SEP did not change significantly. We reexamined 

P-arrival picks and found no evidence that the changes in 

arrival-time differences were a result of analyst bias (multiple 

PNSN analysts were involved in picking earthquakes; Qamar 

and others, this volume, chap. 3) or decreased P-arrival 

impulsiveness at distant stations (which could result in late 

picks). The observed pattern is best explained by either a 

decrease in seismic velocity near station SEP or a decrease in 

earthquake depths below SEP. Although injection of magma 

and associated fracturing of country rock could result in lower 

velocities, shallowing of hypocenters is the simplest explanation for the increase in P-wave arrival-time differences. This 

shallowing is in part supported by the appearance on September 26 of cracks in the crater glacier (Dzurisin and others, 

2005; Scott and others, this volume, chap. 1; Walder and 

others, this volume, chap. 13).

At roughly 1600 PDT on September 26, a series of 

earthquakes with similar waveforms (an earthquake “family” 

or multiplet; for example, see Frémont and Malone, 1987) 

began, at first mixed with other events with different waveforms (Thelen and others, this volume, chap. 4). By early 

morning of September 27, seismicity was dominated by this 

family, with magnitudes increasing through midday and reaching a maximum of Md

 1.9 by late afternoon. This was the first 

clear instance of repetitive earthquakes, something that was 

seen frequently before dome-building eruptions at Mount St. 

Helens in the 1980s (Frémont and Malone, 1987) and has also 

been observed during other dome-building eruptions, including Usu volcano4

 in Japan (Okada and others, 1981), Augustine Volcano (Power, 1988) and Redoubt Volcano (Power 

and others, 1994) in Alaska, Guagua Pichincha in Ecuador 

(Villagómez, 2000), Soufrière Hills volcano in Montserrat 

(Rowe and others, 2004), and Galeras Volcano in Columbia 

(R. White, oral commun., 2006). Preliminary fault-plane solutions for the larger events in this family have reverse-faulting 

4

 Capitalization of “Volcano” indicates adoption of the word as part of 

the formal geographic name by the host country, as listed in the Geographic 

Names Information System, a database maintained by the U.S. Board on Geographic Names. Noncapitalized “volcano” is applied informally—eds.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 33

-1

-0.5

0

0.5

1

1.5

9/23/04 9/25/04 9/27/04 9/29/04 10/1/04

DATE (PACIFIC DAYLIGHT TIME)

P-WAVE ARRIVAL DIFFERENCES, IN SECONDS

HSR

YEL

SHW

STD

-1.5

mechanisms, in contrast to the mix of mechanisms generated 

during September 23–25 (appendix 1). By 1500 on September 

27, the family had become less dominant, but similar events 

continued through September 28.

The event rate increased relatively rapidly, to 1 event per 

minute, around 1700 PDT on September 27, as smaller events 

began occurring between the larger earthquakes (fig. 8). The 

size of the larger earthquakes also increased, with maximum 

magnitudes reaching Md 2.0–2.4 by midday on September 

28. Most fault-plane solutions were dip-slip with a steeply 

dipping fault plane, in contrast to the reverse-faulting mechanisms dominant during the previous 24 hours (appendix 1). 

At this point, clipping of near-field (<4 km) stations occurred 

frequently. As a result, smaller events were frequently overwhelmed by the codas of the larger events, making it difficult 

to distinguish many individual events. It also became increasingly difficult to manually pick earthquakes and determine 

accurate coda magnitudes, because the coda from one event 

would be obscured by the onset of the next event. At this 

point, RSAM became the principal tool for monitoring seismicity (fig. 9), along with helicorder records from stations 

increasingly distant from Mount St. Helens.

Earthquakes further intensified between 0500 and 0800 

on September 29 (fig. 8), when event rates increased to ~3 

earthquakes per minute and maximum magnitudes rose to Md

2.4–2.8, causing CVO to issue a Volcano Advisory (or Alert 

Level 2) at 1040 PST (Dzurisin and others, 2005). A second 

rise in RSAM levels occurred 12 hours later, between 1700 

and 2100 (fig. 9), when maximum magnitudes increased to Md

2.8–3.3, with Md >2 events occurring at a rate of ~1 per minute 

by the evening of September 29. Most fault-plane solutions 

for these events had normal-faulting mechanisms, although 

there were also occasional events with many, or all, down 

first-motions, consistent with reverse faulting (appendix 1). 

One side effect of the high rate of larger earthquakes was that 

they began to dominate seismic recordings at other Cascade 

volcano seismic networks (in particular at Mount Adams, 

Mount Hood, Mount Rainier, and Three Sisters), as well as at 

regional seismic stations in Washington and Oregon, resulting 

in a temporary reduction in PNSN sensitivity to smaller events 

in a broad area around Mount St. Helens.

Following the second intensification on September 29, 

the rate of larger earthquakes declined somewhat, resulting in 

a decline in RSAM values, which reached a relative minimum at roughly 0500 on September 30 (fig. 9). The rate of 

larger earthquakes increased and decreased two more times on 

September 30, with RSAM values always between the minimum and maximum levels achieved on September 29. The 

increases did not correspond with any known changes in the 

crater, although temporal resolution of visual observations and 

measurements from other geophysical instrumentation were 

insufficient to detect changes on an hourly scale. Fault-plane 

solutions from the PNSN catalog switched from dominantly 

normal-faulting to reverse-faulting solutions for a short time 

between 2300 September 30 and 0500 October 1 (appendix 1); 

then, three hours later at 0800, RSAM levels again increased 

as a result of an increase in event rate (figs. 8, 9). RSAM values peaked at 1100 at levels similar to previous maxima. The 

first phreatic explosion occurred an hour later at 1202 PDT, 

8.5 days after the first precursory seismicity on September 

23. This interval was one day longer than the 7-day interval 

between the first earthquakes and first phreatic explosion in 

March of 1980 (Endo and others, 1981; Malone and others, 

1981). This similarity could indicate that a characteristic time 

interval exists between initiation of precursory activity and initial explosive activity at Mount St. Helens. If this is the case, 

then the time interval presumably is governed by the structure 

of the shallow conduit system and surrounding country rock, 

as well as the composition of the intruding magma.

Aside from the RSAM increase, there were no obvious 

precursors to the October 1 explosion. The explosion began 

aseismically, with no obvious changes in seismicity associated 

with the start of the explosion (Moran and others, this volume, 

chap. 6). Timing constraints for the explosion are excellent, 

because a CVO crew was flying around the vent when the 

explosion began (Schneider and others, this volume, chap. 

17). The most obvious change in seismicity occurred at 1203, 

1 minute after the explosion began, when earthquakes suddenly stopped (figs. 8, 9). This cessation was so dramatic that 

it caused one CVO crew working at seismic station STD (fig. 

1) to look up at the crater when the sound of the earthquakes, 

heard through a scanner tuned to the frequency of a nearby 

seismic station, suddenly stopped. A relatively broadband 

Figure 7. Plot of P-wave arrival-time differences between 

station SEP (located on the 1980–86 lava dome) and four other 

stations including HSR (magenta), YEL (dark blue), SHW (green), 

and STD (red), between September 23 and October 1, 2004. (see 

fig. 1 for station location). P-wave arrival-time differences were 

computed only for high-quality picks, with an estimated picking 

error of ≤0.05 s. Note that the absolute value of differences 

increases on all station pairs between September 25 and 

September 27. This increase was not apparent on other non-SEP 

station pairs and is consistent with shallowing of hypocenters.
34 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

09/27

09/28

09/29

09/30

10/01

10/02

10/03

10/04

10/05

0 30 60 90 120

TIME, IN MINUTES

October 2

tremor

October 3

tremor

October 1

explosion

Wind/traffic

noise

10/06

Figure 8. Seismic record from station JUN (6.5 km southeast of vent; fig. 1) from September 27 through October 5, 2004, showing 

seismicity trends during the vent-clearing phase.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 35

0

200

400

600

800

1,000

1,200

1,400

1,600

09/23 09/25 09/27 09/29 10/01 10/03 10/05 10/07 10/09

DATE (PACIFIC DAYLIGHT TIME)

RSAM COUNTS,

AS 10-MINUTE AVERAGE

0

200

400

600

800

1,000

1,200

1,400

1,600

09/23 10/25 11/26 12/28 01/29 03/02 04/03 05/05 06/06 07/08 08/09 09/10 10/12 11/13 12/15

DATE (PACIFIC DAYLIGHT TIME)

RSAM COUNTS, AS 10-MINUTE AVERAGE

10/02/04 

tremor

10/03/04 

tremor

Explosions

2004 2005

(0.5–3 Hz) tremor signal accompanied the explosion after 

the last earthquake at 1203 (fig. 8), with tremor amplitudes 

roughly doubling between 1216 and 1221. At 1221 the tremor 

and the explosion stopped abruptly.

The October 1 explosion destroyed seismic station SEP, 

the closest station to the earthquakes (fig. 1). The loss of this 

station decreased the quality of PNSN catalog locations, with 

average hypocentral errors for well-constrained earthquakes 

(P waves on at least 7 stations, azimuthal gap <135°, nearest 

station within 3 km) increasing from 0.2 km in the 24 hours 

before October 1 to 0.34 km in the 24 hours after the explosion. This impact illustrates the importance of having seismic 

stations close to the source. Indeed, the loss of SEP was one of 

the primary motivations for the development of a seismic sensor package that could be remotely deployed at locations too 

hazardous for crews to work on the ground (McChesney and 

others, this volume, chap. 7).

October 1–October 5, 2004

Seismicity levels dropped precipitously following the 

1202–1221 explosive event, with the first small earthquake 

not occurring until roughly 3 hours later at 1515. This drop 

indicated that seismicity was occurring in response to elevated 

pressure in the conduit system, pressure that was temporarily reduced by the October 1 explosion. Hybrid earthquakes 

slowly increased in number and size and then rapidly 

increased between 1700 and 2000 on October 1, with RSAM 

levels eventually exceeding those of September 29–October 1. 

As with previous intensifications, RSAM levels flattened and 

then gradually declined on October 2 (fig. 9). At roughly 1217 

PDT on October 2, earthquakes were replaced abruptly by 

relatively broadband seismic tremor with dominant frequencies in the 0.5–3-Hz range (fig. 10). Tremor rapidly intensified 

over a period of 20 minutes, to peak reduced-displacement 

values of 28–43 cm2

, and was energetic enough to be recorded 

on seismic stations more than 240 km away. The energetic 

tremor, in combination with the October 1 explosion and 

continuing energetic earthquake activity, caused CVO to issue 

a Volcano Alert, the highest alert level. No explosions or other 

obvious surficial events (such as increased steaming or small 

ash emissions) were observed in association with the tremor. 

Tremor amplitudes began decreasing shortly after the alert 

level change, and, by 1315, tremor had virtually disappeared. 

As with the October 1 explosion, earthquake rates and sizes 

decreased significantly after tremor ceased, although the drop 

was not as abrupt or as long-lived (figs. 8, 9, 10). Earthquake 

rates and sizes began increasing again around 1400, and by 

1900 RSAM values had returned to levels just below those 

before the October 2 tremor episode (fig. 9). A smaller tremor 

episode (reduced displacement 10–20 cm2

, peak frequencies 

0.5–2.5 Hz) occurred between 0250–0315 PDT on October 

Figure 9. Plot of 10-minute real-time seismic amplitude measurement (RSAM) values from station SHW (3 km west of vent; fig. 1) 

between September 23, 2004, and December 31, 2005. Inset shows SHW RSAM between September 23 and October 10, timing of four 

phreatic explosions, and occurrence of two tremor episodes associated with a vent-clearing phase. Horizontal-axis tick marks in 

inset indicate start of a day, with gaps of 2 days between each tick mark.
36 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

3 (fig. 8), and, as with the October 2 episode, no eruptive 

plume was detected (M. Guffanti, written commun., 2004). In 

contrast to the October 2 episode, earthquakes continued to 

occur during the October 3 tremor episode. Although earthquake rates and RSAM levels dropped after the episode and 

remained low for the next 12 hours, the decrease was not as 

great as those that followed the October 1 explosion and the 

October 2 tremor.

Around 1500 PDT on October 3, earthquake rates again 

began increasing, with RSAM levels peaking at 1650 (fig. 9). 

After a brief drop, earthquake rates increased again, with seismicity dominated by many small-amplitude events. By 2100, 

earthquakes were occurring so close together in time that they 

became difficult to distinguish, forming a spasmodic tremorlike signal (fig. 8). Continuous, small-amplitude tremor was 

likely also occurring, although this signal can only be seen in 

the short gaps between individual earthquakes. Earthquake 

rates and the tremorlike signal decreased substantially around 

2305, following a small steam-and-ash explosion at 2240 

reported by U.S. Forest Service observers at the Coldwater 

Ridge Visitor Center, and RSAM dropped to levels similar to 

those from early in the day on October 3 (fig. 9). Following 

this small explosion, the number of larger (Md

 >2) earthquakes 

decreased, with many smaller hybrid earthquakes occurring 

in the gaps between larger events (fig. 8). Earthquake rates 

and RSAM levels gradually increased again until another 

steam-and-ash explosion at 0943 PDT on October 4, and, as 

before, earthquake rates and RSAM levels briefly dropped 

following the explosion (figs. 8, 9). Shortly afterward, RSAM 

levels increased again, peaked at 1700 on October 4, and then 

gradually declined to a relative low at 2330. As with roughly 

two-thirds of the previous RSAM increases, the peak and 

subsequent decline were not associated with explosions or 

other anomalous surface activity. RSAM levels increased once 

again beginning at 0000 PDT on October 5, reaching a peak at FREQUENCY, IN HERTZ

1200 1210 1220 1230 1240 1250 1300 1310 1320 1330 1340

1

0

2

3

4

5

PACIFIC DAYLIGHT TIME

Figure 10. Tremor episode on October 2, 2004, as recorded 

on vertical component of broadband station JRO (8 km north of 

the vent; fig. 1). Time series (above) and spectrogram (below) 

for same time period. Note that discrete earthquakes stopped 

when tremor started at 1217 PDT and were significantly 

diminished after it ended at 1310.

0600 that was maintained until the final explosion of the ventclearing stage began at 0903. Earthquake sizes and rates began 

declining ~15 minutes after the explosion began and continued to decline over the next hour as the explosion continued 

through 1015. After the explosion, RSAM levels were lower 

than at any time since September 29, with the exception of the 

lull following the October 1 explosion. Although RSAM levels 

increased and decreased several times over subsequent weeks, 

they never again reached levels achieved before the October 5 

explosion (fig. 9).

Between September 29 and October 5, earthquakes were 

too large and occurring too close together in time for more 

than a fraction of all earthquakes to be detected and located 

by the PNSN. Using the average event rates given above, we 

estimate that ~30,000 earthquakes occurred over these seven 

days. The largest located event was a Md 3.9 on October 1, 

and 9 events had Md ≥3.5 between October 1 and October 

4. To estimate the total amount of seismic energy expended 

during the vent-clearing phase, we calibrated RSAM values 

at station FMW (91 km north of Mount St. Helens) using 20 

of the largest earthquakes and then integrated RSAM values from FMW between September 23 and October 5. The 

result was a cumulative magnitude equivalent to a Md

 5.5 

earthquake, an order of magnitude lower than the cumulative 

magnitude of Md 6.5 estimated for the March 28–May 18, 

1980, precursory seismicity.

October 5, 2004–December 31, 2005: 

Seismicity Accompanying Dome 

Building

The steep decline in seismicity following the October 

5 explosion caused CVO to lower its alert level to Volcano 

Advisory (Alert Level 2), where it remained through the end 

of 2005. Although RSAM levels increased again between 

October 6–7 and October 8–9, 2004, as a result of increasing event sizes (largest were Md

 2.5–3.0), the increases were 

much more gradual and didn’t reach September 29–October 5 

levels (fig. 9). Because there had been a number of more rapid 

increases in seismicity between September 29 and October 5 

that were not directly associated with eruptive activity, we did 

not change our alert status at the volcano during these more 

gradual intensifications. We noted during both those periods 

that earthquakes were dominated by a single event family, an 

attribute of Mount St. Helens seismicity that continues at the 

time of this writing (Moran and Malone, 2004; MacCarthy and 

Rowe, 2005; Thelen and others, this volume, chap. 4).

The October 8–9 earthquakes were a particularly spectacular example of an event family (fig. 11). Event magnitudes systematically increased over a ~12-hour period starting 

0200 PDT, then stabilized at Md

 2.5–3.0 for almost 24 hours 

before systematically declining over a ~12-hour period back 

to Md 1–1.5 (fig. 5). We generated hourly event stacks (fig. 
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 37

11) to increase the signal-to-noise ratio of the P-wave arrivals 

and found that first motions were down at all stations, consistent with either very shallow reverse faulting or a deeper, 

nondouble-couple source (such as motion along a curved 

surface or an implosion). The appearance of event families, 

taken together with the overall decline in seismic levels, 

the change in event character to dominantly LF and hybrid 

events, the relatively low gas levels (Gerlach and others, this 

volume, chap. 26), and the observed deformation of the crater 

floor (Dzurisin and others, this volume, chap. 14), led CVO 

and the PNSN to believe that the most likely outcome was a 

new episode of lava dome building.

Seismicity declined significantly following the October 

8–9 earthquake sequence. From October 10 to October 16, 

seismicity consisted of small (Md <1) events occurring at a rate 

of 1–2 per minute, as well as slightly larger events with longer 

between-event intervals. At 0200 on October 16 the number of 

discrete Md

 1–1.5 events started decreasing, and between these 

events many tiny LF earthquakes appeared, spaced roughly 20 

s apart (fig. 12). Event magnitudes of the tiny LF earthquakes 

steadily increased through October 17 before stabilizing at Md

~0.5. When first observing these smaller events we noted a 

strong similarity to short-lived episodes of tiny, closely spaced 

LF events that were observed at Mount St. Helens in three 

Figure 11. Plot of hourly stacks of events from earthquake family 

recorded on station HSR (2.5 km south of vent; fig. 1) between 

October 8 and October 10, 2004.

separate instances associated with the 1983–84 continuous 

dome-building eruption (R. Norris, oral commun., 2006). On 

the basis of this similarity, we inferred that these LF events 

likely represented the extrusion of a new dome.

“Drumbeat” Earthquakes

The appearance of regularly spaced, small earthquakes on 

October 16, 2004, marked the beginning of the next seismic 

phase of the eruption, a phase that continues at the time of 

this writing. Seismicity in this phase was dominated by these 

events, which occurred at such remarkably constant intervals 

that we dubbed them “drumbeat” earthquakes. The drumbeats 

consisted of repetitive LF and hybrid earthquakes and were 

one of the hallmarks of the eruption in 2004–5. Indeed, we 

know of no instances at other volcanoes when drumbeatlike 

events have been observed over such a long period of time. 

The drumbeats accompanied the regular rate of extrusion of 

several dacite spines (in this volume: Schilling and others, 

chap. 8; Vallance and others, chap. 9; Herriott and others, 

chap. 10; Poland and others, chap. 11; Major and others, chap. 

12). Many of the drumbeats had waveforms similar to other 

drumbeats (Thelen and others, this volume, chap. 4), a similarity that can be used to relatively relocate events within individual earthquake families with high precision, as was done in 

the 1980s at Mount St. Helens by Frémont and Malone (1987). 

Thelen and others (this volume, chap. 4) present a thorough 

analysis of a number of well-recorded families. Identifying 

and relocating all earthquake families occurring within the 

million-plus earthquakes of late 2004 and 2005 is a large task 

that will likely take years to complete, well beyond the scope 

of this paper.

Several investigators have proposed models that link 

drumbeats and extrusion rate (Harrington and Brodsky, 2006; 

Iverson and others, 2006; Iverson, this volume, chap. 21; Mastin and others, this volume, chap. 22), with most invoking stick 

slip as the source mechanism. Horton and others (this volume, 

chap. 5) note that spectral peaks from drumbeats recorded 

on broadband stations deployed in 2004–5 are suggestive of 

resonance within a structure such as a crack or conduit, with 

resonance initiated by a stick-slip event (Neuberg and others, 

2006) or by a pressure transient in a crack (Chouet, 1996). 

Research into the drumbeat source mechanism is ongoing, and 

it is not our intent to present a thorough discussion of possible 

drumbeat source mechanisms. Instead we focus on describing 

general trends in drumbeat event spacing, size, and frequency 

content and how changes in these characteristics correlated 

with dome extrusion.

Event Spacing

Figure 13 shows that, to a first approximation, the time 

between drumbeats, or event spacing, changed little between 

October 16, 2004, and April 11, 2005, with daily averages 

ranging between 40 and 80 s. An important exception is that 

0 5 10 15 20 25 30 35 40 45

TIME, IN SECONDS

10/08 02:00

10/09 00:00

10/10 00:00
38 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

TIME, IN MINUTES

OCTOBER 15 OCTOBER 16 OCTOBER 17

5 10 15 20 25 30

0000

1200

0000

1200

0

Figure 12. Seismic record from station HSR (2.5 km south of vent; fig. 1) from 1700 PDT on October 15 through 1800 PDT on 

October 17, 2004, showing onset of small, regularly spaced events (“drumbeats”) occurring between larger events early on 

the morning of October 16. Drumbeat amplitudes increased gradually through the day.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 39

between October 16 and November 12, 2004, events were 

occurring only ~20 s apart and were small enough that many 

went undetected by our detection algorithm (appendix 2). 

After November 12, event spacing increased to ~40 s, and 

visual inspections show that most events were detected. 

Between April 12 and July 8–10, 2005, the average event 

spacing gradually increased to more than 560 s, then briefly 

got shorter; on July 20 it became irregular, with spacing 

progressively increasing. In early August the events had 

become too small to be detected reliably by our algorithm on 

station HSR. Even on stations located within 1 km of the vent, 

earthquakes were occurring more as random earthquakes than 

drumbeats. By early October 2005, they had become so infrequent, emergent, and small that we began questioning whether 

the seismic events we were seeing were earthquakes or other 

surface sources, such as ice quakes and rockfalls. However, 

on October 9, 2005, earthquakes began to increase in size, and 

event spacing became remarkably regular again, with events 

occurring every ~70 s. Initially, event sizes were bimodal 

(fig. 13), with a population of larger-magnitude events (peak 

magnitude Md ~2.5) occurring alongside a population of more 

frequent smaller events. The larger-magnitude events gradually decreased in size until by mid-November 2005 there was 

just a single dominant event size. Event spacing began increasing in mid-December and, by the end of December 2005, had 

again become irregular.

One important qualification to the drumbeat story is 

that many small earthquakes recorded on stations located 

inside the crater in 2004–5 were not drumbeats. This was 

particularly the case from May through July of 2005, when 

the spacing between drumbeats increased significantly (fig. 

13). As drumbeats decreased in size and number starting in 

mid-June, there was a factor of 15 difference between numbers 

of detected events at station MIDE (located ~200 m from the 

vent) and station HSR (2.5 km from the vent) between May 

20 and July 18, 2005 (fig. 14). Some of the events detected at 

MIDE were rockfalls, but most were earthquakes that were too 

small to be recorded on stations >0.5 km from the vent. The 

Figure 13. Plot showing ESAM (top), inter-event spacing (middle), and peak amplitudes (bottom) 

for detected events at station HSR (fig. 1) between September 23, 2004, and December 31, 2005 (see 

appendix 2 for event detection details). Roughly 370,000 events were detected during this time period, 

including occasional noise glitches and other false triggers.

Ch. 2, Moran, FIGURE 13

PEAK AMPLITUDE, 

IN COUNTS

EVENT SPACING, 

IN SECONDS FREQUENCY, IN HERTZ

200

0

400

600

09/23 11/09

2004 2005 2005

12/26 02/11 03/30 05/16 07/02 08/18 10/04 11/26

DATE (PACIFIC DAYLIGHT TIME)

0

1,000

2,000

3,000

800

6

2

10

14
40 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

events recorded on MIDE and other near-field stations lacked 

the regular spacing characteristic of drumbeats, and relative 

amplitudes and arrival-time patterns between MIDE and other 

near-field stations indicate that many were occurring within 

the extruded spines.

Event Size Distribution

There was a complex relation between drumbeat event 

spacing and event size. In general, increases in event spacing 

corresponded to increases in event size and vice versa (fig. 

15). For example, on December 16, 2004, average drumbeat magnitudes were Md

 ~1.5 and events were spaced ~90 s 

apart, whereas on December 23, 2004, average magnitudes 

had decreased to Md ~0.5 and events were spaced ~40 s 

apart (although occasional Md

 1.0–1.5 events still occurred). 

If drumbeats are a result of stick-slip motion along a common fault plane and the steady extrusion rate corresponds 

to a constant slip rate (or stress accumulation rate) along the 

fault plane, then smaller drumbeats must be closer together if 

extrusion rate and seismic energy release are to remain constant. However, for seismic energy to have remained constant 

between December 16 and December 23, the event spacing 

would have had to decrease to ~3 s on December 23, a factor 

of 13 less than the observed ~40 s spacing. Thus there was a 

net reduction in seismic energy release between December 

16 and December 23. Another drop in seismic energy release 

occurred in the evening of December 29, 2004, when drumbeat magnitudes dropped to Md

 0.0–0.5 while event spacing 

remained roughly constant (fig. 13).

05/20 06/01 06/13 06/25 07/07

500

1,000

1,500

2,000

500

1,000

1,500

2,000

PEAK AMPLITUDE, IN COUNTS

DATE

PEAK AMPLITUDE, IN COUNTS

HSR - 16,104 events

MIDE - 234,646 events

Figure 14. Detected events vs. time at stations MIDE (top) and HSR between May 20 and July 

18, 2005. Note that the decline in events at HSR starting around June 13 is not seen at MIDE, 

evidence that many events detected on MIDE were occurring close to the station and were 

due to a process distinct from that responsible for the drumbeats.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 41

11/16/04 - 12/31/04 01/01/05 - 02/16/05

INTER-EVENT TIME, IN SECONDS

0 40 80 120 160 200 0 40 80 120 160 200

0

400

800

1,200

0

400

800

1,200

02/17/05 - 03/31/05 04/01/05 - 05/15/05

PEAK AMPLITUDE, IN COUNTS PEAK AMPLITUDE, IN COUNTS

INTER-EVENT TIME, IN SECONDS

These and other decreases in seismic energy release could 

have been caused by decreases in extrusion rate, decreases in 

seismic coupling, or other changes in the mechanics of extrusion. Both December 2004 seismic energy drops (Dec. 23 and 

Dec. 29) occurred during the breakup of spine 3 and transition 

to spine 4 as the active locus of extrusion, a transition that 

began November 16 and continued through mid-January 2005 

(Vallance and others, this volume, chap. 9). A mid-April 2005 

drop in energy release, when event rates declined noticeably 

(fig. 13), coincided with the breakup of spine 4 and transition 

to spine 5, and another energy drop in mid-July coincided with 

the breakup of spine 5 and transition to spine 6 (Vallance and 

others, this volume, chap. 9). The correlation between transitions in extrusion style at the surface and changes in seismic 

energy release suggests that drumbeat seismicity was controlled to a certain extent by the mechanics of extrusion at the 

surface. Support for this interpretation comes from the observation that extrusion rates determined from digital elevation 

models (DEM) computed periodically throughout the eruption 

(Schilling and others, this volume, chap. 8) remained relatively 

constant despite fairly substantial changes in seismic energy 

release. In particular, the extrusion of spine 6 was associated 

with a significant drop in event rate and size from mid-August 

through early October of 2005, as described above. However, 

three DEMs spanning August–October 2005 show that the 

rate of extrusion remained roughly constant during the lull 

in seismicity. This lack of correspondence indicates that the 

rate and size of drumbeats may have been more a function of 

changes in extrusion mechanics than changes in magma flux 

through the conduit. Changes in extrusion mechanics that 

could have influenced earthquake occurrence include changes 

in the angle of extrusion (which would affect clamping 

forces), changes in the locus of extrusion, and (or) changes in 

the frictional properties along the conduit margin through rate 

hardening/softening or other processes.

A further complexity in the relation between event 

sizes and event spacing is that, although there clearly were 

extended periods of time when larger drumbeat events were 

followed by longer gaps, there also were many instances 

when larger events were followed by short gaps and smaller 

events followed by larger gaps. This complexity is not 

predicted by models that marry drumbeats with the regular extrusion rate (Iverson and others, 2006; Iverson, this 

volume, chap. 21; Mastin and others, this volume, chap. 22). 

If the drumbeats were a result of a stick-slip source, then 

this complexity indicates that drumbeats at times occurred at 

multiple locations within a complex network of shear zones, 

with individual faults having individual loading and unloading histories somewhat independent of those on other faults. 

It remains to be seen if individual drumbeat families (Thelen 

and others, this volume, chap. 4) show a more direct relation 

between event size and event spacing.

A final general characteristic of drumbeat magnitudes 

was their magnitude-frequency distribution over time. Most 

commonly, drumbeats occurred with a single dominant event 

size and, in general, did not follow a Gutenberg-Richter relaFigure 15. Plot of peak 

amplitudes vs. event spacing 

for detected events at HSR 

(fig. 1) for four time periods 

(all four plots have same 

vertical and horizontal scales). 

Although a general relation is 

seen between event size and 

inter-event time, particularly 

through mid-February (top two 

plots), the slope decreases 

over time, indicating that 

seismic energy release was 

declining as event spacing 

was increasing. Two linear 

features in the top two plots 

correspond to a sufficiently 

strong secondary phase arriving 

~20 s after the P-wave at HSR 

to be occasionally detected as 

an event.
42 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

tion. This is consistent with the repetitive nature of the drumbeats and implies that source dimensions, as well as location, 

varied little between individual drumbeats. However, there 

were also time periods when event amplitudes appeared to 

follow a Gutenberg-Richter relation (fig. 13), in the sense that 

a greater number of small events occurred than large events. 

We note that there is no relation between event size and event 

spacing during many of these time periods, which also implies 

that multiple seismic sources were active at the same time.

P-Wave Spectra

With a thousand or more events occurring per day, it 

was difficult to notice changes in event character over time, 

although obvious variations from the norm, such as the 

sporadic occurrence of VT events, were noted, usually within 

hours. One means we used for tracking changes in drumbeat 

character was to plot event spectra over time (see appendix 2 

for details). Earthquake spectral amplitude (ESAM) plots in 

figures 5 and 13 show that seismicity transitioned from hybrid 

events following the October 5 explosion to LF events by 

October 9 and was dominated by LF events (peak frequencies 

in the 2–3-Hz range) through early January 2005. Between 

January 8 and 25, hybrid events occurred along with LF events 

(fig. 16), with peak hybrid P-wave frequencies of 8–16 Hz on 

station HSR. Hybrids also occurred with LF events between 

mid-April and late May 2005 and were the dominant event 

type between July 20 and August 11, 2005 (fig. 16). In both 

cases, peak P-wave frequencies were in the 12–16-Hz range 

on station HSR. Hybrids again occurred alongside LF events 

between October 9 and November 18, with peak P-wave 

frequencies in the 4–6-Hz range and a weaker band between 

5 s 5 s 5 s

01/19/05 0116

01/03/05 1516 05/10/05 0652

05/10/05 0648 07/17/05 0356

07/20/05 0754

14 and 16 Hz (fig. 13). These frequency changes are most 

notable on stations HSR and SHW, two stations located 2.5–3 

km from the vent. Changes in event appearance are also apparent on station YEL, ~1 km from the vent (the closest station 

to have operated throughout the entire eruption), but peak 

frequencies for the hybrid events are 4–5 Hz, suggesting that 

the path to YEL attenuates much of the high-frequency energy 

apparent on flank stations.

The four instances of concentrated hybrid occurrence all 

coincide with transitions in spine extrusion, as well as with 

several changes in event amplitude and (or) event spacing 

described above. The best example of this was July–August 

2005, when seismicity switched from dominantly LF to dominantly hybrid events. A time-lapse sequence of photos (fig. 17) 

from a remote camera located on the northeast edge of the crater (Poland and others, this volume, chap. 11; Major and others, this volume, chap. 12) shows that the hybrid events began 

as the south end of spine 5 stopped moving on July 20, resulting in differential motion within the spine. As time progressed, 

the resultant shear zone moved northward, until finally spine 6 

split away from spine 5 and began moving westwards around 

August 9–10. Other major spine transitions occurred December 2004–January 2005, mid-April 2005, and early October 

2005 (Vallance and others, this volume, chap. 9). Hybrid seismic events occurred to varying degrees during each of these 

transitions, a correlation that suggests a direct relation between 

hybrids and motion of the spines after extrusion from the vent. 

One speculative model is that the hybrids are a result of differential shearing occurring along the base of the spine during 

spine transitions, whereas the LF events, which do not have 

as clear a relation to changes in spine motion after extrusion, 

occur at shallow depths along the conduit margin.

Figure 16. Examples of hybrid (top) and low-frequency drumbeats (bottom) as recorded at station HSR in January, May, and July 2005. 

Hybrid drumbeats were relatively rare and occurred in temporal clusters that are clearly visible on the ESAM plot in figure 13. Such 

clusters corresponded to transitional periods when one spine became inactive and a new spine began to grow.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 43

July 18, 2005 July 24, 2005

August 2, 2005 August 13, 2005

Figure 17. Four pictures taken from a camera station at SUG (fig. 1) showing transition from spine 5 to spine 6 that 

took place in late July–early August. Arrows indicate relative motion of common spots on spine 5. At bottom of each 

picture is an ESAM plot showing earthquake spectra between July 1 and August 31, 2005, with frequency ranging from 

0 to 16 Hz (identical scale to figs. 5 and 13). Black vertical bar on each ESAM plot corresponds to date on which that 

picture was taken.
44 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

“Big” Earthquakes

Another intriguing aspect of seismicity after October 

16, 2004, was the occasional occurrence of larger, or “big” 

(Md >2.0, Mmax = 3.4), shallow earthquakes. The first of these 

was a Md 2.4 earthquake at 2048 PST on November 15, 2004, 

that was followed by 11 others over the next two weeks (fig. 

18). These earthquakes were notable because they were 

much larger than the regular drumbeat events (fig. 19). They 

generated much interest among scientists and with the public, 

because they could be seen easily in the real-time digital 

helicorder plots displayed on the PNSN Web site (Qamar and 

others, this volume, chap. 3). Also, they often caused large 

rockfalls that, in turn, generated ash plumes, which frequently 

rose above the crater rim. The onset of the “big” events coincided with spine 3 encountering the south crater wall, as demonstrated by DEMs (Schilling and others, this volume, chap. 

8). Spine 3 was likely in a state of uniaxial compression at this 

point, and the coincidence of its south-wall impingement and 

the occurrence of larger events suggests a causal relation.

The larger events cannot be distinguished from larger 

drumbeat events on the basis of PNSN catalog magnitudes 

alone, because drumbeats tend to have long codas that yield 

overly large Md

 values relative to magnitudes determined 

using amplitudes. We therefore arbitrarily define larger events 

as those post-October 2004 events with amplitudes large 

enough to clip on station HSR. Through the end of 2005 

there were more than 290 larger earthquakes with a cumulative magnitude of Md

 4.4. Most events occurred in one of four 

temporal clusters: November 15, 2004–January 14, 2005; 

January 25–March 9, 2005; March 22–April 22, 2005; and 

July 15–December 31, 2005 (fig. 18). The first and third 

clusters occurred in association with the breakup of spines 3 

and 4, respectively, and the fourth occurred in association with 

the breakup of spines 5 and 6. In contrast, the second cluster 

occurred during steady extrusion of spine 4. The ends of the 

first and second clusters coincided with explosions on January 

18, 2005, and March 8, 2005, respectively. The largest number 

of events in a single day was 11 on March 27, 2005, when 

several events occurred within an hour of their predecessors, 

and the largest daily moment release occurred April 4, 2005. 

Both peaks occurred during the transition in locus of active 

extrusion from spine 4 to spine 5.

Despite their size, the larger earthquakes are similar in 

several ways to the drumbeat events. Like the drumbeats, 

they are classified as either LF or hybrid events. Roughly 40 

percent have impulsive arrivals on most stations, and, like the 

drumbeats, many of the events with impulsive arrivals have 

most or all down first motions (appendix 1). Many of the 

impulsive larger earthquakes have higher frequencies at the 

Figure 18. Plot showing occurrence of “big” earthquakes between November 1, 2004, and December 31, 2005. Green triangles 

correspond to impulsive events, hollow magenta squares to emergent events. Timing of two explosions on January 16 and March 8, 

2005, are indicated by dashed blue lines.

11/01/04 12/13/04 1/24/05 3/07/05 4/18/05 5/30/05 7/11/05 8/22/05 10/03/05 11/14/05 12/26/05

DATE (PACIFIC DAYLIGHT TIME)

MAGNITUDE

1

2

1.5

3

2.5

3.5

4

EXPLANATION

Emergent

Impulsive

ExplosionExplosion
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 45

TIME, IN MINUTES

0 5 10 15 20 25

1200

0000

1200

0000

30

NOVEMBER 30 NOVEMBER 29 

start of the event and are classified as hybrid events, whereas 

the emergent events are dominantly LF events. Clipping of 

near-field sites makes event-family detection difficult, but 

on-scale recordings at broadband station JRO (~9 km from 

the vent; fig. 1) show that, although many of the larger events 

are dissimilar, there was at least one set of events occurring 

between December 2004 and April 2005 with similar waveforms (fig. 20). The similarity indicates not only that at least 

some of the larger earthquakes occurred in the same place 

with the same mechanism, but that the structure on which they 

occurred remained intact over a 4-month time span.

Spectra for drumbeats and the larger events are also 

similar. Figure 21 shows a set of waveforms and instrumentcorrected spectra for a Md

 1.0 drumbeat and a Md 3.0 “big” 

earthquake as recorded on stations plotted in order of distance 

from the earthquake. Spectra on all stations for both events 

have relatively flat curves for frequencies <1 Hz, a welldefined corner frequency at 1–2 Hz, and a 2

 decrease in 

frequency content at higher frequencies, all characteristics of 

tectonic earthquake spectra (Brune, 1970). Horton and others 

(this volume, chap. 5) found similar characteristics for events 

recorded on their broadband network in 2004–5. Of particular 

note is the nearly identical corner frequency (Fc in fig. 21) 

of 1–2 Hz for all spectra, a frequency that is independent of 

distance and magnitude. One implication of this observation is 

that if the source is stick slip in nature, then the source duration is independent of source area. A source area independent 

of event size indicates that event magnitude is primarily a 

function of source displacement, rather than a combination of 

fault area and displacement as it is for tectonic earthquakes. 

This implies that the principal mechanical difference between 

drumbeats and “big” events is the amount of slip. If the earthquakes reflect stick-slip faulting within or at the periphery of 

the rising, rapidly solidifying magma plug (Iverson and others, 

2006; Thelen and others, this volume, chap. 4; Iverson, this 

volume, chap. 21), then it is not surprising that the source area 

for these events is limited, because ruptures cannot propagate 

far outside of high-strain regions. The source duration of such 

a stick-slip event would then be a combination of the duration 

of motion of any part of the rupture surface and the area over 

which the rupture spreads, with the duration of this spreading 

being a function of the rupture velocity.

Locations for the impulsive larger events are more concentrated than drumbeat locations (fig. 22), although epicenters are still spread over an area 1 km by 1 km centered on the 

vent. Some of this scatter is likely due to changes in station 

Figure 19. Seismic record from station HSR (see fig. 1) for 48 hours between 1700 PST on November 28 and 1700 PST on November 30, 

2004, showing ten “big” earthquakes occurring along with regular drumbeats. Note that no change in drumbeat spacing or size occurs 

before or after the larger earthquakes, indicating that source processes are separate and distinct from one another.
46 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

TIME, IN SECONDS

0 1 2 3 4 5

(VARIED SCALES)

04/09/05 0801

12/18/04 1241

04/05/05 0416

12/19/04 1247

AMPLITUDE

SEP

0.5 km

NED

0.5 km

YEL

1.4 km

SHW

3.4 km

STD

5.3 km

ML FC = 1.2 Hz =1.2

FREQUENCY, IN HERTZ FREQUENCY, IN HERTZ

0.1 1.0 10

TIME, IN SECONDS TIME, IN SECONDS

-5 0 5 10

 VELOCITY, IN MILLIMETERS PER SECOND VELOCITY, IN MILLIMETERS PER SECOND

-0.2

0.2

-0.1

0.1

0.05

-0.05

0.01

-0.01

0.002

-0.002

10-4

10-3

10-4

10-3

10-5

10-4

10-3

10-2

10-3

10-2

AMPLITUDE SPECTRA, IN MILLIMETER SECONDSAMPLITUDE SPECTRA, IN MILLIMETER SECONDS

ML=3.0 FC = 1.1 Hz

0.1 1.0 10

-0.1

0.1

1.0

-1.0

0.05

-0.05

0.002

-0.002

0.002

-0.002

0 10 3020

10-4

10-3

10-4

10-3

10-3

10-2

10-3

10-2

10-2

10-1

SEP

0.5 km

STD

5.3 km

JRO

9 km

ASR

46 km

LON

68 km

A B

Figure 20. P-wave arrivals at broadband station JRO (fig. 1) for four “big” events occurring in 

December 2004 (upper two records) and April 2005 (lower two) with similar waveforms. Waveforms 

are scaled relative to the peak amplitude for each event. Although many “big” events do not have 

similar waveforms, the fact that some do indicates that these events on occasion occur at the same 

location with the same source mechanism. The 4-month time span between these four events also 

indicates that the structure on which these events occurred was not destroyed over that interval. 

Figure 21. Instrument-corrected seismogram and frequency spectra for typical drumbeat (A) and “big” (B) seismic events. 

Stations are arranged in increasing distance from vent, with distance indicated for each station. All stations shown on location 

map (fig. 1) except ASR, located near Mount Adams, and LON, near Mount Rainier. Spectra are for 10.24-s windows starting with 

P-wave arrival. Thin dotted line below each spectral plot is noise estimate based on 10.24-s window taken before first arrival. 

Vertical thick dashed lines indicate corner frequencies (Fc

) for each event. A, Drumbeat event, ML 1.2, on August 20, 2005, at 1345 

PDT. B, “Big” event, ML

 3.0, on March 26, 2005, at 1928 PST.
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 47

geometries and inaccuracies in the shallow velocity model 

(inaccuracies that, because of constantly changing velocity 

structure caused by the ongoing eruption, varied through time). 

Some scatter is also due to picking uncertainties, because 

arrival times, particularly on close-in stations, are difficult to 

pick to within ~0.1-s accuracy for all but the most impulsive 

events. However, some of the location scatter is real, the best 

evidence for which comes from differences in relative P-wave 

arrival times at stations placed close to the vent. As an example, 

figure 23 shows records on four crater stations from two events 

recorded ~9 hours apart on February 10 and 11, 2005. For the 

February 10 event, the P-wave arrived first at station BLIS (250 

m east of the vent) and ~0.4 s later at AHAB (located 250 m 

southeast of the vent on top of spine 4); whereas for the February 11 event, the P wave arrived only 0.02 s later at AHAB. 

There was little change in relative P-wave arrivals at stations 

SEP and NED (~500 m north-northwest and ~500 m northnortheast of the vent, respectively). These patterns indicate that 

the February 11 event was located midway between AHAB and 

BLIS (350 m apart, including the altitude difference), whereas 

the February 10 event was located close to BLIS, a difference of as much as ~175 m (depending on P-wave velocities) 

between the two events. The 0.4-s difference in arrivals for the 

first event also implies remarkably low P-wave velocities (<1–2 

km/s) at shallow depths, which is likely a result of AHAB and 

BLIS being placed on top of hot and highly fractured rock and 

crater-filling deposits (rockfall debris and ashfall deposits from 

the 1980 eruptions).

Hypotheses for the larger earthquake source include 

implosive-type mechanisms, fracturing of new rock (formation of new pathways to the surface), fracturing within spines 

before or after extrusion, and lurching of the active spine in a 

large-scale stick-slip event. Implosive mechanisms are consistent with all-down first-motion patterns observed for many of 

the impulsive events. However, none of these events occurred in 

association with explosions or significant degassing events. In 

addition, the extremely low levels of magmatic gases emanating 

from the vent (Gerlach and others, this volume, chap. 26) make 

it unlikely that sizable gas- or steam-filled cracks or pockets are 

able to accumulate below the surface and collapse to produce 

these events. We note, however, that Waite and others (2008) 

demonstrated that several events occurring in 2005 had a sigFigure 22. Map showing epicenters of drumbeat events (blue dots) and 

impulsive “big” events (red dots) occurring between November 1, 2004, 

and December 31, 2005. Dot sizes vary with magnitude. Drumbeat locations 

include events with P arrivals on at least seven stations, a nearest station 

within 2 km, and an azimuthal gap ≤135° (all impulsive “big” event locations 

meet these same criteria). Yellow triangles correspond to seismic station 

locations as of July 2005.

46°14.5'

122°14.5' 122°12' 122°10.5' 122°8'

46°12'

46°10.5' 0 1 km
48 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

February 10, 2005 19:46:33.5, Md 2.4

TIME, IN MINUTES AND SECONDS TIME, IN MINUTES AND SECONDS

BLIS

AHAB

SEP

NED

46:34 46:36 04:35 04:37 04:39

BLIS

AHAB

SEP

NED

February 11, 2005 05:04:34.6, Md 2.2

Figure 23. P-wave arrivals from two “big” earthquakes on four stations located within 0.5 km of vent (see fig. 1 for station locations). 

Note ~0.4-s difference in relative arrival times between stations AHAB and BLIS, indicating epicentral difference of more than one 

hundred meters for the two events.

nificant volumetric component, and they inferred that the source 

was a periodically collapsing, horizontal steam-filled crack.

Fracturing of new rock is also an unlikely process, as 

we would not expect to see repeating events or events with 

all-down first motions. Also, such a process would presumably not result in the emergent arrivals that we see at most 

stations in roughly half of the larger events. Fracturing within 

the spines is consistent with the observation that three of four 

large seismic-event clusters occur in association with the 

breakup of individual spines. However, events with all-down 

first motions are inconsistent with the observed failure mode 

in the spines, which occurred largely on shear, normal, and 

tensile fractures. Because the spines were, with one exception, 

higher than the seismometers, these failure modes would produce either a mixture of first motions or, in the case of normal 

faulting, dominantly up first motions. In addition, repetitive 

events spanning several months (and different spines) are 

incompatible with in-spine fracturing, because the fractures 

are destroyed in the spine breakup process.

Lurching of the active spine is our preferred hypothesis for 

the larger events. Lurches are consistent with shallow reverse 

faulting or upward thrusting of a piston that is indicated by the 

all-down first motions, as well as with long-lived event families. 

The fact that at least some events are located beneath the spine 

away from the vent indicates that they may have occurred along 

the interface between the base of the spine and the old crater floor. 

Because the 1980 crater floor is composed largely of ash and 

rockfall debris that fell from the crater walls in the years following the May 18, 1980, eruption, this interface likely is irregular 

and separates two poorly consolidated rock bodies. These lithologic and structural aspects could explain the emergent character 

of many larger seismic events and their low-frequency nature.

Volcano-Tectonic (VT) Earthquakes

Following the initial VT-rich days of the vent-clearing 

phase, VT events were a minor component of seismicity 

at Mount St. Helens. On a daily basis we scanned through 

seismic records in search of VT events, because we felt that 

VT events, especially those deeper than 2–3 km, might be 

an indicator either of a surge of new magma entering the 

system or adjustment of the magmatic system to magma 

withdrawal—and thus a potential harbinger of the eruption’s 

end. VT events occurred as isolated events throughout the 

eruption and had small magnitudes (Md <1), shallow (<2 km) 

depths, and epicenters similar to those for the drumbeats and 

larger events. The most notable occurrence of VT events was 

a swarm of more than 70 small (Md

 <0.3) shallow earthquakes 

occurring between December 27, 2005, and January 4, 2006. 

The VT events in this swarm occurred independently of the 

low-frequency drumbeats, in some cases occurring just tens 

of seconds before or after a drumbeat (fig. 24). This swarm 

appears quite clearly in figure 13, as only the VT events were 

large enough to pass our amplitude threshold for the ESAM 

plot during this time period. Poor weather obscured the 

volcano from view during this interval, but tiltmeters located 

on the 1980–86 lava dome showed slight inflation while VT 

events were occurring, in contrast to deflation before and 

after (D. Sherrod and M. Lisowski, oral commun., 2006). 

Other periods of inflation, however, do not coincide with VT 

events, so we are uncertain whether this relation is coincidental. Given the lack of correspondence of VTs to any obvious 

changes in eruption dynamics or mechanics of extrusion, our 

best guess is that the VT events represent fracturing of more 

competent rock in the crater floor, such as basaltic andesite 
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 49

lava flows from the Castle Creek period (2,200–1,895 yr B.P.) 

that underlie much of the 1980 crater (Clynne and others, this 

volume, chap. 28).

Summary—Patterns and Trends in 

Mount St. Helens seismicity, 2004–2005

In several intriguing ways the precursory seismic 

sequence of September 23 to early October 2004 was similar 

to those before the May 18, 1980, eruption and the 20 eruptions that followed during 1980–86. The 8-day-long interval 

between the first earthquakes and the first phreatic explosion 

on October 1, 2004, is close to the 7-day interval seen in 

March of 1980 (Endo and others, 1981; Malone and others, 

1981), suggesting a characteristic time interval for magma 

to reach the surface once it begins to move. In addition, the 

transition from VT events to hybrid and LF events starting 

September 25, 2004, was broadly similar to transitions in 

event type seen before many of the dome-building eruptions 

in the 1980s (Malone, 1983). The occurrence of the event 

families, starting with the first significant family on September 

26, 2004, is also similar to occurrence of event families before 

and during 1983–86 dome-building eruptions (Frémont and 

Malone, 1987), as well as to those in association with many 

other dome-building eruptions (for example, Okada and others, 1981; Power, 1988; Power and others, 1994; Villagómez, 

2000; Rowe and others, 2004). In retrospect, the combination 

of event-type transition and occurrence of event families on 

September 25, 2004, could have given us an early indication 

of the expected nature of eruptive activity at Mount St. Helens, 

although that knowledge would likely not have changed how 

CVO and the PNSN reacted to the developing seismic crisis.

A summary observation from the September 29–October 

5 period is that, of the hours-long increases in seismic intensity, only 4 of 12 occurred before explosions (fig. 9). Furthermore, seismic energy levels had stabilized at least an hour 

before two of the four explosions (October 1 and October 5). 

Thus we find no statistical basis for using short-term increases 

in RSAM or other measures of seismic energy to make shortterm forecasts of explosive activity.

The drumbeats and “big” earthquakes associated with the 

dome-building eruption represent a relatively new type of seismicity at Mount St. Helens that, with a few exceptions, was 

not a dominant characteristic of 1980–86 seismicity. Efforts 

Figure 24. Seismograms from three stations for a drumbeat and a VT event occurring within 50 s of each 

other on December 29, 2005, at 1643 and 1644 PST, respectively. See figure 1 for station locations.

YEL

HSR

SHW

TIME, IN SECONDS

0 20 40 60 80 100

DRUMBEAT VT
50 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

to characterize the nature of the drumbeat source are ongoing and likely will be the subject of research papers for years 

to come. Relatively new results with fascinating implications 

have already emerged from several field experiments run during 2004 and 2005, including the finding of infrasonic pulses 

associated with some drumbeats (Matoza and others, 2007) 

and the discovery of very long period (VLP) earthquakes in 

association with some of the larger drumbeats (Waite and 

others, 2008). These results provide evidence that processes 

operating at both deep (~500 m) and shallow (surface) levels 

are involved in the generation of drumbeats. The complexities 

in the relation between event spacing, event size, event spectra, 

and extrusion rate discussed in this paper, as well as in Thelen 

and others (this volume, chap. 4), further illustrate that the 

drumbeat source process is complex.

Finally, an important lesson that was reemphasized by 

the rapid escalation of seismicity during the first week of the 

eruption is that conditions can quickly become too dangerous to deploy additional seismometers after a volcano awakens. We were fortunate that the Mount St. Helens network 

installed in 1980–1983, particularly crater stations SEP 

(formerly REM) and YEL, was still in place and functional 

in 2004, because many of the events from the first two days 

were poorly recorded on stations outside of the crater. As the 

seismic crisis progressed, however, the limited dynamic range 

of the short-period analog stations was frequently exceeded 

at seismometers in the crater and on the flanks, resulting in 

lost information about seismic energy, spectral content, and 

event similarity. In addition, the absence of broadband threecomponent instruments in the crater meant that information 

was lost about the presence or absence of very low frequency 

signals and about source mechanisms for all types of events. 

As discussed by McChesney and others (this volume, chap. 7), 

seismicity intensified so rapidly during the first several days 

that we were unable to safely install such instruments in the 

crater. An important goal for seismic monitoring networks at 

any potentially active volcano should be to have seismometers, 

including broadband seismometers, installed at locations close 

to the presumed vent well before unrest begins.

Acknowledgments

The number of people involved in the seismic response to 

the renewed eruption in 2004–5 was huge, and it is impossible 

to give proper credit to all who helped us. Peggy Johnson, Josh 

Jones, Ruth Ludwin, Wendy McCausland, Karen Meagher, 

Guy Medema, Bob Norris, Bill Steele, George Thomas, and 

Tom Yelin all played critical roles in the PNSN/CVO seismic 

monitoring team, particularly with the immense tasks of data 

processing, production of Web graphics, around-the-clock 

monitoring of seismicity in the first several months of the 

eruption, and troubleshooting hardware problems brought 

on by the immense amount of seismic data and Web traffic. 

Marv Couchman, Rick LaHusen, Andy Lockhart, Jeff Marso, 

Pat McChesney, and Tim Plucinski all played critical roles in 

developing and installing new seismic and telemetry systems 

on and around the volcano, and we are extremely grateful to 

each of them for the many hours they devoted to improving 

real-time seismic monitoring at the volcano. Wendy McCausland calculated reduced displacement estimates for the October 2 and October 3 tremor episodes, and Bob Norris assisted 

with overnight shifts at CVO for more than a month, provided 

logistical support in the field, and was the first to recognize the 

similarity between 2004 drumbeats and “peppercorn” events 

seen in the 1980s.

We are grateful to Peter Cervelli for providing invaluable assistance in the drafting of all figures showing multiday 

seismic records and to Dave Ramsey for providing us with the 

station map showing crater station locations superimposed on 

the April 19, 2005, DEM in figure 1. We thank Jeff Johnson 

for his assistance with event-family identification during the 

first several weeks of the eruption and Helena Buurman for her 

assistance in processing and error-checking the results from 

our earthquake detection algorithm. We also thank Phil Dawson, John Power, Stephanie Prejean, and Randy White for their 

contributions to analysis and discussions of the significance of 

seismicity during the first critical weeks of renewed activity, for forming invaluable bridges to other researchers in the 

volcano seismology community when we were too busy to do 

so, and for providing greatly appreciated moral support during 

the first incredibly hectic weeks of the eruption. Finally, we 

are grateful to Charlotte Rowe and Randy White for thorough 

reviews that significantly improved this paper.

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2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 55

Appendix 1. PNSN Fault-Plane Solutions, 09/23/04–12/31/05

This appendix comprises numerous plots showing first motions and preferred fault-plane solutions for well-constrained 

events (maximum azimuthal gap <135o, nearest station <1 km, and at least 10 picked phase arrivals) (fig. 25) in the PNSN 

catalog from September 24 to November 27, 2004, and for “big” events from November 27, 2004 to December 31, 2005 (fig. 

26). We did not review fault-plane solutions from the PNSN catalog except those for the “big” events. Only those fault-plane 

solutions with at least 10 polarities are shown. Note that all fault-plane solution times are in Coordinated Universal Time (UTC), 

in contrast to the local time convention used in the rest of this paper. All fault-plane solutions were determined using the FPFIT 

computer program (Reasenberg and Oppenheimer, 1985).

20040924 0308 Z= 1.20 M= 1.40

T

P

20040924 0916 Z= 1.28 M= 1.00

T

P

20040924 1159 Z= 1.12 M= 0.80

T

P

20040924 1708 Z= 0.70 M= 1.30

T

P

20040924 1854 Z= 1.33 M= 0.00

T

P

20040924 1919 Z= 1.48 M= 1.80

T

P

20040924 2211 Z= 0.94 M= 1.10

T P

20040925 0015 Z= 1.17 M= 0.50

T

P

20040925 0412 Z= 1.04 M= 1.60

T

P

20040925 0451 Z= 0.94 M= 1.20

T

P

20040925 0531 Z= 1.12 M= 1.60

T

P

20040925 0702 Z= 0.92 M= 1.60

T

P

20040925 0738 Z= 0.34 M= 1.60

T

P

20040925 0756 Z= 1.12 M= 2.00

T

P

20040925 1452 Z= 1.31 M= 2.00

T

P

20040925 1521 Z= 1.25 M= 2.00

T

P

20040925 1616 Z= 1.37 M= 1.40

T

P

20040925 1907 Z= 0.94 M= 1.30

T

P

20040926 0314 Z= 1.00 M= 2.20

T

P

20040926 0352 Z= 1.53 M= 2.00

T

P

20040926 0417 Z= 1.02 M= 2.20

T

P

20040926 1012 Z= 1.16 M= 2.20

T

P

20040926 1037 Z= 1.33 M= 2.40

T

P

20040926 1223 Z= 0.87 M= 2.80

T

P

20040926 1925 Z= 1.26 M= 2.20

T P

20040927 0219 Z= 0.80 M= 1.90

T

P

20040927 2021 Z= 0.90 M= 1.70

T

P

20040928 0018 Z= 0.79 M= 1.80

T

P

20040928 0107 Z= 1.08 M= 1.50

T

P

20040928 0132 Z= 0.99 M= 1.90

T

P

Figure 25. Fault-plane solutions for 161 earthquakes occurring at Mount St. Helens during the period from Sept. 24 to Oct. 11, 2004. P, 

compressional axis; T, tensional axis. Open circles correspond to dilatational (down) first motions, closed circles to compressional (up) 

first motions. Each solution is coded for date and time (UTC) in this format: YYYYMMDD hhmm. Z, depth, in kilometers; M, magnitude.
Figure 25—Continued.

20040929 0333 

Z= 0.97 M= 2.00

T

P

20040929 0342 

Z= 1.22 M= 1.60

T P

20040929 1752 

Z= 0.70 M= 1.90

T

P

20040929 1847 

Z= 1.16 M= 2.80

T

P

20040929 1922 Z= 0.93 M= 2.70

T

P

20040929 1937 Z= 1.15 M= 2.50

T

P

20040929 2007 

Z= 0.89 M= 2.70

T

P

20040929 2021 

Z= 0.62 M= 2.60

T P

20040929 2026 Z= 0.92 M= 2.60

T

P

20040929 2035 

Z= 1.11 M= 2.60

T

P

20040929 2042 Z= 0.98 M= 2.60

T

P

20040929 2116 Z= 0.74 M= 2.60

T

P

20040929 2122 

Z= 1.07 M= 2.90

T

P

20040929 2154 

Z= 1.09 M= 3.00

T

P

20040929 2207 Z= 0.49 M= 2.60

T P

20040929 2231 

Z= 0.90 M= 2.50

T

P

20040929 2242 

Z= 1.10 M= 2.90

T

P

20040929 2341 Z= 1.07 M= 2.70

T

P

20040930 0

Z= 1.04 M= 2

T

P

20040930 0

Z= 1.09 M= 2

T

P

20040930 0Z= 1.16 M= 3

T

P

20040930 0

Z= 1.02 M= 2

T

P

20040930 0Z= 0.89 M= 3

T

P

20040930 0

Z= 1.11 M= 2

T

P

20040928 0145 

Z= 0.86 M= 1.80

T

P

20040928 0509 

Z= 0.90 M= 1.50

T

P

20040928 0611 

Z= 0.49 M= 1.90

T

P

20040928 0639 

Z= 0.04 M= 2.00

T

P

20040928 1610 

Z= 1.15 M= 2.20

T

P

20040928 1640 

Z= 0.88 M= 2.00

TP

20040928 1655 

Z= 0.92 M= 2.30

T

P

20040928 1734 

Z= 1.04 M= 2.09

T

P

20040928 1817 

Z= 0.96 M= 2.50

T

P

20040928 1854 

Z= 1.29 M= 2.30

TP

20040928 1905 

Z= 0.89 M= 2.09

T

P

20040928 1925 

Z= 0.91 M= 2.00

T

P

20040928 1947 

Z= 0.85 M= 2.20

T

P

20040929 0013 

Z= 0.94 M= 2.00

T

P

20040929 0119 

Z= 0.80 M= 1.50

T

P

20040929 0153 

Z= 0.74 M= 1.80

T

P

20040929 0214 

Z= 0.79 M= 2.00

T

P

20040929 0321 

Z= 0.93 M= 2.00

T

P
20041001 0720 

Z= 0.51 M= 3.50

T

P

20041001 0739 

Z= 0.52 M= 3.00

T

P

20041001 0751 

Z= 0.80 M= 3.20

T

P

20041001 1003 

Z= 1.28 M= 3.00

T

P

20041001 1041 Z= 0.52 M= 3.10

T

P

20041001 1051 Z= 0.89 M= 3.30

T

P

20041001 1143 

Z= 0.70 M= 2.80

T

P

20041001 1160 

Z= 0.88 M= 3.10

T

P

20041001 1205 Z= 1.19 M= 3.00

T P

20041001 1221 

Z= 1.14 M= 2.80

T

P

20041001 1252 Z= 1.04 M= 2.80

T

P

20041001 1258 Z= 1.05 M= 3.10

T

P

20041001 1322 

Z= 0.92 M= 2.90

T

P

20041001 1352 

Z= 0.71 M= 2.90

T

P

20041001 1403 Z= 1.04 M= 2.09

T

P

20041001 1504 

Z= 0.26 M= 3.30

T

P

20041001 1544 

Z= 0.36 M= 3.20

T P

20041001 1804 Z= 0.04 M= 3.10

T

P

20041002 05

Z= 0.19 M= 3.

T

P

20041002 11

Z= 1.07 M= 2.

T P

20041002 12Z= 0.79 M= 3.

T

P

20041002 14

Z= 0.89 M= 3.

T

P

20041002 14Z= 1.08 M= 3.

T

P

20041002 15

Z= 0.05 M= 2.

T

20040930 0723 

Z= 1.08 M= 3.10

T

P

20040930 0733 

Z= 1.08 M= 3.20

T

P

20040930 0753 

Z= 0.95 M= 3.10

T

P

20040930 1847 

Z= 1.16 M= 2.90

T

P

20041001 0043 Z= 0.87 M= 3.20

T

P

20041001 0134 Z= 0.71 M= 2.40

T

P

20041001 0136 

Z= 1.16 M= 3.00

T

P

20041001 0150 

Z= 0.35 M= 3.00

T

P

20041001 0253 Z= 0.52 M= 2.90

T

P

20041001 0323 

Z= 1.13 M= 3.20

T

P

20041001 0328 Z= 0.94 M= 3.30

T

P

20041001 0438 Z= 0.41 M= 3.10

T

P

20041001 0504 

Z= 0.94 M= 3.20

T P

20041001 0513 

Z= 0.97 M= 2.70

T

P

20041001 0529 Z= 1.11 M= 3.50

T

P

20041001 0551 

Z= 1.13 M= 2.90

T

P

20041001 0614 

Z= 0.80 M= 3.60

T

P

20041001 0645 Z= 1.02 M= 3.50

T

P

Figure 25—Continued.
Figure 25—Continued.

20041009 0514 

Z= 1.23 M= 2.50

T

P

20041009 0538 

Z= 1.31 M= 2.80

T P

20041009 0553 

Z= 1.14 M= 2.60

T

P

20041009 0611 

Z= 1.25 M= 2.60

T

P

20041009 0657 

Z= 0.98 M= 2.70

T

P

20041009 0725 

Z= 1.00 M= 2.60

T

P

20041009 0759 

Z= 0.64 M= 2.60

T P

20041009 0813 

Z= 1.30 M= 2.80

T

P

20041009 0823 Z= 0.91 M= 2.30

T

P

20041009 0826 

Z= 0.93 M= 2.30

T

P

20041009 0834 

Z= 1.13 M= 2.40

T

P

20041009 0850 

Z= 1.07 M= 2.70

T

P

20041009 0911 

Z= 1.25 M= 2.70

T

P

20041009 1016 

Z= 0.82 M= 2.50

T

P

20041009 1101 Z= 0.92 M= 2.90

T

P

20041009 1144 

Z= 1.00 M= 2.50

T

P

20041011 0758 

Z= 0.04 M= 2.20

T

P

20041002 2143 

Z= 0.33 M= 3.10

T

P

20041002 2151 

Z= 0.65 M= 2.30

T

P

20041002 2202 

Z= 1.00 M= 3.10

T

P

20041002 2212 

Z= 0.93 M= 2.20

T

P

20041002 2231 

Z= 1.38 M= 3.00

T

P

20041002 2247 

Z= 1.02 M= 3.00

T

P

20041002 2320 

Z= 1.03 M= 1.90

T

P

20041003 0559 

Z= 0.04 M= 2.60

T

P

20041003 0624 

Z= 0.88 M= 2.40

T

P

20041003 0635 

Z= 1.05 M= 3.50

T

P

20041003 0714 

Z= 1.17 M= 3.00

T

P

20041004 0637 

Z= 0.89 M= 2.40

P T

20041004 0757 

Z= 0.86 M= 2.90

T

P

20041005 0518 

Z= 0.03 M= 2.60

TP

20041009 0045 

Z= 1.02 M= 2.40

T

P

20041009 0425 

Z= 1.12 M= 2.40

T

P

20041009 0432 

Z= 0.70 M= 2.50

T

P

20041009 0442 

Z= 1.10 M= 2.50

T

P
2. Seismicity Associated with Renewed Dome Building at Mount St. Helens, 2004–2005 59

Figure 26. Fault-plane solutions for 55 earthquakes (Md 0.7–3.4) occurring at Mount St. Helens during the period from Nov. 27, 2004, 

to Oct. 7, 2005. P, compressional axis; T, tensional axis. Open circles correspond to dilatational (down) first motions, closed circles 

to compressional (up) first motions. Each solution is coded for date and time (UTC) in this format: YYYYMMDD hhmm. Z, depth, in 

kilometers; M, magnitude.

20041127 1301 

Z= 0.79 M= 3.10

T

P

20041129 1514 

Z= 0.60 M= 2.70

T P

20041217 0124 

Z= 0.04 M= 3.10

T

P

20041217 1009 

Z= 0.58 M= 3.10

T

P

20041217 1716 

Z= 0.17 M= 3.10

T

P

20041218 1018 

Z= 0.04 M= 2.90

T

P

20041218 1241 

Z= 0.98 M= 3.10

T

P

20041219 0011 

Z= 0.40 M= 3.10

T

P

20041219 1247 

Z= 1.12 M= 3.10

T

P

20050305 0728 

Z= 0.63 M= 2.80

T

P

20050305 1142 

Z= 0.58 M= 3.00

T

P

20050322 0950 

Z= 0.96 M= 2.50

T

P

20050323 0457 

Z= 1.02 M= 2.80

T

P

20050326 0021 

Z= 0.38 M= 3.20

T

P

20050326 1415 

Z= 0.63 M= 3.10

T

P

20050327 0128 

Z= 0.35 M= 3.20

T

P

20050327 0431 

Z= 0.05 M= 3.00

T P

20050327 0852 

Z= 0.88 M= 2.70

T

P

20050327 1146 

Z= 0.66 M= 3.00

T

P

20050327 1401 

Z= 0.51 M= 2.70

T

P

20050327 1726 

Z= 0.97 M= 3.00

T

P

20050327 2121 

Z= 0.03 M= 2.60

T

P

20050402 2217 

Z= 0.84 M= 2.80

T

P

20050403 1641 

Z= 1.19 M= 3.00

T

P

20050404 0543 

Z= 0.94 M= 3.40

T

P

20050404 0558 

Z= 0.68 M= 2.80

T

P

20050404 0620 

Z= 0.95 M= 2.40

T

P

20050404 1056 

Z= 0.62 M= 3.20

T

P

20050405 0416 

Z= 1.03 M= 3.30

T

P

20050406 2216 

Z= 1.11 M= 2.90

T

P

20050408 0619 

Z= 0.96 M= 2.60

T

P

20050409 0801 

Z= 1.03 M= 3.00

T

P

20050409 1545 

Z= 0.54 M= 2.50

T P

20050412 0715 

Z= 0.83 M= 2.90

T

P

20050424 0452 

Z= 0.24 M= 2.09

T

P

20050506 1326 

Z= 0.10 M= 0.70

T

P
60 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 2. Automated Event Detection

Analyst-driven earthquake location quickly became infeasible as a means for tracking the course of the eruption because 

it was too slow (Qamar and others, this volume, chap. 3), so over subsequent months we developed methods for automatically 

detecting events and recording several event parameters of interest. Figures 5, 13, and 14 show results from one such method, 

a standard triggering algorithm that uses the ratio of average seismic amplitudes computed over short- and long-term time 

windows, as applied to station HSR (fig. 1). We chose HSR over other stations for three principal reasons: (1) it is a sensitive 

site that usually records even small earthquakes cleanly; (2) wind noise is usually insignificant; and (3) after a gain reduction on October 4, 2004, no changes were made to the site (McChesney and others, this volume, chap. 7). We used averaging 

windows of 1 s and 8 s for the short- and long-term averages and a triggering ratio of 2.3 that was determined to be optimal 

for HSR through trial and error. To avoid multiple triggers from the same event, we skipped forward 6 s after each trigger. We 

also devised station-specific algorithms for distinguishing between noise glitches (including telemetry and/or Internet dropouts, 

radio interference, and calibration pulses) and events, although some glitches still made it through our filters. To test the impact 

of such noise-induced event detections, we visually scanned through helicorder plots for selected time periods and removed 

glitches that had been recorded as events. Although the resulting plots were cleaner, the fundamental trends were unchanged; in 

essence, noise was overwhelmed by the shear number of earthquakes.

For each detected event we recorded the peak amplitude over the next 6 s (usually the S wave at HSR), the time gap 

between successive events, and the normalized power spectra for the first 256 samples following the P-wave arrival (or 2.56 s 

with the 100-Hz sample rate used by the PNSN and CVO for all short-period stations). We chose to look only at spectra for the 

first 2.56 s (256 samples) following the event trigger, as full-event spectra were dominated by surface-wave energy and were 

thus not useful in detecting subtle changes in source spectra, particularly between hybrid and LF events. We only show spectra 

for events with amplitudes greater than 300 counts (Md

 ~1.0), because our algorithm commonly triggered on the S-wave for 

events smaller than this threshold. The resulting spectra plots, which we referred to as ESAM (Earthquake Spectral AMplitude) 

plots, proved to be more useful for assessing changes in earthquake source frequency than standard spectrograms of continuous 

seismic data, particularly because eruption-related seismicity was dominated by discrete seismic events.

20050724 1035 Z= 0.05 M= 3.30

T

P

20050729 0151 Z= 0.22 M= 2.40

T

P

20050729 1704 Z= 0.51 M= 3.20

T

P

20050731 0934 Z= 0.69 M= 3.30

T

P

20050821 0428 Z= 0.02 M= 2.60

T

P

20050827 1629 Z= 0.02 M= 2.09

T P

20050512 1914 Z= 0.71 M= 2.30

T

P

20050609 1151 Z= 0.38 M= 1.60

T

P

20050712 1205 Z= 0.50 M= 1.70

T

P

20050715 1222 Z= 0.37 M= 3.10

T

P

20050716 2008 Z= 0.45 M= 3.20

T

P

20050717 0637 Z= 0.89 M= 2.50

T

P

20050717 2103 Z= 0.18 M= 1.90

T

P

20050719 0354 Z= 0.02 M= 3.20

T

P

20050720 0519 Z= 0.02 M= 3.30

T

P

20050721 0960 Z= 0.22 M= 3.10

T

P

20050722 1747 Z= 0.19 M= 3.10

T

P

20050723 0938 Z= 0.23 M= 3.20

T P

20051007 1848 Z= 8.22 M= 1.70

T

P

Figure 26—Continued.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 3

Near-Real-Time Information Products for Mount St. 

Helens—Tracking the Ongoing Eruption

By Anthony I. Qamar1*, Stephen D. Malone1, Seth C. Moran2, William P. Steele1, and Weston A. Thelen1

1

 Pacific Northwest Seismic Network, Department of Earth and Space Sciences, University of Washington, Box 351310, Seattle, WA 98195

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

* Deceased

Abstract

The rapid onset of energetic seismicity on September 

23, 2004, at Mount St. Helens caused seismologists at the 

Pacific Northwest Seismic Network and the Cascades Volcano Observatory to quickly improve and develop techniques 

that summarized and displayed seismic parameters for use by 

scientists and the general public. Such techniques included 

webicorders (Web-based helicorder-like displays), graphs 

showing RSAM (real-time seismic amplitude measurements), 

RMS (root-mean-square) plots, spectrograms, location maps, 

automated seismic-event detectors, focal mechanism solutions, automated approximations of earthquake magnitudes, 

RSAM-based alarms, and time-depth plots for seismic events. 

Many of these visual-information products were made available publicly as Web pages generated and updated routinely. The graphs and maps included short written text that 

explained the concepts behind them, which increased their 

value to the nonseismologic community that was tracking 

the eruption. Laypeople could read online summaries of the 

scientific interpretations and, if they chose, review some of 

the basic data, thereby providing a better understanding of the 

data used by scientists to make interpretations about ongoing eruptive activity, as well as a better understanding of how 

scientists worked to monitor the volcano.

Introduction

The renewed activity of Mount St. Helens, which started 

in September 2004, caused the Pacific Northwest Seismic Network (PNSN) and the U.S. Geological Survey’s Cascades Volcano Observatory (CVO) to rapidly adjust routine monitoring 

procedures and activities in order to accommodate additional 

data volume, data types, and real-time monitoring requirements. The speed with which the precursory earthquake swarm 

developed, the sheer number of earthquakes involved, and the 

need for rapid analysis and interpretation resulted in many 

changes being made to what had become routine processing 

procedures developed over the preceding years. The goal of 

providing accurate and timely interpretation of the evolving 

sequence was driven by the need of the emergency-response 

community to mitigate possible hazards and by intense public 

interest, as indicated by the horde of media that descended on 

any site where information might be available (Driedger and 

others, this volume, chap. 24). Through the precursory seismic 

swarm, the initial minor explosions, and the subsequent domebuilding phase, seismic processing and display techniques 

were developed or expanded to help provide relevant information to the wide variety of users. This paper documents these 

techniques, focusing on why and how they were developed 

and what they contributed at the time to monitoring efforts.

Preexistence of Webicorders

Following the end of the previous eruption of Mount St. 

Helens in October 1986, seismicity at the volcano was monitored using standard seismic techniques, including earthquake 

locations and visual monitoring of continuous seismic records 

in the form of helicorders and webicorders. Data from 12 

short-period, analog seismic stations were telemetered to the 

Pacific Northwest Seismic Network (PNSN) headquarters at 

the University of Washington (UW) for recording and analysis. A subset was also telemetered to the Cascades Volcano 

Observatory (CVO) for monitoring purposes.

The PNSN routinely processed events from Mount St. 

Helens through an event-triggered recording system as part 
62 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

of the routine catalog generation for the whole Pacific Northwest. These catalogs were originally only released in quarterly 

reports and distributed by surface mail. With the rapid evolution of the World-Wide Web, in the early 1990s the PNSN 

created a public Web site (www.pnsn.org) for online distribution of epicenter catalogs, maps, and other derived information 

products on earthquakes and volcanoes. As is common for 

many seismic networks, the PNSN also provided near-realtime Web images of selected seismograms, referred to as 

webicorder plots (fig. 1).

By 2004, webicorder plots were being produced for more 

than 60 of the approximately 250 PNSN seismic stations, 

including several Mount St. Helens stations. As soon as seismic unrest was recognized at Mount St. Helens (September 

23, 2004), additional stations there were added to the volcano 

webicorder list at the expense of stations at other volcanoes; 

1800

1700

PDT

Station HSR, south flank, Mount St. Helens

TIME, IN MINUTES

1900

2000

2100

2200

2300

September 26 September 27

DATE AND HOUR (2004)

0000

0100

0200

0300

0400

0 1 2 3 4 5 6 7

Figure 1. Part of a webicorder plot from station HSR on the south flank of Mount St. Helens from September 26 and 

27, 2004. Time advances from left to right and top to bottom, as if reading a book. Each line begins a new 10-minute 

cycle (the last 3 minutes of each cycle are clipped off here), and gray vertical lines mark 1-minute intervals; color 

variations are merely to aid the eye in separating traces. Webicorder plots such as this one were updated in real time, 

available online before and during the 2004 restlessness and ensuing eruption, and were very popular with the public.

processing constraints limited the total number of webicorders 

that could be accommodated at any one time.

The webicorder plots were quickly discovered by many 

members of the public. By September 28, the number of hits 

on the PNSN departmental Web server was causing a substantial slowing of computer processing and increased Web-page 

delivery times. Because of the obvious public interest and the 

PNSN’s mission to provide near-real-time information to the 

public, the PNSN requested assistance from the University of 

Washington Computers and Communications (CAC) group. 

Within 18 hours of our request the CAC had a high-capacity 

Web server connected directly to external high-volume routers 

to mirror our main pages. Within a few days this server was 

servicing requests at rates as high as 250,000 per hour, and the 

CAC group expanded the system to dual server machines with 

load-balancing routing. The first week’s activity totaled more 
3. Near-Real-Time Information Products for Mount St. Helens—Tracking the Ongoing Eruption 63

than 23 million hits, which translates to roughly 5 million 

complete pages of information served.

In addition to establishing a mirror site, the PNSN also 

established a semiprivate (nonlinked but otherwise open) 

Web server for use by the scientific staff at UW, CVO, and 

other institutions involved in monitoring the evolving crisis. 

A password-protected Web server at CVO hosted additional 

webicorders and other seismicity-related plots. The mix 

of public and private Web servers enabled us to satisfy the 

public’s need for information while maintaining real-time 

data-sharing capabilities between CVO and the PNSN. Given 

the physical separation between the PNSN (located in Seattle, 

Wash.) and CVO (290 km south in Vancouver, Wash.), the 

capability for real-time data sharing was critical for effective 

telephone discussions between CVO and UW staff regarding 

the evolution of the unrest and eruption.

Real-Time Seismic Amplitude 

Measurements (RSAM)

Webicorders were valuable for seismologists to qualitatively gauge changes in event type, frequency, and size over 

the previous minutes to hours, but other tools were necessary to provide a quantitative basis for tracking seismicity 

changes through time. One such tool, which was quickly 

implemented at CVO (September 23), was the plotting of 

real-time seismic amplitude measurement (RSAM) values (Endo and Murray, 1991) (fig. 2). RSAM plots show 

the rectified amplitude of ground motion averaged over 

specified time intervals, commonly 10 minutes. Rather than 

focusing on individual events, RSAM provides a simplified 

but useful measure of the overall level of seismic activity. 

RSAM has become a widely used tool in volcano observatories around the world because it readily and quantitatively 

reflects changes in number and size of earthquakes, tremor, 

and background noise, each or all of which may be related 

to changes in volcanic activity. As a result of their previous 

widespread use, RSAM plots were generally well understood 

by all scientists at CVO and the PNSN.

The Earthworm seismic-data acquisition and processing 

system (Johnson and others, 1995; U.S. Geological Survey, 

2006) is used both at CVO and at UW for the collection and 

analysis of seismic data. The “ew2rsam” module, a part of the 

Earthworm system, was used in the Mount St. Helens crisis to 

generate RSAM values in real time and plot them on multistation graphs at intervals ranging from several days to several 

months. Such multistation plots make it easier to determine 

whether seismic-amplitude changes result from local effects at 

each station (such as wind noise) or from volcanic processes. 

The RSAM plots were particularly important for recognizing 

trends when seismicity intensified on September 26, because 

webicorder plots for stations on or near the edifice became 

increasingly saturated and difficult to interpret (Moran and 

others, this volume, chap. 2).

14 15 16 17 18 19 20 21

DATE IN OCTOBER 2004

22 23 24 25 26 27

10

20

30

40

50

RSAM AMPLITUDE, IN COUNTS RMS AMPLITUDE, IN COUNTS RMS AMPLITUDE, IN COUNTS

Station HSR: 10-minute RMS

100

50

0

150

200

100

50

0

150

200

Station HSR: 1-minute RMS

Station HSR: 10-minute RSAM

Figure 2. Real-time seismic amplitude measurement (RSAM, 

top) and root-mean-square (RMS, middle and bottom) plots for 

seismic signal levels from station HSR during mid- to late October 

2004, early in the eruption. Besides the different algorithm to 

generate the points, plotting characteristics are also different 

to emphasize different aspects of the activity. Note how the 

10-minute and 1-minute RMS plots differ in resolution for the 

variability of seismic activity.

Root-Mean-Square (RMS) Plots

The usefulness of the RSAM plots inspired the creation 

at the PNSN of root-mean-square (RMS) plots (fig. 2). Rather 

than just plotting the average of the rectified signal, these plots 

show the square root of the sum of squares of the signal averaged over fixed time intervals. Also, different plotting styles 

for the CVO RSAM and the PNSN RMS were used to emphasize different aspects of seismicity changes. Whereas the 

RSAM plots connect the points of each average with a line, 

the RMS plots just show individual averages over a fixed window length as isolated points. Such plots were routinely made 

for averaging windows of 1 minute and 10 minutes and posted 
64 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

in nearly real time on the PNSN public Web page. Although 

the RSAM and RMS plots generally show similar trends, the 

RMS plots can isolate changes caused by a few large events 

from cases in which there were an increased number of 

moderate-size events. From what we know, the inclusion of 

these plots on the PNSN public Web pages was the first time 

anywhere that reduced data, other than earthquake locations, 

were automatically released to the public in real time over the 

course of a volcanic eruption.

A modification of the RMS plotting routine was developed later in the eruption to show the largest and smallest 

RMS values determined over a specific time interval. For 

these plots, RMS values were computed every 5 s and only the 

largest and smallest during a 30-minute period were selected 

to plot. In effect these plots track the largest earthquakes and 

the lowest background level. The maximum-value RMS plots 

were used for tracking changes in maximum earthquake size 

over periods of days to weeks, whereas the minimum-value 

RMS plots were used for detecting elevated background levels 

that could have represented volcanic tremor (fig. 3).

Spectrograms

Spectrograms, sometimes called “Seismic Spectral 

Amplitude Measurements” (SSAM) (Rogers and Stephens, 

1995) are plots of signal frequency versus time in which color 

or intensity is used to display the relative strength of a signal 

in a frequency band (fig. 4). Seismic spectrograms are generated by taking the Fourier spectrum of a fixed time window 

of signal (1 minute on 12-hour spectrograms at UW and 2.5 s 

on 10-minute spectrograms at CVO), smoothing the spectrum, 

and then plotting the amplitude as a color (warm for high 

amplitude, cool for low) on the vertical frequency axis.

10

0

20

30

40

AMPLITUDE, IN COUNTS

14

DATE IN OCTOBER 2004

15 16 17 18 19 20 21 22 23 24 25 26 27

Minimum RMS amplitude

100

0

200

300

400

500

600

Maximum RMS amplitude

Figure 3. Maximum root-mean-square (RMS, top) 

and minimum RMS (bottom) plots for seismic-signal 

levels from station HSR during mid- to late October 

2004, early in the eruption. Time period same as in 

figure 2. Minimum RMS shows background levels, 

which increased and diminished over roughly 

weekly intervals. Maximum RMS shows a substantial 

decrease in large earthquakes during this period.

Spectrograms for different sets of stations were being 

created at UW (12 hours per plot) and at CVO (10 minutes 

per plot) before September 23, 2004. The spectrograms added 

little to the interpretation of event types in the first few days of 

the seismic buildup. By September 26, however, they became 

useful for recognizing the decrease in higher frequencies 

as the earthquake sequence progressed from predominantly 

volcano-tectonic events, with their wide spectrum, through 

hybrids to dominantly low-frequency events, described by 

some as long-period (LP) events (Moran and others, this volume, chap. 2; Thelen and others, this volume, chap. 4). These 

changes of earthquake character were obvious on seismograms 

of individual events, but the spectrograms allowed us to track 

changes over time as well. Spectrograms also aided in distinguishing rockfalls, which have broad spectra and long durations, from small, emergent earthquakes. Spectrograms have 

continued to play an important role in the subsequent years of 

dome building, because they display variations in seismicity 

that correspond to different types of events.

Event Processing

At the PNSN, locations and magnitudes of earthquakes 

in the Pacific Northwest are routinely determined by manual 

review of waveforms selected by the Earthworm automatic 

event-triggering-and-recording system. This system uses a 

detector algorithm that determines a station “trigger” based on 

the ratio of short-term average over long-term average (STA/

LTA) exceeding a specific threshold. Several station triggers 

must be active at roughly the same time from a subnetwork 

of nearby stations for the system to declare an event. This 

criterion helps to discriminate seismic events of real interest from noise bursts on individual stations. The triggering 

algorithm is tuned to be sensitive, so that very few 

events of interest are missed, at the cost of having 

many false triggers. Such a triggering system, in 

one form or another, has worked well since 1980, 

having detected for processing more than 100,000 

earthquakes between 1980 and 2004.

With the onset of the Mount St. Helens seismic swarm in late September 2004, the routine 

processing procedures for the PNSN quickly 

were swamped with data. The automatic system 

triggered and recorded the early earthquakes with 

magnitudes as small as Md

 = –1.0. Event trigger-
3. Near-Real-Time Information Products for Mount St. Helens—Tracking the Ongoing Eruption 65

ing rates for the whole network went from an average of 2–5 

events per hour (more than 50 percent noise triggers) before 

the swarm onset to 6–10 per hour by the end of the second 

day and to 25–40 per hour by the end of the third day. Soon 

thereafter the system was in an almost continuously triggered 

state. By the end of the second day, more than 250 events had 

been processed, and by the third day the analysis staff could 

not keep up with manually processing all triggered events. 

In addition to the burden of reviewing the triggered events, 

the volume of waveform data was filling available disk 

space, forcing us to move unprocessed data to tape and other 

media. Triage was necessary to ensure that at least a reasonable sampling of events were manually reviewed and that no 

trace data were lost. The task was twofold: first, we trained 

additional staff and students in the basic process and divided 

review tasks among them; then we experimented with different trigger parameters to achieve a representative, but not 

complete selection of events for manual processing.

After the first two months, with seismic activity continuing at a fairly high rate and the potential for a long-lasting 

eruption becoming evident, we changed the procedures in such 

a way that we could review the data for significant changes 

1720

0

5

10

0

5

10

0

5

10

0

5

10

0

5

10

0

PACIFIC STANDARD TIME (hhmm) ON MARCH 8, 2005,

IN 24-HOUR NOTATION

Seismic trace, vertical scale arbitrary

Spectrogram, showing frequency

5

10

1721 1722 1723 1724 1725 1726 1727 1728 1729 1730 MIDE

SEISMIC STATIONS

STDZ SHWZ JUN

FREQUENCY, IN HERTZ

HSR CDF

Figure 4. Sample spectrograms from six 

stations for a 10-minute period on March 

8, 2005 (local date). Shown for each station 

are conventional seismogram (upper line 

trace) and corresponding spectrogram (color 

spectrum). Periodic earthquakes characterize 

the record until about 17:25:30 (1:25:30 UTC), 

when an explosion at the vent marks onset of 

sustained, low-frequency tremor. Station MIDE, 

lowermost of graphs, lost signal sporadically 

and then completely, owing to thick, airborne 

ash that interrupted radio transmission.

but process in detail only a selection of the larger or impulsive 

events. For this we started generating hour-long, sequential 

artificial trigger files of only Mount St. Helens stations, with 

a display order based on distance from the vent so that the 

waveforms could be easily scanned for events of interest and 

for changes in event or background signal characteristics. 

The trigger threshold for the automatic triggering system was 

raised for the Mount St. Helens subnet so that only larger 

events would generate a complete event trigger. This procedure sped up the routine analysis and provided a better way 

of detecting subtle changes in event types or characteristics, 

because an analyst could quickly scan through all of the data 

and easily see changes in both large and small events.

The hypocenter distribution determined by standard 

processing of well-recorded earthquakes changed very little 

over the course of the seismic restlessness and ensuing eruptive sequence. Hypocenter depths decreased during the first 

two weeks of restlessness from 2–3 km deep to less than 1 km. 

The scatter in epicenters is comparable to variation arising 

from picking errors. Some of the variation in locations can be 

attributed to changes in station configuration. Several times 

during the sequence, no station was located within 1.5 km of 

the source area because of station outages 

resulting from explosions, and thus depth 

control, in particular, was poor. Subsequent 

specialized analysis shows a much tighter 

clustering of hypocenters than that determined 

during the height of the sequence (Thelen and 

others, this volume, chap. 4).

Automatic Trace Processing

Once it became apparent that the dominant seismic signals were from very regular 

earthquakes—so regular that they were named 

“drumbeat” earthquakes—we felt a need for 

tools to track the rate and type of individual 

earthquakes. A modification of an eventtriggering algorithm was developed to detect 

characteristics of individual events. This algo-
66 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

rithm was based on a ratio of short-term and long-term averages 

of RMS values computed at 5-s intervals for a key subset of 

close-in stations with good signal-to-noise ratios. The onset of 

an event was reliably detected, but determining the end of one 

event so that the algorithm would de-trigger and be ready for the 

next event required some tuning of parameters. Once an event 

was declared, several parameters, including its peak amplitude, 

average RMS value, and peak frequency, were determined from 

the original waveforms. The event was not located in the traditional sense, but an effective automatic event catalog was generated that included time, size, and general frequency content.

Focal Mechanisms

Focal mechanisms are traditionally difficult to obtain in 

volcanic areas because the emergent arrivals typical of lowfrequency earthquakes and the high attenuation within the 

edifice result in low signal-to-noise ratios at more distant stations. Additionally, at Mount St. Helens, earthquakes located 

at shallow depth in the edifice are mostly small, contributing 

to poor signal-to-noise ratios. Shallow hypocenters mean that 

sampling the entire focal sphere, a requirement for a well-constrained focal mechanism, is difficult. Determination of focal 

mechanisms is therefore restricted to the relatively few large 

events for which good first motions are available on many stations. Typically focal mechanisms were generated at the PNSN 

during the event-location process by picking polarities (up 

or down) of the first motion. A modification of the program 

“FPFIT”( Reasenberg and Oppenheimer, 1985) was used to 

fit focal planes for a double-couple solution when there were 

10 or more picked polarities. An unusual aspect of Mount St. 

Helens earthquakes is that a vast majority of first motions are 

dilatations. Many events have dilatations on all stations, giving 

the impression of an implosive or tensile source. However, 

because of poor sampling of the focal sphere, it is sometimes 

possible to fit a double-couple solution even to these cases. 

Figure 5 shows examples of four typical events with predominantly “all down” first motions. Of these four events only one 

clearly does not have a possible double couple solution.

Event Magnitudes

Routine earthquake processing by the PNSN uses coda 

duration to compute magnitudes (Crosson, 1972). For most 

tectonic earthquakes a consistent linear relation exists between 

the log of the coda duration and the local magnitude, ML. 

However, this relation breaks down for earthquakes recorded 

on Cascade Range volcanoes. For shallow earthquakes in 

particular, the coda duration is much longer than expected for 

a given magnitude. Two calibrated broadband stations were 

installed near Mount St. Helens in October 2004, early in 

the eruption (McChesney and others, this volume, chap. 7). 

By deconvolving their known instrument responses and then 

Compression

Dilatation

T P

12/01/2006 1609 PST M=3.4

T

P

07/23/2005 0238 PDT M=3.2

T

P

07/18/2006 0955 PDT M=3.5

T P

11/27/2004 0501 PST, M=2.1

Figure 5. Examples of focal mechanisms determined from 

four of the many larger, well-recorded events. All such 

events are located within a few hundred meters of the vent. 

Note that while almost all polarities are dilatations, doublecouple solutions can be fit in some cases. In other cases, 

such as the event of July 18, 2006, no double-couple solution 

is possible without many (five) inconsistent polarities.

convolving a Wood-Anderson response, equivalent maximumtrace amplitudes on standardized Wood-Anderson seismograms could be calculated. The resulting local magnitude (ML) 

was determined for a suite of earthquakes over the magnitude 

range 0.5–3.4 (fig. 6). Coda durations for the larger events, 

mostly picked on stations distant from the volcano, gave duration magnitudes (Md

) comparable to the local magnitude (ML). 

For smaller events, however, Md is often grossly overestimated 

owing to the extended codas generated by the volcanic earthquakes recorded at stations on volcanic rocks.

Even though the coda duration magnitudes overestimate 

the true magnitude of the earthquake, coda magnitudes were 

still measured for located earthquakes to give a quantitative 

measure of the relative sizes of earthquakes at Mount St. Helens. Codas for the calculation of magnitude are measured only 

on stations off the edifice (beyond a 5-km radius) to minimize 

the effect of the extended codas.

Special Web Pages

For many years the PNSN Web site has maintained topical Web pages for all of the monitored Cascade Range volcanoes. Besides some general descriptive text for each volcano 

and links to the more extensive descriptions on CVO Web 
3. Near-Real-Time Information Products for Mount St. Helens—Tracking the Ongoing Eruption 67

1 2 3

1

2

3

CODA DURATION MAGNITUDE (Md)

LOCAL MAGNITUDE (

ML)

Figure 6. Coda duration magnitude versus local magnitude for a 

selection of Mount St. Helens earthquakes illustrates how use of 

coda duration magnitudes overestimates the local magnitudes for 

these earthquakes, particularly at lower magnitudes.

pages, these pages provide plots of epicenters, depth versus 

time, and energy release versus time at three different time 

scales. The plots have always been based on routine processing of earthquakes located at a volcano. During inter-eruption 

periods and for minor seismic swarms, these pages continue 

to be useful for seismologists and informative for the general 

public. As the current seismic sequence escalated, however, 

the plots often became outdated because the routine processing 

(and thus the catalog) was greatly delayed or incomplete. Thus 

the utility of these special pages to assist in the interpretation 

of changes was severely compromised. The automatic RSAM 

and RMS plots quickly became favored as the data product 

most used for tracking changes in the seismicity at Mount St. 

Helens. However, public interest in the special pages remained 

high, even though we frequently posted disclaimers concerning the accuracy and completeness of these plots.

By adopting the changes already described in this paper, 

earthquake processing again became routine enough that the 

special volcano-related pages for Mount St. Helens represented a near-real-time picture of seismicity, even though only 

a fraction of detected earthquakes were logged in the catalog 

of located events. For example, time-depth plots (fig. 7) of 

located events eventually made it obvious that the depths of 

earthquakes taking place beneath Mount St. Helens since 

the 1980s changed dramatically with the advent of the 2004 

eruption. Earthquakes at depth of 2–3 km, with occasional 

deeper swarms, had been the norm at Mount St. Helens since 

the late 1980s. These events, particularly the deeper swarms, 

were interpreted as representing the repressurization of the 

magmatic system by recharge (Moran, 1994; Musumeci and 

others, 2002). With the advent of the 2004 (and ongoing) eruption, these deeper events ceased. This depth-pattern shift hints 

that if magma recharge is continuing during the eruption, it is 

not pressurizing the deeper system.

RSAM Alarms

In addition to magnitude-based earthquake alarms 

routinely used by the PNSN, several eruption-specific alarms 

were developed by CVO during the first 6 months of the eruption. RSAM alarms tuned to detect both large, discrete events 

(“event alarm”) and smaller amplitude but longer duration 

events (“tremor alarm”) were employed by CVO in mid-October 2004, shortly after the end of intense seismicity accompanying the vent-clearing phase. These alarms had previously 

been developed by the USGS Volcano Disaster Assistance 

Program as part of the “Glowworm” package of modules that 

operate as an add-on to the Earthworm acquisition system 

(Marso and others, 2003a). Before 2004, RSAM alarms had 

been used for more than a year to monitor the eruption of Anatahan volcano in the Commonwealth of the Northern Mariana 

Islands (Marso and others, 2003b).

The “event alarm” tracks RSAM values determined 

over a 2.56-s window (for 100-Hz data), with alarms issued 

if thresholds are exceeded at a prescribed number of stations 

for longer than 8 s. The “tremor alarm” uses RSAM values 

for 1-minute windows. When an alarm is issued, a separate 

module sends alerts via Short Message Service (SMS) to cell 

phones carried by on-call scientists. A potential weak link in 

this chain is the use of SMS, which requires that e-mail servers 

and SMS services be fully functional. Alerts were occasionally 

delayed by several minutes or even an hour before reaching 

individual phones. To partly address this limitation, SMS messages were sent to multiple cell phones (13 as of the summer 

of 2006) using multiple carriers. Nevertheless, these problems 

kept us from relying completely on SMS messaging, which we 

augmented by periodic data checks during off-hours.

Alert thresholds for the “event” and “tremor” alarms 

required refinement throughout the first 6 months of the 

eruption. Seismicity was so energetic initially that the alarms 

had to be desensitized to minimize the number of alerts. In 

mid-November 2004, larger earthquakes began occurring 

daily (Moran and others, this volume, chap. 2), and the “event 

alarm” was set so that thresholds were just barely exceeded 

by these events. Thresholds for the tremor alarm were much 

harder to determine, as we had no quantitative means for 

establishing thresholds. They initially were set at nominal 

values that were sufficiently high to prevent false triggers from 

wind and other noise sources. As a result, no alarms were 

issued for the January 16, 2005, explosion, which was accompanied by low-amplitude tremor that was barely detectable 

on stations outside the crater (Moran and others, this volume, 

chap. 6). Tremor alarm thresholds were reset on the basis of 
68 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

3

6

9

12

DEPTH, IN KILOMETERS

1997 1998 1999 2000 2001

YEAR

1996 2002 2003 2004 2005 2006

Figure 7. Time versus depth plot for a decade of seismicity at Mount St. Helens. Plot shows approximately 24,000 earthquakes 

located directly under the volcano (within 6-km radius of the new dome) between January 1997 and the end of 2006. Size of circles 

is proportional to event magnitude. Green circles, events in 2006; red circles, events in December 2006. 

recordings of this explosion on stations located within the 

crater. However, the closest station became intermittent shortly 

after the January explosion and had stopped transmitting altogether by the time of the March 8, 2005, explosion. Seismicity 

associated with the March 8 explosion was more energetic 

(Moran and others, this volume, chap 6), but thresholds were 

not exceeded on enough stations to generate an alarm.

The failure of these alarms to generate alerts during 

explosions motivated us to develop a new alarm in the event 

of a loss of analog telemetry. Of the six explosions occurring 

during 2004–5, four, including the January 16 and March 8, 

2005, explosions resulted in temporary or permanent loss of 

telemetry from at least one station. Traditional RSAM alarms 

have always used RSAM values from which the electronic bias 

or direct-current (DC) component has been removed, in large 

part to avoid false alarms resulting from common misalignment 

problems with analog telemetry circuitry that can impose a 

small DC offset on seismic data. However, the USGS J120 discriminator includes circuitry that detects the loss of the subcarrier and sets the discriminator output to a static 0.7-V DC level 

to distinguish loss of telemetry from a quiet seismic signal.

We created a new Earthworm module (based on the 

Earthworm ew2rsam module) that does not remove the DC 

component from RSAM values. Named FLOTSAM (not an 

acronym), this new alarm system sends alerts whenever values 

exceed a preset threshold at any station. The FLOTSAM 

system has sent out several alerts, some of which have been 

generated by loss of telemetry from crater stations destroyed 

during large rockfalls. Other alerts have resulted from stations 

that have stopped transmitting because of power outages at the 

site or weather-related disruptions of the radio signal.

The FLOTSAM alarm has become the alarm that the 

PNSN and CVO rely upon the most for detecting explosions, 

because no fine-tuning of amplitude and duration thresholds is 

required. All that is required is at least one seismometer placed 

close enough to the vent to have its radio signal disrupted by 

ash or by physical damage to the station. The event and tremor 

alarms are employed as well, with the event alarm proving to 

be a reliable detector of large earthquake-generated rockfalls, 

which sporadically produce ash clouds that reach from hundreds to several thousand meters above the crater rim.

Real-Time Traces for Remote Use

Early in the Mount St. Helens 2004 eruption, we became 

aware that researchers at other volcano observatories and academic institutions were interested in our seismic data. Besides 

the public Web pages, we provided access over the Internet to 

our real-time trace data wave-servers on a case-by-case basis. 

We also set up a direct export of selected data to the Alaska 

Volcano Observatory (AVO), where researchers were interested in following the seismic sequence in detail. At AVO and 

Stanford University the real-time data integration tool, VALVE 

(Cervelli and others, 2002), was used to track changes in 
3. Near-Real-Time Information Products for Mount St. Helens—Tracking the Ongoing Eruption 69

seismic activity along with other parameters. The seismic-trace 

display program, SWARM (Alaska Volcano Observatory, 2006) 

was used at several different institutions to examine continuous 

seismic waveforms and their frequency content. In particular, 

researchers at the following institutions became near-real-time 

collaborators participating to some degree in the interpretation 

of specific aspects of Mount St. Helens seismicity:

•	 Alaska Volcano Observatory (Anchorage and Fairbanks)—event character, multiplet analysis

•	 U.S. Geological Survey (Menlo Park, Calif.)—event 

character, frequency shift

•	 University of California, Los Angeles—strong-motion 

portable monitoring, explosion analysis

•	 University of Memphis (Tennessee)—broadband portable monitoring

•	 Los Alamos National Laboratory (New Mexico)— 

multiplet analysis

•	 University of New Hampshire—multiplet analysis, 

acoustic signals.

Public Use and Response

The availability of relatively detailed information on the 

Web in nearly real time for this eruption of Mount St. Helens 

provided a new and powerful tool for public understanding. The availability of near-real-time original data (seismic 

traces) and analyzed products (RMS plots, maps, and timedepth plots) resulted in a change in the public perception of 

the eruption and their interaction with scientists. In previous 

eruption crises, the public has learned about the process primarily through the eyes of the news media, who may witness 

or photograph the physical events and talk with scientists 

about their interpretations. In this latest eruption of Mount 

St. Helens, the public could see for themselves on the Web 

many of the same data and products that the scientists were 

using. This availability improves the public understanding 

in two significant ways: Although the public still obtained 

the basic summary information from the news media, those 

wanting more or not trustful of the interpretations they were 

hearing could look at some of the data themselves. 

The current eruption of Mount St. Helens is remarkable for the near absence of rumors and conspiracy theories 

purporting that critical information was being deliberately 

withheld from the public by scientists and emergency managers. The public also demonstrated the ability to digest a 

variety of sophisticated information products in order to follow the course of the eruption. Providing rich, open sources 

of eruption data enabled the public and media to monitor the 

ongoing eruption and may have encouraged them to seek 

further clarification from the scientific community when 

issues of interpretation arose. This public availability limited 

the kind of wild speculations that have sometimes occurred 

in the past.

Acknowledgments

The many staff and students associated with the Pacific 

Northwest Seismic Network (PNSN) and Cascades Volcano 

Observatory are too many to acknowledge individually but 

are greatly appreciated for their tireless help in the field and 

lab, often going much beyond what was expected to make 

the acquisition and processing of the huge amount of seismic 

data possible. Tom Murray, Peter Cervelli, Randy White, Jeff 

Johnson, Jackie Caplan-Auerbach, Chris Newhall, Giuliana 

Mele, Stephanie Prejean, and Steve Horton are thanked for 

their contributions to software, Web design, and general discussions relevant to our information products. David Sherrod 

and Jackie Caplan-Auerbach provided excellent reviews of 

our preliminary manuscript. The University of Washington’s 

Office of Computing and Communications is acknowledged 

for their very rapid and effective Web server upgrades and 

extra telephone services. Partial support was provided by U.S. 

Geological Survey Cooperative Agreement 04HQAG005 for 

the operation of the PNSN.

References Cited

Alaska Volcano Observatory, 2006, Alaska Volcano Observatory software development page, SWARM version 1.2.3: 

U.S. Geological Survey online-only publication, including 

tutorial, manual, and software [http://www.avo.alaska.edu/

Software/swarm/; last accessed Mar. 17, 2009].

Cervelli, D.P., Cervelli, P., Miklius, A., Krug, R., and 

Lisowski, M., 2002, VALVE; Volcano Analysis and Visualization Environment [abs.]: Eos (American Geophysical 

Union Transactions), v. 83, no. 47, Fall Meeting Supplement, U52A-07, p. F3.

Crosson, R.S., 1972, Small earthquakes, structure and tectonics of the Puget Sound region: Seismological Society of 

America Bulletin, v. 62, p. 1133–1171.

Driedger, C.L., Neal, C.A., Knappenberger, T.H., Needham, 

D.H., Harper, R.B., and Steele, W.P., 2008, Hazard information 

management during the autumn 2004 reawakening of Mount 

St. Helens volcano, Washington, chap. 24 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. Geological Survey Professional Paper 1750 (this volume).

Endo, E.T., and Murray, T., 1991, Real-time seismic amplitude 

measurement (RSAM); a volcano monitoring and prediction 
70 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

tool: Bulletin of Volcanology, v. 53, no. 7, p. 533–545.

Johnson C.E., Bittenbinder, A., Bogaert, B., Dietz, L., and 

Kohler, W., 1995, Earthworm—a flexible approach to seismic 

network processing: IRIS Newsletter, v. 14, no. 2, p. 1–4.

Marso, J.N., Murray, T.L., Lockhart, A.B., and Bryan, C.J., 

2003a, Glowworm, an extended PC-based Earthworm system for volcano monitoring [abs], in Cities on Volcanoes 3: 

Hilo, Hawaii, International Association of Volcanology and 

Chemistry of the Earth’s Interior (IAVCEI), Hilo, Hawaii, 

July 14–18, 2003, Abstract volume, p. 82.

Marso, J.N., Lockhart, A.B., White, R.A., Koyanagi, S.K., 

Trusdell, F.A., Camacho, J.T., and Chong, R., 2003b, The 

Anatahan volcano-monitoring system [abs.]: Eos (American 

Geophysical Union Transactions), v. 84, no. 46, Fall Meeting Supplement, V32B-1020.

McChesney, P.J., Couchman, M.R., Moran, S.C., Lockhart, 

A.B., Swinford, K.J., and LaHusen, R.G., 2008, Seismicmonitoring changes and the remote deployment of seismic 

stations (seismic spider) at Mount St. Helens, 2004–2005, 

chap. 7 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., 

eds., A volcano rekindled; the renewed eruption of Mount 

St. Helens, 2004–2006: U.S. Geological Survey Professional Paper 1750 (this volume).

Moran, S.C., 1994, Seismicity at Mount St. Helens, 1987–

1992: Evidence for repressurization of an active magmatic 

system: Journal of Geophysical Research, v. 99, no. B3, p. 

4341–4354, doi: 10.1029/93JB02993.

Moran, S.C., Malone, S.D., Qamar, A.I., Thelen, W.A., 

Wright, A.K., and Caplan-Auerbach, J., 2008a, Seismicity associated with renewed dome building at Mount St. 

Helens, 2004–2005, chap. 2 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Moran, S.C., McChesney, P.J., and Lockhart, A.B., 2008b, 

Seismicity and infrasound associated with explosions at 

Mount St. Helens, 2004–2005, chap. 6 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Musumeci, C., Gresta, S., and Malone, S.D., 2002, Magma 

system recharge of Mount St. Helens from precise relative 

hypocenter location of microearthquakes: Journal of Geophysical Research, v. 107, no. B10, 2264, p. ESE 16-1–ESE 

16-9, doi:10.1029/2001JB000629.

Reasenberg, P.A., and Oppenheimer, D., 1985, FPFIT, 

FPPLOT, and FPPAGE: FORTRAN computer programs for 

calculating and plotting earthquake fault-plane solutions: 

U.S. Geological Survey Open-File Report 85–739, 109 p.

Rogers, J.A., and Stephens, C.D., 1995, SSAM: Real-time 

seismic spectral amplitude measurement on a PC and its 

application to volcano monitoring: Seismological Society of 

America Bulletin, v. 85, no. 2, p. 632–639.

Thelen, W.A., Crosson, R.S., and Creager, K.C., 2008, 

Absolute and relative locations of earthquakes at Mount St. 

Helens, Washington, using continuous data; implications for 

magmatic processes, chap. 4 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

U.S. Geological Survey, 2006, Earthworm documentation v. 

7.0: U.S. Geological Survey online-only publication, including overview, release notes, and software [http://folkworm.

ceri.memphis.edu/ew-doc/; last accessed Nov. 6, 2006].
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 4

Absolute and Relative Locations of Earthquakes at 

Mount St. Helens, Washington, Using Continuous Data: 

Implications for Magmatic Processes

By Weston A. Thelen1, Robert S. Crosson2, and Kenneth C. Creager2

Abstract

This study uses a combination of absolute and relative locations from earthquake multiplets to investigate the 

seismicity associated with the eruptive sequence at Mount St. 

Helens between September 23, 2004, and November 20, 2004. 

Multiplets, a prominent feature of seismicity during this time 

period, occurred as volcano-tectonic, hybrid, and low-frequency earthquakes spanning a large range of magnitudes and 

lifespans. Absolute locations were improved through the use 

of a new one-dimensional velocity model with excellent shallow constraints on P-wave velocities. We used jackknife tests 

to minimize possible biases in absolute and relative locations 

resulting from station outages and changing station configurations. In this paper, we show that earthquake hypocenters shallowed before the October 1 explosion along a north-dipping 

structure under the 1980–86 dome. Relative relocations of 

multiplets during the initial seismic unrest and ensuing eruption showed rather small source volumes before the October 1 

explosion and larger tabular source volumes after October 5. 

All multiplets possess absolute locations very close to each 

other. However, the highly dissimilar waveforms displayed by 

each of the multiplets analyzed suggest that different sources 

and mechanisms were present within a very small source 

volume. We suggest that multiplets were related to pressurization of the conduit system that produced a stationary source 

that was highly stable over long time periods. On the basis 

of their response to explosions occurring in October 2004, 

earthquakes not associated with multiplets also appeared to be 

1

 Pacific Northwest Seismic Network, Department of Earth and Space Sciences, University of Washington, Box 351310, Seattle, WA 98195

2

 Department of Earth and Space Sciences, University of Washington, Box 

351310, Seattle, WA 98195

pressure dependent. The pressure source for these earthquakes 

appeared, however, to be different from the pressure source of 

the multiplets.

Introduction

Because more than 1 million earthquakes occurred 

between the beginning of the latest eruption of Mount St. 

Helens in 2004 and the end of 2005 (Moran and others, this 

volume, chap. 2), there is a need to locate a representative set 

of earthquakes to search for first-order patterns and trends that 

may aid in future monitoring and research. Although locating 

most of the 1 million earthquakes would be possible and even 

somewhat practical with current methods and parallel computing, such a study is beyond the scope of this paper. Herein we 

concentrate on the seismicity leading up to the early October 

explosions and through about 1 month of dacitic dome building in order to characterize the seismicity associated with each 

eruptive phase.

One of the most striking features in this eruption’s seismicity was the prominence of earthquake multiplets, or repetitious earthquakes. Earthquake multiplets are characteristically 

defined by their highly similar waveforms, often exceeding 

cross-correlation coefficients of 0.8 (Geller and Mueller, 

1980; Frémont and Malone, 1987; fig. 1). This phenomenon is 

likely caused by earthquakes occurring very close to the same 

location with a similar source process (Geller and Mueller, 

1980). Earthquake multiplets are a common occurrence in 

tectonic and volcanic areas worldwide, and many authors have 

exploited earthquake multiplets to determine highly precise 

relative relocations (for example, Poupinet and others, 1984; 

Fréchet, 1985; Frémont and Malone, 1987; Got and others, 

1994; Waldhauser and Ellsworth, 2000). Every relative relocation method operates on the basic idea that closely spaced 
72 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

earthquakes recorded at a common station will have similar 

path and site effects. As long as the source region is homogeneous (or reasonably so) with respect to the dominant wavelength of the waveforms, the relative time lag between the 

earthquakes is a measure of the separation distance between 

the sources (Wolfe, 2002).

The classification and description of large suites of 

multiplets is also fairly common on volcanoes worldwide. Got 

and others (1994) relocated 250 earthquakes beneath Kïlauea, 

Hawai‘i, to define a nearly horizontal plane of seismicity at 

8-km depth. Stevens and Chouet (2001) analyzed long-period 

multiplets beneath Redoubt Volcano before its 1989 eruption 

and found smooth changes in amplitude and cross-correlation 

coefficient within multiplets. Rowe and others (2004) relocated approximately 17,000 similar earthquakes on Soufrière 

Hills volcano, Montserrat, by identifying multiplets through a 

hierarchical clustering algorithm and using those multiplets as 

stacks to refine phase picks. On Mount Pinatubo, Philippines, 

Battaglia and others (2004) used a similar approach to that of 

Got and others (1994) to relocate almost 11,000 events.

When locating similar earthquakes, a common approach 

is to use manually phase-picked events to seed stacks of similar events. Short windows of data are normally used, usually 

between 1.28 and 2.56 s after the P-wave arrival. Often such 

studies use triggered waveform data (records started by an 

amplitude trigger at the seismometer) instead of continuous 

waveform data, which are used here to identify multiplets. Past 

data sets are thus limited by what is triggered or located manually by the local seismic network and may be largely incomplete (Rowe and others, 2004).

Beginning in 1980, Mount St. Helens has been the most 

seismically active volcano in the Cascade Range. Increases in 

moment release in the shallow edifice are characteristically 

associated with explosive and effusive eruptions (Swanson and 

others, 1983). Earthquake multiplets at shallow depths were 

commonly observed at Mount St. Helens after 1982 and were 

used to obtain highly precise relative locations of earthquakes 

during the dome-building eruptions of September 10, 1984, 

and May–June 1985 (Frémont and Malone, 1987). Musumeci 

and others (2002) used similar events during the late 1990s to 

establish continuing magma recharge after 1992 into a storage 

zone at 5.5–10-km depth.

Since the beginning of the latest eruption in September 

2004, less than 1 percent of the earthquakes detected have been 

routinely processed by the Pacific Northwest Seismic Network 

(PNSN) (Qamar and others, this volume, chap. 3), although 

care has been taken by PNSN analysts to extract a representative sampling of waveforms for the PNSN earthquake catalog. 

In this paper we (1) relocate the triggered earthquakes from the 

PNSN catalog to establish absolute locations, (2) identify and 

locate multiplets from the continuous seismic record, and (3) 

discuss the volcanological significance of our findings. We do 

not claim to classify every earthquake that has occurred during 

this period; instead, we present the absolute locations of the 

PNSN catalog data with a new velocity model and the characteristics of a representative set of multiplets that exhibit large 

populations and/or long lifespans. Especially of interest in this 

study are the relative locations and characters of the individual 

multiplets, how those characteristics change through time, and 

the relation of those characteristics to other phenomena. This 

information is important because we would like to assess the 

ability of multiplets to forecast eruptive activity or style.

This paper follows the time divisions used by Moran and 

others (this volume, chap. 2) to present the seismic evolution 

of the eruption. As it happens, station outages or station installations necessitated the use of additional time breaks to present 

consistent locations within a time period. Care has been taken 

across changes in network configurations to present a consistent and coherent picture of the seismicity.

Methods

Absolute Locations

Mount St. Helens has the densest seismic network of any 

volcano in the Pacific Northwest. Before the eruption in 2004, 

there were 13 permanent stations within 20 km of the volcano 

(fig. 2). Except for 11 days after the first phreatic explosion on 

October 1, a seismometer was positioned within 1 km of the 

new vent (McChesney and others, this volume, chap. 7), an 

arrangement that allowed an unusually accurate determination 

of absolute locations beneath the edifice compared to standard 

regional network locations. In this paper, we consider only 

earthquakes with eight picked P-wave arrival times, a maximum azimuthal gap of 135 degrees, and an arrival time pick 

within 3 km of the epicenter. The 3-km criterion is necessary 

for times when stations on the 1980s dome were saturated or 

destroyed. Additionally, we use a new velocity model (Thelen 

and others, 2006) to significantly reduce station corrections 

and improve locations (table 1, appendix 1). Absolute locations are determined with the program “SPONG,” a PNSN 

adaptation of FASTHYPO (Herrmann, 1979).

Uncertainties in absolute event location are given by the 

program SPONG, which was benchmarked against Hypoinverse (Klein, 1989; S. Malone, oral commun., 2006). These 

include uncertainties in phase picks and in the earthquakelocation inversion scheme itself. To evaluate the effect of 

changing station configurations, we performed jackknife tests 

(Efron and Gong, 1983) by systematically removing seismic 

stations. Generally, removing crater stations reduced depth 

constraint but did not change the distribution of epicenters 

significantly (appendix 2). The reduced resolution in depth 

often produced “airquakes,” earthquakes with phase picks that 

resulted in depths above the velocity-model datum used in 

the solution. Airquakes were common in the existing PNSN 

earthquake catalog throughout the eruptive sequence at Mount 

St. Helens beginning in 2004. The uncertainties presented 

here do not include uncertainties in the velocity model, do not 

quantitatively account for changes in station configuration, and 

therefore are likely underestimates, especially in depth.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 73

2 4 6 8 10 12 14 16

Oct. 08, 2004 15:02:09

0.88

0.92

0.94

0.96

0.93

0.91

0.91

0.88

0.87

0.83

0.82

0.95

Multiplet D34: Station HSR

TIME, SECONDS

A

B

C

D

E

F

4 6 8 10 12

Sep. 25, 2004 09:28:13

0.98

0.94

0.96

0.95

0.96

0.95

0.98

0.95

0.94

0.97

0.98

Multiplet A27: Station HSR

TIME, SECONDS

CROSS-CORRELATION VALUE

2 4 6 8 10

Sep. 27, 2004 02:17:22

0.61

0.91

0.98

0.99

0.98

0.90

0.92

0.85

0.80

0.77

0.76

0.91

Multiplet B1: Station HSR

TIME, SECONDS

CROSS-CORRELATION VALUE

8 10 12 14 16 18 20

Oct. 02, 2004 22:08:54

0.84

0.85

0.80

0.85

0.94

0.84

0.84

0.81

0.92

0.85

0.81

0.94

Multiplet C15: Station CDF

TIME, SECONDS

CROSS-CORRELATION VALUE

4 8 12 16 20

Oct. 16, 2004 18:44:30

0.85

0.83

0.86

0.89

0.89

0.89

0.88

0.86

0.90

0.80

0.87

0.94

Multiplet E7: Station HSR

TIME, SECONDS

4 8 12 16 20

Nov. 04, 2004 18:02:13

0.89

0.77

0.89

0.84

0.85

0.80

0.68

0.79

0.75

0.75

0.74

0.85

Multiplet F1: Station HSR

TIME, SECONDS

Figure 1. Example of waveforms from six multiplets analyzed in this study. Top trace in each is seed event for multiplet, 

and last trace is stack of all waveforms within multiplet at given station. Date and time on first trace are when seed event 

occurred. Arrow in bottom trace shows first arrival with polarity (up is compressional, down is dilatational). Station CDF 

has reversed polarity, so the first motion in C is actually down. Cross-correlation value of each trace with respect to seed 

event is shown on the y-axis. Waveforms are unfiltered and chosen at constant time intervals within the lifespan of their 

respective multiplets.
74 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Classification of Multiplets

The PNSN catalog contains less than 1 percent of the total 

number of earthquakes, so entire groups of earthquakes have 

no representative in the catalog. As a result, we could not use 

the PNSN catalog as a starting point in searching for multiplets. Instead, our method for identifying multiplets started with 

a simple, highly sensitive threshold autopicker for which we 

selected a window of data around each trigger on the basis of 

the characteristic length of earthquakes in a given time period 

(10–20 s). Our autopicker typically missed the P-wave first 

arrival, so we included an empirically determined amount of 

time before the trigger (usually 1–3 s) in order to capture the 

entire event.

After the triggers were selected from the record, we crosscorrelated, in the time domain, the first event to all other events 

at a particular station. Any event that was equal to or above the 

threshold cross-correlation coefficient (0.8) was considered 

part of that multiplet. The multiplet was then taken from the 

triggered events and stored. The next event that was not part 

of an existing multiplet was cross-correlated against all the 

remaining events and evaluated for events exceeding the crosscorrelation coefficient cutoff. Any event that did not correlate 

with another event was considered an orphan and stored elseFigure 2. Map of permanent 

Pacific Northwest Seismic 

Network (PNSN) stations 

around Mount St. Helens from 

September to November 2004. 

Modified from Moran and 

others (this volume, chap. 2).

46°15'

46°00'

122°22'30" 122°15' 122°7'30" 122°00'

46°7'30"

46°22'30"

0 5 10 km

WASHINGTON

AREA OF MAP

500 m

SHW

JUN

STD 

CDF

FL2

MTM

TDL

LVP

EDM

SOS

ELK

YEL

SEP

BLIS

NED

HSR
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 75

Depth

(km)

P-wave 

velocity

(km/s)

S-wave 

velocity

(km/s)

0 2.7 1.51

1.0 3.5 1.97

1.3 4.3 2.42

1.6 5.1 2.87

3.9 6.0 3.37

6.5 6.2 3.58

10.5 6.6 3.82

18.5 6.8 3.93

34.5 7.1 4.10

43.5 7.8 4.51

Table 1. Depth and velocity for P-wave 

model of Thelen and others (2006).

[Depth to top of layers defined by P-wave 

velocities; datum is altitude 2,200 m above sea 

level. Column 3 is the velocity model used for 

the relative locations, which assumes a Poisson 

solid in order to convert to S-wave velocity from 

the P-wave velocity of the one-dimensional 

model (appendix 1).]

where. This process was repeated until all events were either 

classified into a multiplet or as an orphan. The multiplets were 

visually inspected to remove those with systematic noise such 

as calibration pulses and telemetry spikes, which can produce 

repeatable waveforms. Computer memory restrictions and the 

large number of triggers allowed us to process only about 10 

days of data at a time. We therefore treated each time period 

individually and then combined the time periods by stacking 

multiplets from each time period and running the same analysis 

described above with event stacks instead of triggered data. 

This method served to combine time periods; however it also 

allowed for events with similar stacks to be combined even if 

individual events did not correlate at the initial threshold level. 

This technique of tracking multiplets has a disadvantage of 

using only one station to identify and classify multiplets, which 

makes it susceptible to near-site effects. However, Stephens and 

Chouet (2001) successfully used only one station to classify the 

multiplets before the 1989 eruption of Redoubt Volcano. We 

also visually checked other stations at the same distance range 

for the presence of the larger multiplets and ensured that nearsite effects were not dominating the multiplet analysis.

Relative Location of Multiplets

After the identification of a multiplet, a representative 

event from that multiplet was chosen and cross-correlated 

against each hour of continuous data. Again, cross-correlation 

windows of 10–20 s were used, a choice based on the length 

of high signal-to-noise ratios throughout the waveform. Crosscorrelation coefficients above a threshold were considered triggers. Three stations were required to have coincident triggers 

in order for the earthquake to be classified as a valid event. 

After the triggered events were compiled, all of the events 

were cross-correlated against each other at each station, and 

lags were noted at the maximum cross-correlation coefficient. 

Subsample estimates of lags were achieved in the time domain 

through polynomial interpolation of the cross-correlation coefficient peak. Lags were weighted by their normalized crosscorrelation coefficient, and only lags with a cross-correlation 

above 0.8 were considered for earthquake locations. The lags 

for each event were then used to compute double-difference 

hypocenter locations with the hypoDD software (Waldhauser, 

2001). For the initial location of the earthquakes within each 

multiplet, we stacked the multiplet and picked the first arrival 

of the stack at each station. Those picks were then applied to 

each earthquake in the stack on the basis of the lag derived 

from the cross-correlation. The picks were used to get absolute 

locations for each earthquake in the multiplet using the identical process as discussed above.

The use of long time windows meant that we were correlating dominantly S waves and Rayleigh waves. Most of 

the stations used in this study had a difference of 2 s or less 

between the P- and S-wave arrivals, meaning that at least 80 

percent of the correlation window was composed of S waves 

and Rayleigh waves. Consequently, cross-correlating long time 

windows produced lags of phases that are traveling at S-wave 

velocities between the source and the receiver. We therefore 

elected to use an S-wave model for our relative locations. The 

S-wave model was obtained by assuming a Poisson’s ratio 

for the crust and applying it to the P-wave model used in the 

absolute locations (table 1).

Errors in relative locations have several sources of uncertainty. One source is waveform alignment. Poorly aligned 

waveforms at a particular station may generate lineations in 

events toward that station (Battaglia and others, 2004). We 

minimized this source of uncertainty by using large time 

windows, which minimizes the occurrence of poorly aligned 

waveforms (Schaff and others, 2004). Another source of error 

in relative locations is uncertainty in the velocity model that 

the initial hypocenters start in (Michelini and Lomax, 2004). 

The path is canceled out by calculating the lag between two 

events, but the takeoff angle and azimuth are affected by the 

choice of an initial velocity model. In our analysis, we use 

a velocity model that is homogenous in the region of our 

cluster centroid. Changes in velocity, and therefore takeoff 

angle, only stretch or compress our cloud of hypocenters and 

do not affect the overall shape (appendix 3). Finally, artifacts 

in relative earthquake locations can be introduced through 

inhomogeneous phase lags on all stations in a solution. In 

other words, some stations will have a larger effect on the 

earthquake location than others, depending on the number of 

phase picks on that station. We tested this with the use of a 

jackknife test (Efron and Gong, 1983), a method in which one 

station is removed at a time and then all of the earthquakes 

within the multiplet are relocated, using the reduced station 

set (appendix 3). The change in location of each earthquake 
76 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

was tracked, and the maximum offset of all of the stations 

in the north, east, and vertical directions is presented as the 

uncertainty in relative location.

September 23–October 5, 2004: VentClearing Phase

Absolute Locations, September 23, 0000 PDT, 

Through September 25, 1200 PDT

Relocating a subset of catalog PNSN events with the 

updated velocity model revealed a tight cluster of earthquakes 

with a radius of 250 m centered immediately southeast of station SEP (fig. 3A). The depth cross section showed a cluster 

of hypocenters that was elongate toward the future vent. The 

cluster spanned depths of approximately 400 to 1,100 m below 

the future vent. The average number of phases used in the 

earthquake location during this time period was 9.32, with 

a maximum of 52. The mean azimuthal gap in this subset of 

earthquake locations was 90 degrees. A total of 229 earthquakes were relocated with formal uncertainties averaging 

132, 92 and 84 m in the north, east, and vertical directions, 

respectively. Reductions in estimated uncertainties for the 

north, east, and vertical directions were 20, 35, and 74 percent 

when compared to the routine network locations.

Multiplets, September 23, 0000 PDT, Through 

September 25, 1200 PDT

The buildup of multiplets occurred at an accelerating 

pace until approximately 0200 on September 24 (fig. 4). All 

multiplets were composed of volcano-tectonic (VT) events 

and, in most cases, many multiplets were occurring simultaneously. Here we distinguish VT earthquakes from hybrid and 

low-frequency earthquakes, as was done by Moran and others 

(this volume, chap. 2). Examples of frequency spectra of 

hybrid versus low-frequency earthquakes versus VT earthquakes, as defined in this paper, are shown in figure 5. The 

maximum number of multiplets with populations greater than 

20 occurring concurrently was eight at 0200 PDT on September 24, coincident with the maximum of earthquake amplitudes on station JUN during this time period (fig. 4). After 

the maximum in earthquake amplitudes on September 24, 

seismic activity waned until September 26. None of the early 

multiplets recurred after seismic activity ramped up again on 

September 26, 2004.

The largest magnitude multiplet during this period, A27 

(figs. 1, 6), was a group of 59 VT events that occurred over 

approximately 3 hours beginning at 1104 PDT on September 

25. Picking the first arrivals of the stacked traces of A27 on 

each station and locating the multiplet with standard absolutelocation techniques resulted in a location at the south end of 

the cluster of epicenters in figure 3A, at depths near the top 

of the cluster. The absolute location of the stack used only six 

phase picks, giving the location of the hypocentroid the largest 

uncertainty of any multiplet analyzed. Double-difference relocations of this swarm reveal a tight cluster, spanning less than 

10 m in depth and about 5 m horizontally (fig. 6A). Orthogonal cross sections show mostly vertical orientations, which are 

an artifact of different stations’ picks being used for different 

events. Uncertainty estimates are on the order of 3 m horizontally and 20 m vertically. In figure 6, the generally small 

average errors are masked by a few larger errors as great as 40 

m. The cross-correlation values of the first event to subsequent 

events in the multiplet are nearly constant. Considering the 

location errors and constant cross-correlation values, we suggest that the earthquakes in this multiplet occurred at nearly 

the same location. The end of this multiplet coincided with the 

waning seismic activity of the swarm on September 25 (Moran 

and others, chap. 2, their figs. 4, 5).

Absolute Locations, September 25, 1200 PDT, 

Through October 1, 1200 PDT

Absolute locations of 313 PNSN catalog events during 

this time period showed a distinct shift up and to the south 

compared to those of the previous period (figs. 3B, 7). The 

decrease in earthquake depth was consistent with phase-pick 

differences identified by Moran and others (this volume, 

chap. 2; fig. 7) between September 25 and September 27. 

This cluster, too, was tightly spaced in depth (250 m radius) 

and in epicenter (300 m radius), with respect to other absolute locations. The hypocentroid shift compared to the swarm 

beginning September 23 was 133 m south, 69 m east, and 221 

m up. The average number of stations used was 11, the average gap was 78 degrees, and the average distance from the 

epicenter to the closest station was 115 m. Uncertainties in 

event locations averaged 109 m north-south, 140 m east-west, 

and 119 m in depth.

Multiplets, September 25, 1200 PDT, Through 

October 1, 1659 PDT

Multiplets were present throughout this time period and 

occurred in two phases before the explosion on October 1 

(fig. 4). The multiplets had lifespans of 1–2.5 days. Their 

first buildup started at approximately 1200 PDT on September 25, 2004. Accompanying an increase in the number and 

lifespan of the multiplets was an increase in the maximum 

amplitudes of the continuous data averaged over 1-minute 

intervals at station JUN (fig. 4). The second buildup started 

around midday on September 29 and continued until the 

explosion on October 1. The second phase is identified by 

the end of multiplets from the first phase of buildup and the 

occurrence of new multiplets. The transition coincided with 

a 250 percent increase in the maximum amplitude of signals 

recorded at station JUN (fig. 4). Thirteen multiplets that 
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 77

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0

0.5

1

1.5

2

2.5

3

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0.5

1

1.5

2

2.5

3

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0.5

1

1.5

2

2.5

3

46.185°

46.2°

46.215°

YEL

SEP

46.185°

46.2°

46.215°

YEL

SEP

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

500 m

YEL

SEP

SEP

SEP

LATITUDE (DEGREES NORTH)

LATITUDE (DEGREES NORTH)

LATITUDE (DEGREES NORTH)

ALTITUDE, IN KILOMETERS ABOVE SEA LEVEL ALTITUDE, IN KILOMETERS ABOVE SEA LEVEL ALTITUDE, IN KILOMETERS ABOVE SEA LEVEL

A

B

C

N

09/23/04–09/25/04

229 EVENTS

09/25/04–10/01/04

313 EVENTS

10/01/04–10/05/04

149 EVENTS

S

S N

S N

0

0

122.205° 122.19° 122.175°

500 m

122.205° 122.19° 122.175°

500 m

Figure 3. Absolute locations for Pacific Northwest Seismic Network (PNSN) catalog events for three time 

periods. Earthquakes plotted have eight or more phase picks, azimuthal gap of less than 135 degrees, and a 

phase pick at a station closer than 3 km from epicenter. Error bars are from SPONG location output. Permanent 

seismic stations shown on maps by black and white triangles. Cross sections show earthquake depth plots and 

topographic surface profile north-south through station SEP. White dot in map view and arrow in cross section 

show approximate location of vent.
78 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

9/23 9/25 9/27 9/29 10/01

100

50

150

200

250

300

350

400

450

2004 DATE, PACIFIC DAYLIGHT TIME

MULTIPLET NUMBER

500

1,000

1,500

2,000

2,500

3,000

3,500

MAXIMUM RMS AMPLITUDE, COUNTS

Explosion

0

Clusters

Hybrid

event

Low-frequency

event

0 5 10 15 20 25

FREQUENCY, IN HERTZ

NORMALIZED AMPLITUDE, IN COUNTS (ARBITRARY SCALE)

0 5 10 15 20 25

0 5 10 15 20 25

High-frequency

event

A

B

C

were active before the October 1 explosion recurred after the 

explosion. Multiplets occurring after September 25 at 1200 

PDT and through November 2004 were overwhelmingly 

hybrid and low-frequency earthquakes.

One multiplet, B1, started September 27 and continued 

for approximately 2.5 days (fig. 6B). The absolute location 

of the stacked traces formed a hypocentroid at 520-m depth, 

with an epicentroid centered in the cluster in figure 3B. The 

absolute location solution used 12 stations and had impulsive 

arrivals at all stations, making it one of the most reliable locations of the multiplets studied. The polarity of first arrivals 

was mixed and was consistent with reverse faulting on a northstriking, vertical or steeply dipping (85o–90o) fault. Relative 

locations of hypocenters are concentrated inside a sphere less 

than 15 m in diameter (fig. 6B). Location errors derived from 

jackknife techniques reveal uncertainties in the north, east, and 

vertical directions of 1.9 m, 1.3 m, and 1.3 m, respectively. 

The cross-correlation coefficient of the first event relative to 

each subsequent event decays with time, indicating a systematic temporal change in either the path between the earthquake 

and the receiver or a very small shift in location (Gret and 

others, 2005).

Absolute Locations, October 1, 1700 PDT, 

Through October 5, 1659 PDT

The absolute locations of 149 PNSN catalog events 

between the October 1 explosion and the end of October 5 

apparently varied more widely in depth than those of the two 

previous time periods (fig. 3). However, with the destruction 

of station SEP during the October 1 explosion, the depths 

Figure 4. Timeline of multiplets from September 24, 2004, until 

explosion of October 1, 2004 (green line), detected on station 

HSR. Each earthquake is plotted as a circle, and multiplets are 

identified by a horizontal line through events that define multiplet. 

Only multiplets that have more than 20 events are plotted. 

Maximum root-mean-squared (RMS) amplitude averaged over 1 

minute at station JUN shown in red.

Figure 5. Examples of frequency spectra. High-frequency 

energy is seen between 12 and 20 Hz in high-frequency 

and hybrid events but not in low-frequency event. A, Highfrequency event from September 25, 2004. B, Hybrid event 

from seed event of multiplet B1. C, Seed event of lowfrequency multiplet F1 on station HSR. 

were less well constrained than in previous earthquake sets. 

In addition, emergent P-wave arrivals and overlapping codas 

made picks at the remaining crater station (YEL) and flank 

stations (HSR, EDM, SHW, and STD) sparse and imprecise, 

further affecting the depths. The effect of this can be seen 

as a cluster of seismicity that was falsely centered at 400-m 
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 79

DISTANCE, METERS

20

0

-20

-40

-60

40

60

20

0

-20

-40

-60

40

60

360

400

440

480

20

0

-20

-40

-60

40

60

DISTANCE, METERS

DEPTH, METERS

0 50

DISTANCE, METERS

DEPTH, METERS

460

500

540

580

DEPTH, METERS

300

340

380

420

DEPTH, METERS

A

B

C

DISTANCE, METERS

-50 0 50

-50

DISTANCE, METERS

-50 0 50

DISTANCE, METERS

450 400 350

400 350 300

DEPTH, METERS

500550 450

DEPTH, METERS

Multiplet B1

238 Events

Average Error X: 1.35

Average Error Y: 1.85

Average Error Z: 1.34

Multiplet A27

59 Events

Average Error X: 3.3

Average Error Y: 3.3

Average Error Z: 18.8

Multiplet C15

404 Events

Average Error X: 4.9

Average Error Y: 5.8

Average Error Z: 5.1

9/23 9/24 9/25 9/26 9/27 9/28 9/29 9/30

0.6

0.8

1.2

1.0

1.4

1.6

1.8

2.2

2.3

2.0

2004 DATE, PACIFIC DAYLIGHT TIME

ALTITUDE, KILOMETERS ABOVE SEA LEVEL

Figure 6. Relative relocations for three multiplets 

from the beginning of unrest through the period of 

phreatic explosions. For each, upper right square is 

map view (north at top), and left and bottom plots are 

cross sections. Plots are presented as cube with two 

sides folded out. Errors determined through jackknife 

techniques. Directions X, Y, and Z correspond to east, 

north, and vertical, respectively.

Figure 7. Earthquake altitude as function of time, September 

23–30, 2004. Black line is running average altitude of 20 nearest 

earthquakes in time. The apparent gap in earthquakes between 9/26 

and 9/28 is artificial, caused by a lack of earthquakes with location 

parameters that met our criteria and not by diminished seismicity.

altitude owing to a lack of picks at station YEL (fig. 3C). 

Earthquake-location solutions used an average of 10 stations 

with an average gap of 81 degrees and an average delta of 

1.52 km. Uncertainties in x, y, and z were 124, 146, and 244 

m, respectively.

Multiplets, October 1, 1700 PDT, Through 

October 5, 1659 PDT

Several small-population multiplets were present during this time period (fig. 8). These multiplets consist of 

hybrid and low-frequency event families. There were three 

explosions and two episodes of tremor during this time 

period, in some cases concurrent with the start or stop of 

a multiplet. In most cases, however, multiplets during this 

time period continued largely unaffected by the occurrence 

of tremor or explosion.

One large low-frequency multiplet, C15, started 7 hours 

after the explosion on October 1. The absolute location of 

this multiplet using event stacks revealed a hypocentroid 

100 m south and 160 m shallower than B1. Using a similar 

station configuration for both event stacks (no station SEP) 

shows approximately the same offset in location. Polarities 

from first arrivals of the multiplet stack were all dilatational, 

a characteristic for other earthquakes in this sequence, especially after October 1 (Moran and others, this volume, chap. 

2; Qamar and others, this volume, chap. 3). This multiplet 

possessed highly stable amplitudes and cross-correlation 

coefficients in time. The relative relocations of 404 events 

from the C15 multiplet show a cluster of seismicity spanning 

approximately 20 m in diameter (fig. 6C). Some events of 
80 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

10/03 10/05 10/07 10/09 10/11

100

50

0

150

200

250

300

350

400

450

500

2004 DATE, PACIFIC DAYLIGHT TIME

MULTIPLET NUMBER

0

500

1,000

1,500

2,000

2,500

3,000

3,500

Explosion

E E EE Tremor

MAXIMUM RMS AMPLITUDE, COUNTS

10/01

Clusters

Figure 8. Timeline of multiplets from explosion of October 

1 (leftmost green line) through October 11, 2004, detected 

on station ELK. Each earthquake is plotted as a circle, and 

multiplets are identified by a horizontal line through events that 

define multiplet. Only multiplets having more than 20 events 

are plotted. Maximum root-mean-squared (RMS) amplitude 

averaged over 1 minute at station JUN shown in red. Explosions, 

vertical green lines; tremor episodes, vertical gray lines.

the main cluster may have indicated a south-dipping structure; however, the overwhelming majority (75 percent) of 

the earthquake hypocenters were within the circular cluster. 

Using jackknife error techniques, we determined the locational uncertainties in north, east, and depth to be 5.8, 4.9, and 

5.1 m, respectively.

October 5–November 20, 2004: 

Seismicity Associated with Dome 

Building

Absolute Locations, October 5, 0000 PDT, 

Through October 11, 1659 PDT

As with those between October 1 and October 5, 

earthquake arrivals at crater and flank stations were highly 

contaminated with earthquake codas from previous earthquakes and plagued with emergent signals. Thus, large 

uncertainties existed in the absolute locations, particularly in 

depth (fig. 9A). The spread of earthquake depth is suspect, 

owing to the absence of station SEP and the poor character of 

the first arrivals on station YEL. During this time period the 

absolute-location solutions of 227 events used an average of 

10 stations, with an average gap of 89 degrees. The average 

distance to the nearest station was 1.15 km. Average uncertainties were 142 m in the north-south direction, 130 m in the 

east-west direction, and 226 m in depth.

Multiplets, October 5, 0000 PDT, Through 

October 11, 1659 PDT

Although four multiplets continued briefly through the 

explosion on October 5, a new temporal group of multiplets 

began about 1200 PDT on October 5. These multiplets generally had longer lifespans than previous multiplets, and fewer 

multiplets were occurring at the same time. Hybrids and 

low-frequency earthquakes dominated the seismic record. 

Multiplets during this time period did not obviously correspond to changes in the maximum amplitudes of signals 

recorded on station JUN.

The D34 hybrid multiplet discussed by Moran and others 

(this volume, chap. 2, their fig. 11) occurred during this time 

period. This multiplet consisted of 681 events starting October 

8 at 0200 PDT. Absolute locations determined from the first 

arrivals of event stacks revealed a hypocentroid in the same 

location as multiplet C15. As was the case for C15, the polarities of the first arrivals were all dilatational. This multiplet 

methodically increased its amplitude before reaching a plateau 

and then decreased again to near the initial amplitude level 

(Moran and others, this volume, chap. 2). Magnitudes at the 

beginning were in the range Md

 1–1.5 and at the maximum 

were Md 2.5–3.0. The cross-correlation coefficient of the first 

event to all others declined until the plateau in amplitude, and 

then remained constant for the rest of the sequence. Relative 

relocations of this cluster reveal a different pattern from that 

in previous multiplets. The hypocenter volume was tabular, 

striking to the southwest (fig. 10A). Uncertainties derived from 

jackknife tests showed errors of 2.7, 2.3, and 2.5 m, in the 

north-south, east-west, and vertical directions, respectively.

Absolute Locations, October 11, 1700 PDT, 

Through November 4, 1559 PST

During this period, low-frequency earthquakes dominated 

the seismic record. These earthquakes possessed emergent 

arrivals and thus had high uncertainties in absolute locations 

(fig. 9B). Earthquakes were extremely difficult to pick at 

stations outside the crater because of coda interference from 

previous events and poor arrivals. On October 12, station BLIS 

was installed on the saddle between the old and new domes 

(McChesney and others, this volume, chap. 7; fig. 2 for location). This station was used in less than half of the locations, 

however, owing to consistently emergent arrivals. Over the 

period of a month, only 130 earthquakes located by the PNSN 

met our selection criteria. For these earthquakes, an average of 

10 stations were used for location solutions, with an average 

gap of 96 degrees and an average distance to the closest station of 431 m. Uncertainties in the north-south, east-west, and 

vertical directions were 142, 130, and 227 m, respectively.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 81

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0.5

1

1.5

2

2.5

3

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

YEL

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0.5

1

1.5

2

2.5

3

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

YEL

46.185 46.19 46.195 46.2 46.205 46.21 46.215

0.5

1

1.5

2

2.5

3

ALTITUDE ABOVE SEA LEVEL, KILOMETERS

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

0.5 km

0.5 km

SEP

SEP

ALTITUDE ABOVE SEA LEVEL, KILOMETERS

LATITUDE (DEGREES NORTH)

LATITUDE (DEGREES NORTH)

LATITUDE (DEGREES NORTH)

ALTITUDE ABOVE SEA LEVEL, KILOMETERS

A

B

C

N

N

10/05/04–10/11/04

227 EVENTS

10/11/04–11/04/04

130 EVENTS

11/04/04–11/22/04

286 EVENTS

YEL

BLIS

SEP

0.5 km

N

0

0

0

S N

S N

S N

Figure 9. Absolute locations for Pacific Northwest Seismic Network (PNSN) catalog events for given time periods. 

Map views shown on left, cross sections on right. Permanent PNSN seismic stations shown with white triangles. Cross 

sections are north-south through station SEP, which was destroyed by explosion on October 1 and not reinstalled until 

November 5, 2004 (shown in top two plots only for reference). Earthquakes selected using same criteria as in figure 3. 

White dot (map view) and black arrow (cross section) show approximate location of vent.
82 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

50

100

150

DISTANCE, METERS

DEPTH, METERS

0 100

DISTANCE, METERS

DEPTH, METERS

0

20

40

60

DISTANCE, METERS

DEPTH, METERS

0 50

DISTANCE, METERS

DEPTH, METERS

0

20

40

60

DISTANCE, METERS

DEPTH, METERS

0 50

DISTANCE, METERS

DEPTH, METERS

300

320

340

360

380

400

420

-50

400 350 300

-20

-40

-60

-20

-40

-60

180

200

220

240

260

280

300

320

300 250 200

-50

 -50

-100

-150 150

200

250

300

350

400

350 250 150

-100

Multiplet D34

681 Events

Average Error X: 2.3 m

Average Error Y: 2.7 m

Average Error Z: 2.5 m

Multiplet E7

753 Events

Average Error X: 5.1 m

Average Error Y: 4.0 m

Average Error Z: 3.8 m

Multiplet F1

3269 Events

Average Error X: 15.1 m

Average Error Y: 7.0 m

Average Error Z: 10.2 m

A

B

C

Multiplets, October 11, 1700 PDT, Through 

November 4, 1559 PST

October 11, 2004, marked the first time new lava was 

seen at the surface (Vallance and others, this volume, chap. 

9). Between then and November 4, three different spines were 

extruded, accompanied by comparatively low seismicity rates. 

Multiplets in the beginning of this time period had lifespans 

comparable to multiplets beginning after the October 5 explosion (figs. 8, 11), but multiplets with lifespans more than 1 

month long began to occur later in the study period. In general, 

multiplets starting after October 22, 2004, had substantially 

longer lifespans than earlier multiplets. However, there was no 

obvious correlation between spine extrusion and the occurrence of multiplets during this time period.

One multiplet, E7, consisted of 753 events and started 

on October 16. Peak-to-peak amplitudes were stable and the 

cross-correlation coefficients of the first event compared to all 

others were nearly constant. The absolute hypocentroid from 

first-arrival picks of event stacks resulted in a location similar 

to that of multiplets C15 and D34. The polarities of first arrivals from event stacks all showed dilatation. The distribution 

of hypocenters was tabular, like that of D34, but had a more 

westerly strike (fig. 10B). The tabular distribution of epicenters had a long axis of approximately 40 m and a short axis of 

10 m or less. The relative hypocenters spanned depths from 

240 to 260 m. Uncertainties derived from jackknife analyses 

revealed average errors in location of 4.0, 5.1, and 3.8 m in the 

north-south, east-west, and vertical directions, respectively.

Absolute Locations, November 4, 1600 PST, 

Through November 22, 1559 PST

On November 5, station SEP was reinstalled (McChesney 

and others, this volume, chap. 7), increasing to three the number 

of stations within 2 km of the actively extruding spine. This 

addition proved important for locating earthquakes, as there 

were reliable picks from three close-in stations for almost all 

events. Absolute locations for 286 PNSN catalog events showed 

a 300-m-radius cluster with an epicentroid located only 160 m 

from the active vent (fig. 9C). The hypocentroid was located 

less than 400 m below the surface. The addition of the closest 

station, BLIS, on October 12, improved the sensitivity of the 

network to earthquakes, but comparing locations determined 

with and without BLIS is problematic. By ignoring station 

BLIS, we simulated the station configuration before October 1, 

2004. The effect of taking out station BLIS was to make events 

slightly deeper on average, but the epicenters remained nearly 

unchanged. Details of these tests can be found in appendix 2. 

These tests suggest that the hypocentroid during this time period 

was at the same depth as the hypocentroid for the time period 

between September 25 and October 1, 2004, within error (figs. 

3B, 6C). During the time period between November 4 and 

November 22, 2004, the earthquake-location solutions used an 

average of 10 stations, with an average gap of 96 degrees and 

Figure 10. Relative relocations for three multiplets from early 

October to late November 2004. See figure 6 for explanation of 

views and errors.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 83

100

0

200

300

400

500

600

700

800

900

MAXIMUM RMS AMPLITUDE, COUNTS

10/12 10/17 10/22 10/27 11/01 11/06 11/11 11/16 11/20

100

50

0

150

200

250

300

2004 DATE, PACIFIC DAYLIGHT TIME

MULTIPLET NUMBER

Spine

1

Spine

2

Spine

3

Clusters

Figure 11. Timeline of multiplets detected on station ELK from October 12 through November 20, 2004. Each earthquake is 

plotted as a circle, and multiplets are identified by a horizontal line through events that define multiplet. Only multiplets that 

have more than 20 events are plotted. Maximum root-mean-squared (RMS) amplitude averaged over 1-minute intervals 

at station JUN shown in red. Date tickmarks are beginning of day, 0000 hr (midnight). Note change in amplitude scale 

compared to figures 4 and 8. Bar on top shows dome extrusion chronology (from Vallance and others, this volume, chap. 9).

an average distance from the epicenter to the nearest station of 

45 m. Average uncertainties in the north-south, east-west, and 

vertical directions were 134, 107, and 122 m.

Multiplets, November 4, 1600 PST, Through 

November 22, 1559 PST

As was the case for the previous time period, multiplets 

during this time period possessed lengthy lifespans that included 

long periods of quiescence (fig. 11). As before, many multiplets 

were concurrent. One of the most prominent multiplets during 

this period, F1, consisted of 3,269 events that began on November 4 at approximately 1700 PST and lasted for more than 20 

days. The absolute location derived from first-arrival picks of 

earthquake stacks resulted in a location south of all other multiplets analyzed in this study. The emergent first arrivals in the 

earthquake stacks produced large pick uncertainties. Resulting 

large uncertainties in the absolute locations as a result of the 

large pick uncertainties means that we cannot distinguish the 

absolute hypocentroid of F1 from those of other multiplet locations. Relative relocations of F1 formed a tabular volume similar 

in shape to multiplet E7 (Oct. 16 start) (fig. 10B). The epicentral 

distribution of multiplet F1 had a long axis of approximately 

100 m and a short axis of 20 m. The cluster was confined to 

depths of 260 to 300 m. Average uncertainties derived from 

jackknife analyses in the north-south, east-west, and vertical 

directions were 7.0, 15.1, and 10.2 m, respectively.

Discussion

Absolute Locations

The high-frequency nature of the earthquakes between September 23 and September 25, 2004, suggests that the earthquakes 

during this time period were fracturing brittle rock. Although 

the uncertainties in latitude and longitude are on the same order 

of magnitude as the epicentral shift between September 23 and 

September 25, we believe the phase pick differences outlined by 

Moran and others (this volume, chap. 2) support the observed shift 

in epicentroid location. The upward and southward migration of 
84 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

absolute hypocenters between September 24 and September 26 

and the shift to dominantly hybrid earthquakes suggest that the 

seismicity was occurring in response to a magmatic plug moving 

toward the surface (fig. 7). The shift occurred before any surface 

deformation near the 1980–86 dome (Dzurisin and others, this volume, chap. 14), although the continuous GPS at Johnston Ridge 

Observatory (JRO) did detect deflationary movement toward the 

volcano (Lisowski and others, this volume, chap. 15). An upward 

shift in hypocenters in the 1989 Ito-oki submarine eruption was 

explained as an upward propagating dike (Ukawa and Tsukahara, 

1996). Their model showed that seismicity could occur above the 

crack tip in response to the dilatation of the crack below. Although 

this conclusion is plausible in hindsight at Mount St. Helens, the 

character and absolute locations of the earthquakes in the swarm 

between September 23 and September 25 alone probably would 

have not been enough to forecast an impending eruption.

After the first phreatic explosion on October 1, 2004, 

the loss of station SEP and consistently high seismic amplitudes on station YEL prevented the precise picking of events 

on crater stations. The absolute locations thus show greater 

scatter, have larger uncertainties, and are skewed in depth 

owing to the lessened station coverage. Jackknife tests performed by removing stations SEP and YEL from earthquake 

locations from September 25 to October 1 reveal depth patterns in seismicity similar to those seen for absolute earthquake locations calculated between October 1 and November 

4. In essence, the precision in depth was poorer with the 

absence of stations SEP and YEL. This made relatively deep 

earthquakes deeper and shallow earthquakes become shallower (appendix 2). The locations were also more vertically 

aligned, similar to the absolute locations from October 1 

through October 5 (fig. 3). The epicenter distribution in 

jackknife tests was nearly unchanged. We therefore suggest 

that the absolute locations for earthquakes from October 1 

through October 5 were similar to those occurring before 

the first explosion. Similar tests indicate that locations for 

earthquakes after the October 5 explosion were similar to 

locations between September 25 and October 1 (appendix 2).

Absolute locations from November 5 to November 22 

show that earthquake locations shifted slightly both up and 

south compared to pre-October 1 seismicity. The character of 

seismicity did not change significantly between October 16 and 

November 5, and we suggest that seismicity after the middle 

of October had similar absolute locations. The hypocentroid of 

absolute locations from November 5 to November 22 was ~290 

m below the surface. Patterns in absolute locations between 

September 23 and November 22 indicate a conduit plunging 

45o–60o northward at shallow depth beneath the old dome. 

The shape of the pattern is robust to changes in station layout 

(appendix 2); therefore we are confident the pattern is real.

Multiplets

A striking feature of seismicity during the early months of 

seismic unrest and eruption was the overwhelming presence of 

multiplets. Discrete multiplets occurred simultaneously throughout the study period; however, many can be further grouped into 

multiplet clusters. Here, we define a multiplet cluster as several 

multiplets that occurred simultaneously. Multiplet clusters often 

build slowly and end abruptly (figs. 4, 8, 11). On September 25, 

September 29, October 1, and October 5, the abrupt end of one 

cluster and the start of another was accompanied by changes in 

the amplitude of seismic signals recorded at station JUN (figs. 4, 

8). No cluster recurred at any point during the study period, but 

in at least one case a multiplet within a previous cluster recurred 

after the cluster had ceased. The distinct clustering in time of 

multiplets may indicate that a particular source region was 

active during a cluster, with each multiplet having a different 

mechanism or occurring at a different location within the cluster 

source region. As each source region was altered or destroyed, a 

new cluster began from a new source region. Temporal clusters 

of multiplets were also observed at Soufrière Hills volcano, 

Montserrat, and were found to be related to the time derivative 

of the tilt (Green and Neuberg, 2006). Peaks in the real-time 

seismic amplitude measurement (RSAM), similar to the maximum amplitudes used here, were well correlated with peaks in 

the tilt at Soufrière Hills volcano; however, the relation between 

the RSAM and the occurrence of clusters of multiplets was not 

explicitly stated.

At Mount St. Helens, the absolute hypocentroid of 

each analyzed multiplet was similar when determined from 

stacks of each multiplet recorded at each station. All of the 

analyzed multiplets had locations within ~100 m of station BLIS at depths of 250 to 350 m (fig. 12). Further, the 

cross-correlation coefficients between each of the analyzed 

multiplets had values ranging between 0.17 and 0.50. This 

suggests that highly stable but distinct sources were present 

within a small source region at shallow depth beneath the 

1980–86 dome. Similarly, the absolute location of multiplets 

that occurred at Soufrière Hills volcano in 1997 clustered 

tightly together, and Neuberg and others (2006) used this 

evidence to help argue for a pressure-dependent source for 

the occurrence of multiplets.

Multiplets also varied in lifespan. Generally, multiplets 

earlier in the study period had shorter lifespans than those 

occurring later. Lifespans of more than 1 month were observed 

for some multiplets during the dome-building phase of the 

eruption, suggesting a stable and stationary source region. 

Such long event families have only been reported at Shishaldin 

Volcano, Alaska (Petersen and others, 2006). However, very 

few studies have covered periods of time longer than days 

to weeks with continuous data. Amplitudes of events within 

individual multiplets were highly variable. However, changes 

typically varied smoothly with time. The magnitudes of events 

in large-population multiplets also ranged widely, from Md

–0.2 to Md 3.0.

A steady decline in the cross-correlation coefficient 

between the first and subsequent events is observed in many 

multiplets. An example of this phenomenon is multiplet B1 

(fig. 13). This change in cross-correlation coefficient could 

result from a shift in location or a change in the path between 
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 85

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

400 m

A27

B1

C15

E7

SEP F1

YEL

BLIS

46.185° 46.195° 46.205° 46.215°

1

0.5

0

1.5

2

3

S N

2.5

LATITUDE

ALTITUDE, IN KILOMETERS ABOVE SEA LEVEL

A27

E7 C15

D34

B1

F1

A

B

D34

an earthquake and the recording station. Gret and others (2005) 

exploited the sensitivity of the coda to changes in path to 

distinguish changes in location from changes in path at Mount 

Erebus, Antarctica. They compared the cross-correlation values 

of event families using windows in different parts of the earthquake waveform. A similar analysis of multiplets here reveals 

similar patterns in correlation coefficients of the first event to 

each subsequent event, regardless of the time window used 

(fig. 13). We therefore interpret the change in cross-correlation 

values with time to result from small changes in location that 

are within the uncertainty of our relative locations.

0 2 4 6 8 10 12

TIME, SECONDS

CROSS-CORRELATION COEFFICIENT

Multiplet B1, Station HSR

PACIFIC DAYLIGHT TIME

0

0.5

1

Station SEP

Early

Late

Full

0.6

0.7

0.8

0.9

1 Station HSR

Early

Late

Full

0.8

0.85

0.9

0.95

1

Station SHW

Early

Late

Full

1600 2300 600 1300 2000 0300

1600 2300 600 1300 2000 0300

1600 2300 600 1300 2000 0300

9/26/04 9/27/04

B

C

D

A

Figure 12. Absolute locations, obtained using first-arrival picks 

of event stacks, for multiplets depicted in figures 6 and 10. A, Map 

view. Blue triangles, permanent Pacific Northwest Seismic Network 

(PNSN) seismic stations as of November 4, 2004. B, Locations 

projected on north-south cross section through station SEP.

Figure 13. Cross-correlation analysis of first 220 events 

from multiplet B1. A, Waveforms filtered between 1 and 10 

Hz of every tenth earthquake recorded on station HSR. B–D, 

Cross-correlation values as a function of time at stations 

SEP, HSR, and SHW, respectively. Three parts of waveform 

are used in each plot: first 4 s (early), next 4 s (late), and 

entire waveform (full).
86 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Only the explosion on October 1 and, to a lesser extent, 

the explosion on October 5 significantly affected the occurrence of multiplets. After the October 1 explosion, all earthquakes, including multiplets, stopped for more than 3 hours 

(Moran and others, this volume, chap. 2). Several multiplets 

then restarted, and new multiplets began as the level of seismicity increased. This process repeated itself during periods of 

tremor and explosions occurring between October 1 and October 5; some multiplets already present continued and, in many 

cases, new multiplets began. The reactivation of multiplets and 

general seismicity after the October 1 explosion suggests that 

the phenomenon leading to the generation of multiplets during 

the time period around October 1 may be pressure dependent 

(Moran and others, this volume, chap. 2). The relation between 

the October 5 explosion and the occurrence of multiplets 

is less clear, because four multiplets continued through the 

explosion and another multiplet cluster started shortly after.

Relative locations of multiplets during the period we 

studied define a changing source volume. Between September 23 and October 5, hypocenters of multiplets defined a 

small equidimensional space, but after October 5 they defined 

larger, more tabular bodies striking west-southwest. In all 

cases except multiplet F1, depths spanned 20 m or less. The 

strike of the tabular bodies in E7 and F1 was also approximately perpendicular to the strike of spine 3, which extruded 

from October 24 to December 20, 2004 (Vallance and others, 

this volume, chap. 9). Such a progression is interesting in the 

context of volcanic activity. Moran and others (this volume, 

chap. 2) suggest that seismicity between September 23 and 

October 5 was associated with clearing of the conduit to allow 

subsequent lava extrusion. Smaller, more equidimensional 

source volumes early in the study period have relatively 

smaller surface area for gas escape. If the gas could not escape 

efficiently, excess overpressure may have caused fragmentation to occur in the conduit, resulting in the explosions seen 

between October 1 and October 5. After October 5, seismic 

amplitudes decreased, the lifespan of multiplets and source 

volumes increased, and extrusion of lava at the surface soon 

began. The tabular shape and orientation of the relative locations after October 1 are suggestive of a dike. Presumably, the 

larger source volume (and larger surface area) defined by the 

relative locations facilitated greater gas loss at shallow levels 

in the conduit, thereby contributing to steady degassing and 

suppression of explosions.

The cause for earthquakes at the restricted depths 

indicated by the absolute and relative locations is enigmatic. 

Seismogenic stresses around a plug of magma are dependent 

on the diameter, velocity, and viscosity of the plug (Webb 

and Dingwell, 1990; Goto, 1999). Indeed, it has been shown 

that stick-slip motion along the edges of an extruding plug 

can generate earthquakes that match the rates and sizes of 

observed drumbeat earthquakes at Mount St. Helens (Iverson, 

this volume, chap. 21). We might expect seismogenic zones 

at constrictions in the conduit (where the local magma flux 

increases), and we believe that this may be a reasonable source 

of seismogenic stress. We might also expect earthquakes in 

an area where the overburden pressure is low enough to allow 

solidification of the magma and force the viscosity across 

some threshold value. The viscosity is dependent mostly on 

pressure and, to a lesser extent, on temperature (Blundy and 

others, 2006). Viscosity can vary over many orders of magnitude, making it a difficult parameter to model in a context useful for interpreting the location of earthquakes (Goto, 1999). 

The lithostatic pressure, however, increased only slightly in 

the time period we describe. A decrease in the occurrence of 

earthquakes and multiplets after the October 1 explosion also 

suggests that the seismicity was dependent on pressure. We 

would therefore expect earthquake locations to remain fixed or 

to shallow slightly during the study period if the earthquakes 

were dependent on the lithostatic pressure, and the viscosity 

and velocity of the plug were fixed. The absolute depths of 

the multiplets were consistent with this observation and were 

also coincident with the depth of microlite crystallization of 

the rising magma (Cashman and others, this volume, chap. 

19; Pallister and others, this volume, chap. 30). The depth of 

microlite crystallization is also thought to be a pressure-dependent phenomenon.

Not all earthquakes during the study period occurred 

as part of a multiplet. Indeed, the absolute locations show 

that earthquake hypocenters were spread across a wide range 

of depths, exceeding the depth range of the multiplets. The 

seismic energy declined after each explosion (Moran and 

others, this volume, chap. 6, their fig. 3), but for many of the 

explosions the multiplets persisted (fig. 8). This indicates that 

the earthquakes not associated with the multiplets declined in 

size and or number and therefore may also have been pressure 

dependent. Because the behavior of those earthquakes associated with multiplets and of those not associated with multiplets differs, we suggest that the cause of each phenomenon 

is distinct. In the case of the multiplets, the source is highly 

stable and nondestructive for long periods of time. In the case 

of earthquakes not associated with a multiplet, the source is 

nonrepeatable over short time scales.

Conclusions

Absolute and relative locations allowed us to characterize 

the possible source of earthquakes associated with renewed 

activity at Mount St. Helens beginning in September 2004. 

Absolute locations were improved with the use of a new velocity model. During periods of good depth constraint, absolute 

locations were found to be less than 1 km deep. A shallowing of earthquakes occurred between September 24 and 26, 

supported by changes in the difference of first-arrival picks 

at close-in stations (Moran and others, this volume, chap. 2). 

Absolute earthquake locations revealed a north-dipping structure which likely coincides with the shallow conduit that was 

formed during the initial vent-clearing phase.

Multiplets were prominent during the first months of the 

eruption. They included volcano-tectonic, hybrid, and low-
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 87

frequency events with a wide range of magnitudes. Multiplets 

varied in lifespan, but those of the vent-clearing phase had 

much shorter lifespans than multiplets occurring during dome 

extrusion. Multiplets also occurred in temporal clusters that 

were usually associated with large changes in the maximum 

RMS seismic-signal amplitude. Absolute locations derived 

from earthquake stacks show that the individual multiplets 

likely originated in nearly the same area. Relative locations 

showed source volumes with diameters on the order of 25 m 

before October 5 and larger, tabular source volumes oriented 

to the west-southwest thereafter, a change concurrent with the 

transition from vent clearing to dome building. In some cases, 

the cross-correlation coefficients of the first event compared 

to all subsequent events showed a steady decline, suggesting 

movement of the source location at a level within our uncertainty. The properties and locations of multiplets suggest that 

their occurrence was based on changes in pressure.

Acknowledgments

We wish to thank the staff at the Pacific Northwest 

Seismic Network for their meticulous work in collecting and 

archiving data and generating a consistent catalog throughout 

the volcanic sequence—notably Steve Malone, Tony Qamar, 

Amy Wright, Tom Yelin, Guy Medema, Karen Meagher, Bill 

Steele, Bob Norris, and Ruth Ludwin. We also wish to thank 

our reviewers, Seth Moran and Clifford Thurber, for insightful 

comments that significantly improved the final version of the 

manuscript, and Dave Ramsey for the map in figure 2.

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growth behavior at Soufriere Hills Volcano, Montserrat, 

revealed by relocation of volcanic event swarms, 1995–

1996: Journal of Volcanology and Geothermal Research, v. 

134, no. 3, p. 199–221.

Schaff, D.P., Bokelmann, G.H.R., Ellsworth, W.L., Zanzerkia, 

E., Waldhauser, F., and Beroza, G.C., 2004, Optimizing 

correlation techniques for improved earthquake location: 

Bulletin of the Seismological Society of America, v. 94, no. 

2, p. 705–721.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 89

Stephens, C.D., and Chouet, B.A., 2001, Evolution of the 

December 14, 1989 precursory long-period event swarm 

at Redoubt Volcano, Alaska: Journal of Volcanology and 

Geothermal Research, v. 109, p. 133–148.

Swanson, D.A., Casadevall, T.J., Dzurisin, D., Malone, S.D., 

Newhall, C.G., and Weaver, C.S., 1983, Predicting eruptions 

at Mount St. Helens, June 1980 through December 1982: 

Science, v. 221, no. 4618, p. 1369–1376.

Thelen, W.A., Malone, S.D., Qamar, A.I., and Pullammanappallil, S., 2006, Improvements to absolute locations from an 

updated velocity model at Mount St. Helens, Washington 

[abs.]: Seismological Research Letters, v. 77, no. 2, p. 238.

Ukawa, M., and Tsukahara, H., 1996, Earthquake swarms and 

dike intrusions off the east coast of Izu Peninsula, central 

Japan: Tectonophysics, v. 253, p. 285–303.

Vallance, J.W., Schneider, D.J., and Schilling, S.P., 2008, 

Growth of the 2004–2006 lava-dome complex at Mount St. 

Helens, Washington, chap. 9 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Waite, G.P., and Moran, S.C., 2006, Crustal P-wave speed 

structure under Mount St. Helens from local earthquake 

tomography [abs.]: Eos (American Geophysical Union 

Transactions), v. 87, no. 52, Fall Meeting supplement, 

Abstract V11B-0578.

Waldhauser, F., 2001, HypoDD—a program to compute 

double-difference hypocenter locations (hypoDD version 

1.0): United States Geologic Survey Open-File Report 

01–113, 25 p.

Waldhauser, F., and Ellsworth, W.L., 2000, A double-difference earthquake location algorithm; method and application 

to the northern Hayward Fault, California: Bulletin of the 

Seismological Society of America, v. 90, p. 1353–1368.

Webb, S.L., and Dingwell, D.B., 1990, Non-Newtonian rheology of igneous melts at high stress and strain rates; experimental results for rhyolite, andesite, basalt, and nephelinite: 

Journal of Geophysical Research, v. 95, p. 15695–15701.

Williams, D.L., Abrams, G., Finn, C., Dzurisin, D., Johnson, 

D.J., and Denlinger, R., 1987, Evidence from gravity data 

for an intrusive complex beneath Mount St. Helens: Journal 

of Geophysical Research, v. 92, no. B10, p. 10207–10222.

Wolfe, C.J., 2002, On the mathematics of using difference 

operators to relocate earthquakes: Bulletin of the Seismological Society of America, v. 92, p. 2879–2892.
90 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Velocity Model

A velocity model serves an integral role in earthquake 

locations. Although many regions have three-dimensional 

velocity models available, computing power and speed 

requirements often mean that one-dimensional velocity 

models are used for routine locations. At Mount St. Helens, 

large static station corrections and earthquakes with locations 

above the velocity model datum (airquakes) were common during the beginning of the 2004 unrest and eruption, 

which demonstrated the need for a more accurate velocity 

model in the shallow part of the volcano. The velocity model 

used for routine locations by the Pacific Northwest Seismic 

Network (PNSN) was determined using nearby quarry blasts 

(Malone and Pavlis, 1983). That velocity model is regionally accurate, judging from the small station corrections on 

stations beyond the edifice of Mount St. Helens. The data 

used for the velocity model, however, had poor ray coverage 

through the shallow part of the edifice (<2 km). The result is 

a constant P-wave velocity of 4.6 km/s in the upper 2.2 km 

of the velocity model, far too fast for the pyroclastic deposits, thin lava flows, and fractured dacite lava domes found 

throughout the crater walls of the volcano. Exceedingly fast 

velocities resulted in large station corrections on edifice and 

crater stations used for routine processing (table 2). Threedimensional P-wave velocity models exist for Mount St. 

Helens (Lees, 1989; Waite and Moran, 2006), but they also 

suffer from a lack of ray coverage and coarse resolution at 

the depths we are studying. Here we describe the method we 

used to develop a shallow (<2 km) one-dimensional P-wave 

velocity model to supplement the existing one-dimensional 

model (in use by the PNSN) for the Mount St. Helens area.

Methods

In the summer of 2005, we deployed, at 100-m intervals, 39 RT-125 seismic recorders with alternating 4.5-Hz 

geophones and 1-Hz vertical seismometers (Thelen and others, 2006). The instruments were deployed for three days. 

The array began approximately 1 km north of the 2004–5 

vent and extended northward 4 km onto the Pumice Plain 

(fig. 14).

We used an Occam’s inversion scheme with second-order 

Tikhonov regularization to invert first-arrival data for a velocity model (Aster and others, 2005). The Occam’s inversion 

scheme seeks a solution to G(m) = d subject to the following 

constraints:

G(m) d

Lmmin

2

2

and − ≤  

PNSN Station Existing PNSN 

Model1

 (s)

New Velocity 

Model1 (s)

New Velocity 

Model2 (s)

SEND -0.087 0.000 -0.023

RAFT -0.187 -0.083 0.063

MIBL -0.218 -0.096 0.000

WESG -0.179 -0.091 0.099

MIDE -0.212 -0.190 0.000

SEP -0.151 -0.102 0.071

BLIS -0.393 -0.366 -0.170

NED -0.050 0.000 0.131

SUG 0.200 0.000 0.064

YEL 0.024 0.000 0.078

EDM 0.179 0.059 0.033

HSR 0.192 0.070 0.088

SHW 0.108 0.042 -0.032

STD 0.070 0.057 0.000

SOS 0.017 0.039 0.000

JUN -0.029 0.025 -0.034

FL2 -0.029 0.054 0.000

CDF -0.043 0.134 0.062

ELK -0.226 -0.177 -0.227

TDL -0.163 -0.158 -0.223

LVP -0.078 0.000 -0.060

1

Datum is 1,700 m above sea level.

2

Datum is 2,200 m above sea level.

Table 2. Station corrections calculated for each velocity 

model with permanent Pacific Northwest Seismic Network 

(PNSN) stations used in routine earthquake locations between 

2004 and 2006.

[See McChesney and others (this volume, chap. 7) for locations of seismic 

stations not shown in figure 2. Data shown here for each station are time 

residuals, in seconds, that result from the inadequacy of a one-dimensional 

model to represent the three-dimensional Earth. In theory, station residuals 

converge to zero as the computed velocity model converges toward the actual 

velocity structure between earthquake and station.]

where G(m) is a function of the model parameters that calculates the data, d, from the model. In this formulation, L is 

the roughening matrix (which is equal to the finite-difference 

approximation of the second derivative), and  is the discrepancy parameter, n , where n is the number of degrees of 

freedom and r is the uncertainty of the time picks. Occam’s 

inversion relies on an iteratively applied local linearization 

given by Taylor’s theorem,

mmJmGmmG k k k)()()( ∆+≈∆+ ,

where mk is a trial model and J(mk) is the Jacobian. Using a 

damped least-squares approach, the constraints above, combined with the Taylor approximation, give

2

2

2 2

2 )()(min mmLdmmJmG )( k k k α ∆++−∆+ ,

where the parameter a is adjusted after each iteration so that 

the solution will not exceed the allowable misfit. The inversion 

is terminated when d 2=v 2, where
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 91

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

46.23°

46.245°

500 m

))((

1

2

2

2 2

∑=

− =

m

i i

i Gmd iL

σ

χ .

Our “active sources” were earthquakes associated with 

dome building at Mount St. Helens. Twenty Md > 1.0 earthquakes occurred within the crater from the surface to depths 

of 1 km while our instruments were recording. Of these, we 

chose the best earthquake (August 20, 2005, 21:28:37 PDT; 

Md = 2.6) on the basis of its clarity of impulsive arrivals across 

the transect and the number of PNSN catalog picks used to 

routinely locate the earthquake. The epicenter of the earthquake is well constrained, owing to the excellent azimuthal 

coverage of seismometers around Mount St. Helens. The existing one-dimensional velocity model described above, however, 

allows for a tradeoff between the depth and the origin time 

Figure 14. Instrument layout (blue triangles) with 

earthquake epicenters (red dots) used for P-wave 

inversion. Epicenter of primary earthquake used in 

inversion is shown by white dot.

within the upper 2.2 km of the velocity model. Thus, the depth 

is not as well constrained as the epicenter. With additional 

layers in the shallow part of the velocity model, this tradeoff is 

less important.

Secondary arrivals were observed in the record section 

of the transect, but they were not easily picked across the 

array. We therefore used only the P-wave arrivals. Errors in 

picking first arrivals on the event used for the inversion were 

assumed to be constant and were estimated to be within five 

time samples (0.05 s) of the actual arrival. Our algorithm 

allows for the use of one earthquake at a time. We chose to 

use only one earthquake in our inversion for two reasons: 

(1) there was only one other earthquake (August 22, 2005, 

03:32:45 PDT; Md = 1.4) that possessed impulsive arrivals 

across the entire transect, and (2) that earthquake showed an 

identical pattern in first-arrival picks to the earthquake we 

used for the inversion. Using multiple earthquakes would 

not change or add to the information used in the inversion 

because the earthquakes that occurred during the deployment of the transect occurred in such a tight cluster. First 

arrivals from the other earthquakes that occurred during the 

deployment were picked only if they could be identified with 

a precision of better than 0.05 s. The time spread in picks at 

a given offset was used to define the standard deviation to 

ensure that the inversion did not overfit the data.

We used a one-dimensional travel-time calculator 

parameterized with constant velocity layers. The traveltime calculator allows for a three-dimensional geometry of 

stations and earthquake location. The setup of the inverse 

problem prescribes node depths and solves for the layer 

velocity between nodes. The node number and depths were 

based largely on constraints imparted by the program used 

for earthquake locations, SPONG. The first constraint is 

the use of a total of 10 layers in the velocity model. When 

combined with the existing PNSN velocity model, this only 

allowed for 4 layers. The second constraint is that different station altitudes can be used, but they must be included 

inside the top layer of the velocity model. Stations within 10 

km of the crater, excluding the crater stations, span altitudes 

from 1,268 m to 1,700 m. We thus required a 500-m-thick 

top layer so that we could include the altitudes of these 

closest stations to the crater. Depths of the remaining three 

node-depths layers were found through a grid search with an 

increment of 100 m. The preferred model is the model that 

satisfies the condition above and has the smallest residual 

times (t

observed − tcalculated). Models with low-velocity inversions 

were poorly constrained using first-arrival data and thus were 

not considered.

The initial depth of the earthquake used in the inversion 

was 0.24 km (datum at 1,700 m above sea level). We iterated 

velocity-model solutions, each time updating the station 

corrections using a subset of PNSN catalog earthquakes 

with high-quality arrivals (see station correction discussion below) to get a new earthquake location and origin 

time. The new earthquake parameters were then used to 

update distance and time information of the transect, and the 
92 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

velocity model inversion was run again. Iterations stopped 

when the change in depth between iterations was less than 

the calculated uncertainty of the solution. The final depth 

of the earthquake was calculated to be 0.13 km. By alternately recalculating the location and the velocity model, we 

minimize errors in the location of the earthquake and in the 

velocity inversion results.

Location uncertainties in our nonlinear velocity inversion 

were calculated using the diagonal of the covariance, a method 

that is appropriate in linear inverse problems. We adapted this 

technique to our problem by linearizing the solution through a 

central difference and then calculating the error as if it were a 

linear problem. This approach is often used, though not strictly 

correct, and may underestimate the uncertainties in the model 

(Aster and others, 2005).

Station corrections are static time corrections used in 

earthquake hypocenter inversions using one-dimensional 

velocity models. Station corrections, which compensate for 

the one-dimensional modeling of a truly three-dimensional 

velocity structure, can be minimized by finding the closest 

one-dimensional approximation of the three-dimensional 

velocity structure under the volcano (Pujol, 2000). We do not 

minimize the station corrections here; however, we do use 

decreases in station corrections as a proxy for the quality of 

our velocity model.

Station corrections were calculated using a subset of 29 

PNSN catalog events that occurred between 2004 and 2006. 

These events were selected for their exceptionally sharp 

impulsive arrivals and were already in use for calibrating 

PNSN velocity models for Mount St. Helens (S. Malone, 

oral commun., 2006). The first arrivals of these events were 

picked on the PNSN stations only if the arrival was unambiguous to ±0.05 s. Locations were then calculated using 

the existing velocity model and station corrections. The 

new velocity model was then used with the existing, fixed 

location to calculate new residuals for each station for each 

event. The station residuals were then averaged across events 

to obtain the station correction. This approach is not without 

flaws. Earthquake locations using the old velocity model 

may have errors imparted by the velocity model itself, which 

would lead to a bias in the station corrections. Near-site 

effects may also introduce a bias in the station corrections. 

Despite these disadvantages, we find that comparing station corrections gives us the best measure of spatial quality 

between two velocity models.

Results

P-Wave Model

The final result of our P-wave velocity inversion is 

shown in figure 15. Disappointingly, our inversion did 

not constrain velocities deeper than 1.5-km depth. This is 

because of the lack of deep earthquakes during the deployment and the final length of the transect used. Our deepest 

velocity (5.12 km/s) agrees well with the shallowest velocity of the existing model that spans depths of 0 to 2.2 km. 

Inverting a subset of the 12 closest picks to the hypocenter 

resulted in shallow velocities as low as 2.4 km/s, instead of 

the 2.7 km/s that is calculated in the final model. Residuals 

of the model show trends imparted by the constant velocity 

layers used in the inversion. The 13 picks farthest from the 

hypocenter show apparent velocities that are abnormally high 

(7–8 km/s). In our inversion, we discarded points (seven in 

all) with the largest offsets until inverted velocities at depths 

of 1 to 1.5 km below the transect were reasonable with 

respect to the existing velocity model. Removing time picks 

from the most distant stations resulted in depth constraints 

that were not as deep as we had initially planned, but it was 

required in order to get reasonable results. The apparent 

high velocities in our time picks may be a result of threedimensional variations in wave speed. An alternative explanation may be a large amount of scattering that is shifting 

the energy of the direct arrival later in the signal and masking the energy remaining in the actual first arrival buried in 

the noise. Three-dimensional effects are also seen near the 

2.1-km offset (fig. 15A), where three picks are advanced with 

respect to surrounding picks.

0.2

0.6

1.0

1.4

Model

Data

0 5 10 15 20 25 30 35

-0.04

0

0.04

0.08

3 4 5 6

0.5

1

1.5

0

(2.66, 0.00)

(3.48, 0.50)

(4.30, 0.80)

(5.12, 1.10)

VELOCITY, KILOMETERS PER SECOND

DEPTH,

IN KILOMETERS

TRACE NUMBER

RESIDUALS,

IN SECONDS

DISTANCE FROM HYPOCENTER, IN KILOMETERS

1 1.5 2 2.5 3 3.5 4 4.5

TIME, SECONDS

 SINCE ORIGIN TIME

B

C

A

3D Effects

2–norm Residual: 0.25

Figure 15. Results of final velocity inversion and grid search for 

P-wave model. A, Comparison of data (circles) and model (stars) 

for the final model. B, Plot of residuals. C, Velocity model plotted 

with 90-percent confidence intervals shown as horizontal lines.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 93

Discussion

The results of our P-wave model show velocities of 

2.4–2.7 km/s in the uppermost layer, approximately 1 km/s 

slower than found in the velocity model used at Mount St. 

Helens by Musumeci and others (2002). The discrepancy is 

due to different ray coverage within the crater. In our inversion, we had 39 instruments draping the crater floor and 

flank and recorded earthquakes less than 1 km deep. In the 

velocity model by Musumeci and others (2002), only three 

stations existed inside the crater and only earthquakes deeper 

than 2.5 km were recorded. The P waves used in our study 

thus better sampled the shallow structure of the volcano. Our 

shallowest velocities are similar to those measured at other 

volcanoes, including Augustine Volcano, Alaska (2.24–2.6 

km/s; Pearson and Kienle, 1978), Redoubt Volcano, Alaska 

(2.9 km/s; Lahr and others, 1994), and Soufrière Hills, 

Montserrat (2.17 km/s; Rowe and others, 2004).

The new P-wave velocity model has relatively high velocities (4.3 km/s) at fairly shallow depths (0.8 km), considering 

the volcanic strata exposed within the crater. But such high 

velocities may be warranted, inasmuch as an intrusive suite at 

approximately 1-km depth has been interpreted from gravity 

data (Williams and others, 1987). Velocities of approximately 

5.1 km/s have also been used in layered velocity models at 

Augustine Volcano at 0.9-km depth (Reeder and Lahr, 1987) 

and at Redoubt Volcano at 0.6-km depth (Lahr and others, 

1994). Furthermore, the new P-wave model shows decreases 

in calculated station corrections at all stations on the volcano, 

suggesting that the new model is a closer approximation of the 

three-dimensional velocity structure under the edifice than the 

existing PNSN model (table 2).

By concatenating the existing S4 velocity model from 

the PNSN onto the bottom of the inverted velocity model, 

we are able to calculate improvements to station corrections 

resulting from the new model (table 2). The altitudes of our 

transect ranged between 1,100 m and 1,800 m above sea 

level. At best, our data only constrain the seismic velocities 

up to the maximum altitude of our transect. For consistency 

with the existing PNSN velocity model, we set our initial 

datum to 1,700 m. To allow for earthquake locations inside 

the new dome, we also considered a model identical to the 

final inverted velocity model, except with a datum at 2,200 m 

(datum-adjusted). The layer above the initial datum (1,700 

m) has a constant velocity equal to the uppermost layer (2.66 

km/s), and the layer depths are adjusted according to the new 

datum. Adjusting the datum to 2,200 m above sea level, the 

altitude of the highest seismic station, led to reduced station 

corrections. In particular, the station corrections on the dome 

are greatly reduced with respect to other velocity models we 

considered (table 2). The datum-adjusted model is preferred 

because it allows for earthquake locations within the new 

dome and results in lower station corrections for stations on 

the edifice and inside the crater. Admittedly, no dome material was present at 2,200-m altitude early in the study period, 

but in order to present the most consistent set of locations for 

the entire study period, a consistent velocity model should be 

used throughout the study period.

Appendix 2. Effect of Changing Station 

Configurations

We simulated the results of changing station configurations 

within the crater using PNSN catalog events from November 4 

through November 20, 2004. During this time period, stations 

BLIS, YEL, and SEP were operational (McChesney and others, 

this volume, chap. 7), allowing an opportunity to understand 

the effect of particular stations on the geometry of the absolute 

locations. Absolute locations from September 23 until October 

1 all lie along a north-dipping structure that descends beneath 

the 1980–86 dome. When good station coverage and impulsive 

arrivals were both present starting around November 4, 2004, a 

similar north-dipping structure emerged (fig. 16A). This northdipping structure was found to be fairly robust analytically, 

persisting even when the two closest stations were excluded 

(figs. 16B, C). Only when station YEL is also excluded is there 

a breakdown in the shape of the cluster.

Generally, as we removed the close stations, the hypocentroid of the cluster moved deeper and the range of depths 

grew larger. For example, the original locations had a hypocentroidal depth of ~290 m. Upon removing station BLIS from 

calculation, the hypocentroid deepened to ~350 m. When both 

SEP and BLIS were removed, the hypocentroid was more than 

400 m deep.

Appendix 3. Jackknife Tests

We show here the results of jackknife tests on the multiplet E7. The original relative locations show a tabular body with 

a surface projection spanning nearly 50 m. In depth the locations are restricted to a range of only about 20 m, outlining a 

cigar-shaped source region (figs. 10B, 17A). Multiplets D34 

and F1 also show similar tabular bodies, so we are compelled 

to show the results of our jackknife tests to prove that the locations are not an artifact of station geometry or poor waveform 

alignment. In figure 17, we show the results of removing, one 

at a time, each of the four stations with the most observations. 

In almost all cases there is no change in the geometry of the 

source region. An exception is the source geometry when 

station YEL is removed (fig. 17C). In this test, the orientation 

of the source region is more southwest than in the original 

locations, and the relative hypocenters, viewed to the east, are 

elongate north-south.

We also tested the effect of a P-wave velocity model on 

the relative locations (fig. 17F). As mentioned previously, 

the velocity model that we chose lacks layer boundaries near 

the cluster centroid, so a change to higher velocities simply 

spreads the hypocenters out over a larger area without gross 

change in geometry.
94 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

122.205° 122.19° 122.175°

46.185°

46.2°

46.215°

200 m

SEP

YEL

BLIS

46.185° 46.195° 46.205° 46.215°

0

0.5

1

1.5

2

2.5

3

0

0.5

1

1.5

2

2.5

3

LATITUDE

N

N

46.185°

46.2°

46.215°

SEP

YEL

BLIS

46.185° 46.195° 46.205° 46.215°

LATITUDE

ALTIUTDE ABOVE SEA LEVEL,

IN KILOMETERS

ALTIUTDE ABOVE SEA LEVEL,

IN KILOMETERS ALTIUTDE ABOVE SEA LEVEL, IN KILOMETERS ALTIUTDE ABOVE SEA LEVEL, IN KILOMETERS

N

46.185°

46.2°

46.215°

SEP

YEL

BLIS

46.185° 46.195° 46.205° 46.215°

LATITUDE

N

46.185°

46.2°

46.215°

SEP

YEL

BLIS

46.185° 46.195° 46.205° 46.215°

LATITUDE

N

Original Locations

No Station BLIS 

No Station BLIS and SEP 

No Station BLIS, SEP, and YEL 

A

B

C

D

200 m

200 m

200 m

122.205° 122.19° 122.175°

122.205° 122.19° 122.175°

122.205° 122.19° 122.175°

0

0.5

1

1.5

2

2.5

3

0

0.5

1

1.5

2

2.5

3

S

S N

S N

S N

Figure 16. (A–D) Results of simulations of different station configurations using Pacific Northwest 

Seismic Network (PNSN) catalog earthquakes from November 4, 2004, to November 20, 2004.
4. Absolute and Relative Locations of Earthquakes at Mount St. Helens : Implications for Magmatic Processes 95 DISTANCE, METERS

DEPTH, METERS

300

50

0

-50

DISTANCE, METERS

50

0

-50

DISTANCE, METERS

50

0

-50 DISTANCE, METERS

50

0

-50

180

200

220

240

260

280

300

320

250 200

DEPTH, METERS

300 250 200

300 250 200

300 250 200

-50 0 50

DISTANCE, METERS

-50 0 50

DISTANCE, METERS

-50 0 50

DISTANCE, METERS

DEPTH, METERS

180

200

220

240

260

280

300

320

-50 0 50

DISTANCE, METERS

DEPTH, METERS

DISTANCE, METERS

50

0

-50

DEPTH, METERS

300 250 200

180

200

220

240

260

280

300

320

-50 0 50

DISTANCE, METERS

DEPTH, METERS

DISTANCE, METERS

50

0

-50

DEPTH, METERS

300 250 200

180

200

220

240

260

280

300

320

-50 0 50

DISTANCE, METERS

DEPTH, METERS

180

200

220

240

260

280

300

320

DEPTH, METERS

180

200

220

240

260

280

300

320

DEPTH, METERS

No Station BLIS

No Station YEL

No Station HSR

No Station SHW

P-wave model

DEPTH, METERS

DEPTH, METERS

Original Locations

A

B

C

D

E

F

Figure 17. (A–E) Results of jackknife tests on multiplet E7. F, Results of the use of a P-wave model in the relative 

locations of multiplet E7.
96 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

This page intentionally left blank
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

1

 Center for Earthquake Research and Information, University of Memphis, 

3890 Central Avenue, Memphis, TN 38155

2

 U.S. Geological Survey, at Pacific Northwest Seismic Network, Department of Earth and Space Sciences, University of Washington, Box 351310, 

Seattle, WA 98195

3

U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Chapter 5

Broadband Characteristics of Earthquakes Recorded 

During a Dome-Building Eruption at Mount St. Helens, 

Washington, Between October 2004 and May 2005

By Stephen P. Horton1, Robert D. Norris2, and Seth C. Moran3

Abstract

From October 2004 to May 2005, the Center for Earthquake Research and Information of the University of Memphis 

operated two to six broadband seismometers within 5 to 20 

km of Mount St. Helens to help monitor recent seismic and 

volcanic activity. Approximately 57,000 earthquakes identified during the 7-month deployment had a normal magnitude 

distribution with a mean magnitude of 1.78 and a standard 

deviation of 0.24 magnitude units. Both the mode and range 

of earthquake magnitude and the rate of activity varied during 

the deployment. We examined the time domain and spectral 

characteristics of two classes of events seen during dome 

building. These include volcano-tectonic earthquakes and 

lower-frequency events. Lower-frequency events are further 

classified into hybrid earthquakes, low-frequency earthquakes, 

and long-duration volcanic tremor. Hybrid and low-frequency 

earthquakes showed a continuum of characteristics that varied 

systematically with time. A progressive loss of high-frequency 

seismic energy occurred in earthquakes as magma approached 

and eventually reached the surface. The spectral shape of large 

and small earthquakes occurring within days of each other did 

not vary with magnitude. Volcanic tremor events and lowerfrequency earthquakes displayed consistent spectral peaks, 

although higher frequencies were more favorably excited during tremor than earthquakes.

Introduction

Two classes of earthquakes are commonly observed at 

volcanoes around the world. One class, termed volcano-tectonic (VT) earthquakes by Latter (1981), consists of highfrequency events with sharp, well-defined body-wave phases. 

Volcano-tectonic events are produced by shear failure in solid 

rock; their broadband seismic waveforms are indistinguishable 

from normal tectonic earthquakes (Chouet, 1996). The second 

class of earthquakes observed at volcanoes consists of lowerfrequency events that are thought to originate from a boundary 

between a fluid, such as magma or gas, and the solid surrounding rock (see, for example, Neuberg and others, 2006).

In this study we follow the convention of Moran and others (this volume, chap. 2), dividing lower-frequency earthquakes into two groups, hybrid and low-frequency (LF) earthquakes. Low-frequency events have a weak high-frequency 

onset followed by a harmonic waveform with one to several 

dominant frequencies in the range of 0.5–5 Hz (Chouet, 1996). 

In LF events, the P wave is often emergent, and the S wave is 

often obscured. Hybrid events blend the characteristics of both 

the VT and LF events (Lahr and others, 1994). The onset of 

the hybrid earthquake shows more pronounced high-frequency 

energy than the LF event, whereas the coda of both types is 

dominated by a high-amplitude, low-frequency waveform 

(Chouet, 1996). First-motion polarities are often mixed for 

both VT and hybrid events, whereas LF events generally display the same polarity at all stations (Lahr and others, 1994).

A network of eight broadband seismometers was installed 

at Mount St. Helens in October 2004 (fig. 1) to record seismic 

activity associated with the dome-building eruption that began 

with seismic unrest on September 23. The Cascades Volcano 

Observatory (CVO) installed two broadband seismometers 

northwest of the Mount St. Helens crater in October 2004. The 

Center for Earthquake Research and Information (CERI) at 
98 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

the University of Memphis installed six temporary stations at 

complementary azimuths around the volcano to provide better 

coverage of the eruption at about the same time. The CERI 

stations were removed in May 2005, and a separate temporary 

broadband network, installed by the U.S. Geological Survey 

in May 2005, was operated until October 2005. This analysis 

covers only the time period of the first temporary deployment.

Moran and others (this volume, chap. 2) separated the 

seismic activity associated with the dome-building eruption 

at Mount St. Helens into two temporal phases. The first was 

the vent-clearing phase that began with a swarm of shallow 

VT earthquakes on September 23 and ended with an explosion on October 5. A gradual transition from the VT events 

to lower-frequency earthquakes occurred between September 

25 and October 1. This transition from VT to lower-frequency 

earthquakes was not recorded with broadband seismometers, 

because the first broadband seismometer was installed by 

CVO at Johnston Ridge Observatory (JRO; fig. 1) on October 

1, 2004, about 1600 PDT. After October 5, seismicity was 

dominated by lower-frequency earthquakes. Tens of thousands 

of lower-frequency earthquakes were recorded by the broadband network between then and May 2005, with just one highfrequency (VT) earthquake. A few episodes of volcanic tremor 

were also observed, mostly before the CERI component of the 

broadband network became operational (Moran and others, 

this volume, chap. 2).

Lower-frequency earthquakes at Mount St. Helens were 

observed in association with eruptions in the 1980s (see, for 

example, Endo and others, 1981; Malone, 1983; and Malone 

and others, 1983). Those events were recorded using shortperiod, vertical-component seismometers. Since the mid1990s, broadband seismometers have been used to record 

earthquakes and volcanic tremor on numerous volcanoes 

around the world, but the broadband recordings discussed here 

were the first at Mount St. Helens.

Because hybrid and LF earthquakes form the bulk of the 

signals recorded during the CERI broadband deployment, our 

analysis focuses on those events. We examine aspects of the 

seismicity that can be addressed by continuous, broadband, 

and unclipped recordings of the earthquakes. Following a brief 

description of the temporary broadband seismic network at 

Mount St. Helens, we discuss our search of the dataset for very 

Figure 1. Broadband seismic stations at Mount St. Helens used in this study (from table 1). Black circles 

indicate location of CVO broadband seismometers. Black triangles indicate initial locations of CERI broadband 

seismometers. Red triangles show four CERI instruments redeployed to lower altitudes in early December 2004 to 

allow servicing throughout winter. Red circle indicates centroid of seismic activity. Roads shown in black; dashed 

lines are trails. Contours in meters, contour interval 50 m.

WIND

LAVA

BLUE

WEST

CLIM

SMTH

JRO 

LSCT

SNOW

HRSE

PISH

STD 

46.12°

46.16°

46.24°

46.20°

46.28°

46.32° 0 5 10 KILOMETERS

122.3° 122.25° 122.2° 122.15° 122.1° 122.05° 122° 121.95°

Mt. St. Helens

WASHINGTON

AREA OF MAP

Spirit Lake

Coldwater L. SKAMANIA CO. COWLITZ CO.

1500

1000

2000

2500

1000

1000

1500

1000
5. Broadband characteristics of earthquakes recorded during a dome-building eruption at Mount St. Helens

99 

long period (VLP) pulses as described by Almendros and others (2002). We examine the variation in the rate and the magnitude distribution of the events identified by our trigger algorithm. We then look at the differences and similarities between 

hybrid and LF earthquakes and test a numerical measure of 

event character. Finally some issues are discussed related to 

wave propagation that make it difficult to determine focal 

mechanisms for these events through waveform modeling.

Broadband Network

Broadband seismic monitoring at Mount St. Helens 

began with CVO’s installation of a Guralp CMG-6TD at JRO 

on October 1, 2004, approximately 1 week after the onset 

of increased seismic activity at the volcano (fig. 1). CVO 

installed a second instrument at Studebaker Ridge (STD) on 

October 5. Telemetry for both sites was established in late 

October, providing real-time broadband monitoring of Mount 

St. Helens seismicity (McChesney and others, this volume, 

chap. 7). On October 3, CERI, in cooperation with CVO 

and the Pacific Northwest Seismic Network (PNSN), began 

deploying broadband instruments with onsite recording capability. Station locations were chosen to provide the optimum 

azimuthal coverage, encircling the vent at the closest possible 

distance to Mount St. Helens. Budget, safety, and logistical 

considerations limited us to road-accessible sites. We deployed 

six Guralp CMG-6TD broadband seismometers within about 

5 to 10 km of the seismic activity by October 5 (fig. 1). Four 

strong-motion instruments (Kinemetrics Altus K2 Recorder) 

were also deployed, although earthquakes in the sequence did 

not trigger the latter instruments.

A typical broadband station had a Guralp CMG-6TD 

sensor buried in a hole about 18 inches deep. Timing was 

provided by a GPS antenna. Power was provided by 12-V batteries recharged by a 30-W solar panel on a 1.2-m-tall stand. 

The broadband seismometers recorded three channels continuously at 50 samples per second and had a flat response to 

ground velocity between 0.033 and 25 Hz. Data were stored 

onsite using 2 GB of internal flash memory. With reasonable data compression, and a few sunny days for solar-power 

recharge, a service interval of 6 to 8 weeks was adequate to 

ensure no loss of data. In late October we started removing 

several temporary stations located in remote settings to avoid 

having them snowed under for the winter (table 1). By early 

November only two of the temporary stations, BLUE and 

LAVA, remained. Fortunately, these stations were southwest 

and southeast of the volcanic crater, so they, along with the 

two CVO stations, still provided reasonable azimuthal coverage. At the beginning of December, we moved stations to 

sites where the anticipated snow depth would not exceed 2.4 

m. Three of these sites were winter recreational Sno-Parks 

maintained by the U.S. Forest Service, with plowed road 

access during normal snow years. The fourth site, LSCT, was 

sufficiently low in altitude to be accessible year-round.

Very Long Period Seismicity, OctoberNovember 2004

One of the principal reasons for our rapid deployment of 

broadband seismometers was to record any very long period 

(VLP) seismicity associated with the eruption that would 

otherwise have been missed by the PNSN short-period stations. We looked for evidence of VLP signals as described 

by Almendros and others (2002) in the period range 5–30 s. 

A moving root-mean-squared (RMS) average applied to 

bandpass-filtered seismograms was helpful in identifying 

trends and potential pulses. We checked peaks and trends 

observed in the moving average by visual inspection of the 

time series. For example, a moving RMS average was computed from seismograms recorded from October 5 through 

October 8 (fig. 2A). The two large peaks observed at station 

WIND in figure 2A at about 1600 (PDT) on October 6 and 

0200 on October 8 can be correlated with spikes at other stations in the network. The first peak was produced by surface 

waves following a shallow Mw 6.2 earthquake that occurred in 

Indonesia at 1531 (PDT) on October 6. The second large peak 

was produced by body and surface waves following a shallow 

Mw 6.8 earthquake that occurred in the Solomon Islands at 

0127 on October 8.

Two broad increases in RMS levels occurred between 

October 7 and 9. These trends were too broad to be teleseismic arrivals or VLP pulses but could have been caused by 

very long duration resonance of the volcano’s magmatichydrologic system. We computed the RMS average at station 

LON, approximately 68 km north of Mount St. Helens and 

found that the broad trend increases observed at WIND, only 

7.15 km from the vent, were matched by concurrent increases 

at LON, which suggests the source of the increase may be 

regional. Seismic waves that emanated from a local source 

below the Mount St. Helens crater should have attenuated with 

distance. To investigate amplitude as a function of distance 

from the crater, we plotted seismograms from October 7 to 9 

arranged by the station distance from the crater (fig. 2B). The 

peak amplitudes and general character of the waveforms were 

similar at all stations, including LON. Because attenuation 

was not evident, we concluded that the broad energy increases 

did not result from a local source below the Mount St. Helens 

crater, but rather from a regional source. The increased longperiod energy is best explained as seismic noise (microseisms) 

generated by a low-pressure weather system that approached 

the Pacific Northwest coast on October 7 and remained in 

the coastal area during the following 3 days. Microseisms, a 

nearly harmonic modulated waveform with periods around 6 s, 

have long been associated with storms over water, including 

normal low-pressure systems. Longuet-Higgins (1950) showed 

that the interaction of water waves with similar periods and 

different directions generates pressure waves that efficiently 

couple with the solid earth to produce seismic surface waves. 

These surface waves can propagate with gradual loss of energy 

over broad areas of the continent.
100 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Table 1. Broadband station locations and dates of operation.

[Stations JRO and STD operated by Cascades Volcanic Observatory; others installed by Center for Earthquake Research and Information as part of this report. Datum is WGS84. Altitude is orthometric, interpolated from topographic maps after plotting locations. 

Distance is measured from centroid of seismic activity, the surface trace of which approximates the position of the vent.]

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0.9

Teleseism

WIND Microseism

LON

example

5 6 7 8

DATE IN OCTOBER 2004

A

B

BLUE

JRO

SMTH

WEST

WIND

LON

400 600 800 1,000 1,200 1,400

TIME, IN SECONDS

0

1

2

 1

0

1

2

 1

0

1

2

 1

0

1

2

 1

0

1

2

 1

0

1

2

 1VERTICAL COMPONENT AS RMS AVERAGE,

IN CENTIMETERS PER SECOND x 10-4

VERTICAL COMPONENT,

IN CENTIMETERS PER SECOND x 10-4

Figure 2. Ground motion recorded 

at Mount St. Helens, October 5–8, 

2004. A, Moving RMS average 

(over consecutive 600-s window) of 

filtered (5–30 s) vertical component 

data at station WIND (black) 

and Pacific Northwest Seismic 

Network broadband station LON 

(red). Data from LON plotted 

in counts but scaled for visual 

comparison with velocity data 

from WIND. Date on x axis marks 

start of each day, PDT. Concurrent 

large sharp peaks at both stations 

are teleseismic arrivals. Broad 

increase that peaks about midday 

October 7 was most likely due 

to weather system impinging on 

region. Arrow labeled “example” 

shows approximate time displayed 

in bottom figure. B, Vertical 

component seismograms, all 

scaled to same absolute amplitude, 

for ~16-minute interval of intense 

low-frequency energy as recorded 

at six stations (listed at right edge).

Station Latitude 

(N)

Longitude 

(W)

Altitude 

(m)

Distance from vent 

(km) Dates of operation

JRO 46.2751 122.2178 1,290 8.98 10/01/04–08/10/06

STD 46.2376 122.2240 1,250 5.28 10/25/04–present 1

WIND 46.2504 122.1372 1,230 7.15 10/03/04–10/31/04

LAVA 46.1509 122.0908 850 9.13 10/03/04–12/02/04

BLUE 46.1682 122.2608 980 6.41 10/03/04–12/02/04

WEST 46.1946 122.3143 1,120 9.66 10/04/04–11/01/04

CLIM 46.1470 122.1825 1,140 5.58 10/04/04–11/01/04

SMTH 46.2128 122.0566 1,090 10.34 10/05/04–10/27/04

LSCT 46.1818 122.0540 500 10.52 12/03/04–05/07/05

SNOW 46.1306 122.1708 810 7.51 12/04/04–05/07/05

HRSE 46.1434 122.3244 620 12.02 12/05/04–05/06/05

PISH 46.3256 121.9754 860 21.77 12/07/04–05/08/05

1

 As of March 17, 2008.
 101

Visual inspections of all moving RMS average peaks 

occurring in October and November 2004 revealed no VLP 

signals that originated below the Mount St. Helens crater. The 

absence of observed VLP signals could have been due in part 

to the distance of our stations from the source. Rapid decay 

of signal amplitude with distance is considered characteristic 

of VLP signals (Almendros and Chouet, 2003). VLP signals 

have been observed at other volcanoes at distances of less than 

5 km (for example, Almendros and others, 2002; Aster and 

others, 2003). The closest broadband stations recording during 

the first deployment were STD and CLIM, about 5.28 km and 

5.58 km respectively from the centroid of seismicity (fig. 1). 

Thus we cannot definitively rule out the occurrence of VLP 

signals at Mount St. Helens during this time period.

October 2004–May 2005 Seismicity 

Recorded by the Broadband Network

Event Magnitudes and Rates

Thousands of earthquakes were recorded during the 

7-month broadband network deployment at Mount St. Helens. Because our data were not recorded in trigger mode, we 

developed a trigger algorithm based on the ratio of short-term 

(0.2 s) and long-term (320 s) averages to isolate individual 

events from the continuous records. We first applied a 0.5–3-

Hz bandpass filter to the data and then used a short-term to 

long-term trigger ratio of 6 to identify 57,635 events. We 

assigned a magnitude for each event on the basis of the peak 

velocity (pv) measured on the vertical component at each CERI 

station. The body-wave arrivals of both hybrid and LF events 

were of smaller amplitude than the low-frequency waveforms 

that followed, so we assumed pv to be a surface-wave amplitude. We therefore adjusted the measured pv by the square root 

of the distance from the centroid of seismicity to the station 

(table 1) to account for geometrical spreading. The centroid 

location, 46.1964oN and 122.18882oW, was calculated from the 

PNSN catalog for events with Md ≥2.5 during this time period. 

The magnitude was given by MS=log10(pv)+5.78, where the 

distance-adjusted pv was averaged over all CERI stations and 

the constant (5.78) was calculated so that the largest event had 

MS

 3.4, equal to the PNSN duration-magnitude estimate for 

that event. Our magnitude estimates were reasonably consistent 

with PNSN magnitudes for events above Md 2.6.

Figure 3A shows the distribution of event magnitudes 

selected by our trigger algorithm for the entire period of the 

network deployment. These have a normal distribution, with 

mean magnitude (MS

) of 1.78 and a standard deviation of 

0.24 magnitude units. When comparing event triggers to the 

continuous seismograms, we observed that small earthquakes 

occurring in the coda of larger events were not likely to be 

selected by our trigger algorithm. Therefore, the distribution 

of smaller earthquakes in the sequence is probably not well 

characterized. The distribution of larger events, however, is 

probably well characterized. The distribution of earthquake 

size and the rate of activity varied with time, with larger 

earthquakes occurring irregularly throughout the sequence. 

To help discern patterns, we divided the entire dataset into 

24-hour segments, calculated the magnitude distribution of 

each time interval as in figure 3A, and contoured the resulting 

magnitude distributions over time (fig. 3B). Most of the events 

with MS ≥3 occurred before October 5 and were associated 

with the closing days of the vent-clearing phase (Moran and 

others, this volume, chap. 2). The rate of seismic activity was 

quite high at that time, and the distribution of earthquake sizes 

was bimodal, with peaks around MS 2.5 and MS 2 and a range 

slightly greater than 2 magnitude units.

After October 5 the rate of seismic activity slowed drastically and the modal event size decreased to MS

 1.7 (fig. 3B). 

The minimum event size also increased to about MS 1.5 from 

a minimum MS of ~1.0 during the more active vent-clearing 

10/01 11/01 12/01

2004 2005

01/01 02/01 03/01 04/01 05/01

0.5

1

1.5

2

2.5

3

3.5

DATE, IN MONTH/DAY FORMAT

EARTHQUAKE MAGNITUDE

B

0.5 1 1.5 2 2.5 3 3.5 4

1,000

2,000

3,000

4,000

5,000

6,000

7,000

A

MAGNITUDE

NUMBER OF EVENTS

SCALE SHOWING EVENTS PER DAY

1 50 100 150 200 250 300 350 400 450

Figure 3. Event-magnitude occurrence for the more than 57,000 

earthquakes identified during this deployment. Magnitudes 

are binned in 0.072-magnitude intervals. A, Histogram showing 

magnitude distribution. B, Time-series plot showing magnitude 

distribution. Contour interval is 50 events per day.

5. Broadband characteristics of earthquakes recorded during a dome-building eruption at Mount St. Helens
102 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

period. Because the trigger algorithm should have been more 

sensitive to small events when the rate of larger events was 

low, this increase of minimum event size during a period of 

reduced seismic activity suggests that the observed distribution for small events was not entirely due to lack of detection. 

Instead, small events were simply not plentiful during some 

time intervals.

Both the range and mode of earthquake size and the rate of 

seismic activity increased through the latter part of November 

2004. These increases coincided with the approach and collision of spine 3 with the south crater wall (Schilling and others, 

this volume, chap. 8; Vallance and others, this volume, chap. 9). 

The seismic activity continued at a high rate through the middle 

of December, when 10 events with MS≥3.0 occurred between 

December 17 and 19. Subsequently, both the rate of seismic 

activity and the modal size of events began to decrease. This 

trend continued through the middle of January, as spine 3 was 

disintegrating and being shoved aside by spine 4, and a large 

explosion occurred on January 16, 2005 (Moran and others, 

this volume, chap. 6). The seismicity rate increased through 

the latter part of January 2005. The small number of events 

between January 29 and February 6 was artificial, reflecting a 

loss of network data during that time interval. The seismicity 

rate remained high in March and April, when 14 MS ≥3.0 events 

occurred. It appears that the sporadic occurrence of events of MS 

≥3.0 coincided with intervals of high seismicity rates.

Characteristics of Lower-Frequency 

Earthquakes at Mount St. Helens

Since the mid-1990s, broadband seismometers have been 

used to record earthquakes and tremor on numerous volcanoes 

around the world. Because our data were the first broadband 

recordings at Mount St. Helens, we looked with special care at 

the broadband character of the events recorded on our network. We recorded one high-frequency tectonic earthquake, 

two episodes of volcanic tremor, and thousands of lower-frequency earthquakes during the deployment. In this section we 

describe the characteristics of the lower-frequency events and 

compare our observations to the reported features of earthquakes observed at other volcanoes.

The first step was to compare a lower-frequency event to 

the high-frequency tectonic earthquake, because these form 

the two main types of earthquakes recorded at volcanoes (see, 

for example, Latter, 1981). We did not record volcano-tectonic 

events associated with the vent-clearing phase, because the 

broadband seismometers were only installed at the close of 

that phase. However, a small local tectonic event was recorded 

on November 22, 2004, at 0853 PST. Figure 4 shows seismograms and their associated spectra for that tectonic earthquake 

and an example LF earthquake, MS

 3.3, recorded on November 

27, 2004, at 0601 PST. The tectonic event was not located by 

PNSN. However, a local hypocenter is required by the short 

S-P arrival-time differences at stations LAVA (1.5 s) and 

BLUE (2 s). We consider the event to be tectonic because it 

had sharp and well-defined P and S phases followed by a short 

high-frequency coda. Peak energy for this event was between 

5 and 12 Hz, with very low spectral levels at frequencies less 

than 5 Hz. By contrast, the LF event had a low-amplitude 

onset, with frequencies as high as about 10 Hz, followed by 

a strong, long-duration, and lower-frequency waveform. The 

spectrum had significant energy between 0.3 and 5 Hz, with 

multiple spectral peaks suggestive of harmonic motion as 

observed for lower-frequency earthquakes at other volcanoes 

(for example, Chouet, 1996; Neuberg and others, 2006).

Next we compare the hybrid and LF earthquakes that 

form the two types of lower-frequency earthquakes recorded 

during the first year of renewed eruption at Mount St. Helens 

(Moran and others, this volume, chap. 2). For example, figure 

5A shows the broadband seismograms for a MS

 3.13 hybrid 

earthquake recorded on October 3, 2004, at 1812 (PDT) and 

a MS 3.2 LF earthquake recorded on November 28, 2004, at 

1807 (PST). Figure 5B shows the same seismograms after 

7-Hz highpass filters were applied, and figure 5C shows the 

Fourier displacement spectra for both events. The broadband waveforms of both events are dominated by the lowerTectonic

LF

TIME, IN SECONDS

Tectonic

LF

FOURIER VELOCITY SPECTRUM

VELOCITY, IN

CENTIMETERS PER SECOND

FREQUENCY, IN HERTZ

A

B

0

0 5 10 15 20 25 30 35 40

0

0.01

0.01

10-6

10-5

10-4

10-3

10-2

10-1 2 3 4 5 6 7 89 2 3 4 5 6 7 89 2

100 101

−0.01

Figure 4. Seismograms and spectral characteristics for tectonic 

earthquake recorded November 22, 2004, and a low-frequency 

earthquake recorded November 27, 2004. A, Comparison of 

seismograms from station LAVA, plotted at same absolute scale. 

B, Fourier velocity spectra of associated seismograms. Tectonic 

earthquake is black trace; low-frequency event is red trace.
 103

frequency (from 0.3 to about 5 Hz), nondispersive waves that 

give rise to the name of this class of earthquake. At frequencies between 0.3 and 2.0 Hz, the spectral amplitudes of the LF 

event slightly exceed those of the hybrid event (consistent with 

the slightly larger magnitude estimate), whereas at frequencies above 3.0 Hz, the spectral amplitudes of the hybrid event 

exceed the LF spectral amplitudes. This difference in highfrequency energy is also apparent in the time series, where 

the hybrid earthquake has a much more pronounced highfrequency onset (fig. 5A).

This difference in high-frequency energy is the primary difference between hybrid and LF earthquakes. It may 

result from a difference in the earthquake source process 

or a difference in the rock along the path from source to 

receiver that affects wave propagation (an increase in attenuation related to a reduction in source depth, for example). 

To investigate these possibilities, we filtered the lowerfrequency energy from the signal. The filtered records (fig. 

5B) show that both events actually have a high-frequency 

onset, although the amplitude is about five times larger for 

the hybrid event. Further, the high-frequency component 

of both events is similar in form and of comparable duration, extending well into the low-frequency waveform. This 

similarity in form and duration indicates that attenuation near 

the earthquake sources and along the paths to the receiver did 

not differ significantly. Rather, more high-frequency energy 

was generated at the source for the hybrid than for the LF 

earthquake. This is consistent with a larger component of 

shear faulting during the hybrid earthquake.

We explored the similarity in spectra between small and 

larger events in figure 6, where the velocity spectrum of the 

MS 3.1 hybrid event (October 3, 2004) was compared to a 

randomly chosen MS 1.7 event that occurred on October 11, 

2004. The spectrum of the MS 1.7 event was multiplied by a 

factor of 20 in order to compare the shape of the spectra of 

both events. The spectrum of the smaller event disappears 

into the background noise and is undetectable above 10 Hz 

and below 0.5 Hz. Between 0.5 and 10 Hz, the spectra of 

the large and small events were quite similar in shape and 

amplitude. The similarity in spectral shape suggests that both 

the source and path were similar for the two events. For tectonic earthquakes, the peak in the velocity spectrum would 

be expected to shift to lower frequencies as rupture zone size 

increased for larger earthquakes (Brune, 1970). The consistency in the frequency range of the peak of spectral energy 

for different-size earthquakes favors models such as that of 

Chouet (1996), in which a pressure transient resonates in a 

crack of constant dimension (a few centimeters in width), or 

that of Neuberg and others (2006), in which brittle failure of 

magma provides seismic energy that resonates in a conduit 

section of constant dimension (~30 m wide and several hundred meters long). In both cases the amplitude of the resonance would be determined by the strength of the pressure 

transient or size of the brittle failure, whereas the shape of 

the spectrum would be constrained by the size and properties 

of the resonator.

TIME, IN SECONDS

0

 0.010

0

0.010

 0.010

0

0.010

5 10 15 20 25 30 35 40

VELOCITY,

IN CENTIMETERS PER SECOND

Hybrid

LF

A

B

TIME, IN SECONDS

VERTICAL COMPONENT VELOCITY,

IN CENTIMETERS PER SECOND

 FACTOR x 10-4 FACTOR x 10-3

Hybrid

LF

0

0

6

4

2

2

4

6

0

2

2

5 10 15 20 25 30 35 40

Hybrid

LF

FREQUENCY, IN HERTZ

DISPLACEMENT SPECTRUM

10-6

10-5

10-5

10-4

10-3

10-2

10-1 2 3 4 5 6 7 89 2 3 4 5 6 7 89 2

100 101

C

Figure 5. Comparison of seismic data from station BLUE for 

hybrid and low-frequency (LF) earthquakes. A, Vertical-component 

seismograms, plotted at same absolute scale. B, Corresponding 

7-Hz highpass-filtered seismograms; note different scales. Decay 

of the waveforms is remarkably similar. C, Fourier displacement 

spectra of the original hybrid (black) and LF (red) seismograms. 

Spectrum of a 60-s interval from a “quiet” period on November 

1, 2004, shows representative background noise level (blue). 

(This interval chosen because time intervals just before the two 

earthquakes contain the coda of other earthquakes.)

5. Broadband characteristics of earthquakes recorded during a dome-building eruption at Mount St. Helens
104 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Resonance in a crack or conduit of constant dimensions 

produces amplitude spectra with strong peaks resulting from 

the summation of waves propagating between the boundaries of the structure. The specific frequencies of resonance 

are determined by the dimensions of the structure and 

the properties of the material within the conduit or crack. 

Identifying the frequencies of resonance would therefore 

place useful constraints on the volcanic process. Any given 

station can have a variety of peaks derived from its own 

particular source-receiver path and local site conditions, but 

the resonance peaks associated with a symmetric conduit or 

crack should be common to spectra observed at all stations 

independent of azimuth and distance. By comparing spectra 

at all stations, it should simply be a matter of identifying the 

common peaks.

A feature of the seismicity associated with this eruption 

of Mount St. Helens was the prominence of earthquake multiplets having highly similar waveforms (Thelen and others, this 

volume, chap. 4). Figure 7 shows the Fourier velocity spectra 

at six stations for a multiplet of 29 lower-frequency events 

(2.1 ≤ MS ≤ 2.2) that occurred on October 16, 2004. The 

velocity spectra were normalized and overlaid for each station 

so that peaks that are consistent from event to event tend to 

stand out. Each station exhibits several consistent peaks in the 

range 0.4–2 Hz, but no peak appears consistently at all stations 

to indicate a specific frequency of source resonance. This lack 

of identifiable frequencies of source resonance is inconsistent 

with the resonance of a structure with simple geometry such 

as the rectangular crack of the Chouet (1996) model and the 

rectangular conduit of the Neuberg and others (2000) model. 

It suggests instead that wave propagation dominated the 

response at a given station and that source harmonics played 

Ms 3.2

Ms 1.7

VELOCITY SPECTRA

FREQUENCY, IN HERTZ

103

104

105

106

10-1 2 3 4 5 6 7 89 2 3 4 5 6 7 89 2

100 101

Figure 6. Spectral shapes of large and small events. Verticalcomponent Fourier velocity spectra recorded at station BLUE are 

shown for a MS

 3.2 hybrid event (black) and a randomly chosen 

MS

 1.7 event (red) occurring close together in time. Spectra of 

the small event (scaled up by a factor of 20) and large event are 

similar in shape between 0.5 and 10 Hz.

a minor role. Alternatively, one can appeal to the excitation 

of a more complex, perhaps asymmetric, structure, where the 

harmonic response observed at a given station depends upon 

its location relative to the orientation of the source structure.

Fehler (1983) analyzed short-period records of tremor 

at Mount St. Helens that accompanied eruptions on August 

7 and October 16–18, 1980. He found that the spectra of 

lower-frequency earthquakes were very similar to volcanic 

tremor, and he suggested that tremor was composed of many 

lower-frequency events inseparable in time. Neuberg and others (2000) observed that many episodes of volcanic tremor at 

Soufrière Hills volcano in Montserrat, West Indies, followed 

swarms of lower-frequency earthquakes in which the rate of 

earthquake occurrence continually increased until they merged 

into tremor. Figure 8A shows a seismogram with lowerfrequency earthquakes followed by 58 minutes of volcanic 

tremor at Mount St. Helens on October 2, 2004. In this case, 

the lower-frequency earthquakes appeared to be distinct events 

occurring before and again after the volcanic tremor. There is 

no evidence to suggest that individual earthquakes merged to 

generate the volcanic tremor. Figure 8B shows the normalized 

velocity spectrum of one lower-frequency event and a 60-s 

window from the volcanic tremor. Neither spectrum displayed 

harmonic peaks at whole-number multiples of fundamental 

SMTH 10.3

WEST 9.7

LAVA 9.1

WIND 7.2

BLUE 6.4

CLIM 5.6

NORMALIZED VELOCITY SPECTRA

0.1 1.0 10.0

FREQUENCY, IN HERTZ

Figure 7. Comparison of spectra from six stations to identify source 

peaks. Fourier velocity spectra are plotted for a multiplet of 29 

lower-frequency events (2.1 ≤ MS ≤ 2.2) that occurred October 16, 

2004. Velocity spectra are normalized and overlaid for each station 

so that peaks consistent from event to event become prominent. The 

distance to vent (see table 1) is to right of each station name.
 105

A

B

TIME, IN SECONDS

NORMALIZED VELOCITY SPECTRUM VELOCITY, IN CENTIMETERS 

PER SECOND (x 10-3)

FREQUENCY, IN HERTZ

0

0.2

0.4

0.6

0.8

1

0 1

1,0000 2,000 3,000 4,000 5,000 6,000 7,000

2 3 4 5

-6

-4

-2

0

2

4

frequencies. Both spectra had distinct peaks in common (for 

example, 0.8 and 1.1 Hz), consistent with Fehler’s (1983) 

observations. However, the lower-frequency earthquake spectrum had a prominent peak at 0.6 Hz that was not observed 

in the spectrum of the volcanic tremor. This lower-frequency 

peak would be consistent with resonance of a structure having 

a larger dimension or with favorable excitation of the longer 

dimension of the same structure.

The progression from primarily volcano-tectonic events 

at the beginning of the 2004 eruption to combined hybrid and 

low-frequency (LF) events, then to dominantly LF events, 

reported by Moran and others (this volume, chap. 2), suggests a progressive loss of high-frequency seismic energy as 

magma moved toward and eventually onto the surface. Malone 

and collaborators have pointed out that gradual changes in 

event character were observed in many precursory swarms 

LOG10 FREQUENCY, IN HERTZ

LOG10 POWER SPECTRAL DENSITY

–1 –0.5 0 0.5 1

4

5

6

7

8

9

10

11

12

13

Figure 8. Seismic data showing 2-hour sequence of tremor 

recorded October 2, 2004, at JRO broadband station. A, 

Seismogram showing north component. Record of tremor is 

punctuated by impulsive earthquakes of lower-frequency type. 

B, Normalized velocity spectra for a lower-frequency earthquake 

(red) and the volcanic tremor (black) of October 2.

preceding dome-building eruptions at Mount St. Helens in 

the 1980s (Malone, 1983; Malone and others, 1983). During 

the 1980–86 dome-building eruptions, this transition became 

a diagnostic tool to predict eruption onset time (Malone and 

others, 1983; Swanson and others, 1983; Swanson and others, 

1985). The high-frequency spectral difference between hybrid 

and LF events observed in figure 5C suggests that the slope 

of the high-frequency falloff could provide a measure to track 

the loss of high-frequency energy through time. To investigate 

this, the power spectra of an example hybrid and example 

LF event are compared in figure 9. Lines were fit by least 

squares to the high-frequency spectral amplitudes in the range 

1–10 Hz. The slope of each line quantifies the rate of spectral 

amplitude decay with frequency or the high-frequency falloff 

of the velocity spectrum. The slope of the fitted line was −2.6 

for the hybrid event and −4.9 for the LF event, indicating that 

this slope is a useful measure to quantify the high-frequency 

earthquake character.

We computed the high-frequency falloff for all events 

between October 4 and December 1, 2004 (CERI network 

station locations changed after December 1). Figure 10A

shows earthquake magnitude versus time. Those events with 

magnitude in the range 1.5–2.0 are plotted in black, and other 

magnitudes are plotted in red. Because the high-frequency 

falloff could have a magnitude bias, the high-frequency 

falloff is plotted versus time (fig. 10B) with this same color 

convention. The slope of the high-frequency falloff ranged 

from approximately −1.75 to −5, defining a continuum 

between the ideal hybrid and LF events such as observed at 

Montserrat by Neuberg and others (2000). In early October, 

the slope values ranged from about −1.75 to −3.5, regardless of magnitude. The range was somewhat narrower for 

events with magnitude 1.5–2.0. Over time the slope values 

Figure 9. Power spectra of the vertical velocity components 

of a hybrid event (red) and a low-frequency event (black). Lines 

are least-squares fit to spectral amplitudes between 1 and 10 Hz; 

slopes are −2.6 for hybrid event and −4.9 for the LF event.

5. Broadband characteristics of earthquakes recorded during a dome-building eruption at Mount St. Helens
106 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

decreased to between about −3 and −4 at the end of October. 

The slope values continued at this level or decreased slightly 

until the end of November. If we arbitrarily declared a slope 

of −3.0 as the dividing line between hybrid and LF, then 

this figure would suggest about a 4:1 ratio of hybrid to LF 

events occurred between October 4 and October 5, decreasing smoothly to a 1:1 ratio by October 10, and decreasing to 

a 1:4 ratio by October 17. Other than a small increase around 

October 20, the slope continued to decrease so that by October 24, around 95 percent of the events were low frequency.

The gradual decrease in the slope of the high-frequency 

falloff over time reflects a progressive loss of high-frequency 

seismic energy. The loss of high-frequency seismic energy 

coincides with magma approaching and reaching the surface on October 11 (Scott and others, this volume, chap. 1). 

This lessening of high-frequency energy could be related 

to a decrease in earthquake source depth, with near-surface 

heterogeneity and anelasticity serving to filter out the higherfrequency parts of the seismograms. Alternatively, it may be 

related to a decrease in the static stress drop accompanying 

frictional slip as the magma plug, which initially needed to 

force its way through overlying rock, met with less resistance 

after the path to the surface was cleared.

Waveform Modeling

The record section in figure 11A shows bandpassfiltered (0.1–1 Hz) seismograms for a Ms 2.9 LF earthquake 

that occurred on March 1, 2005, at 0928 PST. For this event, 

the acceleration record at station SEP in the crater was 

available. The SEP record was substantially clipped, but the 

filtered record should still illustrate the general character of 

ground acceleration in this passband. This earthquake was 

assigned a depth of 0.4 km by the PNSN. The seismograms 

all exhibited a low-amplitude onset followed by a longduration, low-frequency waveform even at the closest station 

(SEP) recording acceleration. Signal duration increased with 

distance, from 15–20 s at less than 1 km epicentral distance 

from the source to ~50 s at 20 km distance. The first onset of 

larger-amplitude arrivals propagated at roughly 2 km/s. To 

investigate particle motion we chose a 6-s window having 

Figure 10. Time series showing seismic data from station BLUE. A, Magnitude versus time. Black circles show events in the 

range 1.5 ≤ Ms

≤ 2.0; red circles are all other events. B, Plot of high-frequency slope versus time. Symbol color is determined by 

magnitude as in A (black circles, 1.5 ≤ Ms ≤2.0).

1

2

3

4

10/03 10/10 10/17 10/24 10/31 11/07 11/14 11/21 11/28 12/05

–5

–4

–3

–2

–1

0

EARTHQUAKE MAGNITUDE HIGH-FREQUENCY SLOPE

DATE IN 2004, AS MONTH/DAY

A

B
 107

a group velocity of 2 km/s for each station. At station PISH 

the particle motion in the vertical-radial plane was retrograde 

elliptical motion, characteristic of Rayleigh waves. However, 

the transverse component also had significant energy in this 

window, indicative either of Love waves or of multipathing. 

For the closer stations, the particle motion in the verticalradial plane was more complex, suggesting interaction of 

body and surface waves in this window.

We used a reflection matrix method (Randall and others, 

1995) to compute synthetic seismograms for a one-dimensional (1D) velocity model. The velocity model (table 2) was 

modified from the P-wave velocity model of Lees and Crosson (1989). The S-wave velocity in each layer was calculated 

from the P-wave velocity using a ratio of 1.7, based on the 

work of Sudo and Kong (2001), who found an average VP / VS

ratio of 1.704 fit observed P- and S-wave travel times at all 

stations for a large number of well-recorded earthquakes at 

Aso volcano, Kyushu, Japan. A low-velocity surface layer 

with Swave velocity of 2 km/s was required to have the 

synthetic surface waves propagate at a velocity near 2 km/s. 

Moderate Q values were assumed (see table 2). We assumed 

an implosive source with −1 on the diagonal elements of the 

moment tensor and 0 on the off-diagonal elements. The syn2 km/s

A

B

2 km/s

SEP

STD

SNOW

JRO

LSCT

PISH

HRSE

SEP

STD

SNOW

JRO

LSCT

PISH

EPICENTRAL DISTANCE FROM

SOURCE, IN KILOMETERS

EPICENTRAL DISTANCE FROM

SOURCE, IN KILOMETERS

HRSE

20

0

TIME, IN SECONDS

0 10 20 30 40

0 10 20 30 40

50 60

TIME, IN SECONDS

5

10

15

20

10

0

thetics were convolved with a triangular source pulse of 1-s 

duration and then lowpass filtered at 1 Hz.

The synthetic waveforms in figure 11B did a reasonable 

job of predicting the character of the initial arrivals and the 

beginning of the surface wave train. The synthetic P wave had 

a down first motion and was emergent at all azimuths. The 

synthetics do not have the long duration of the observed surface 

waves but rather a slightly dispersed surface wave traveling at 

approximately 2 km/s. We also tested other reasonable permutations of the layer velocities within a vertically stratified velocity 

model without producing the low-frequency coda. From this 

exercise we concluded that the low-frequency coda observed 

in figure 11A was not produced by a slow surface layer in a 1D 

velocity model. Scattering of low-frequency surface waves by 

the extreme local topography of the volcano would be a likely 

source of some of the coda. The resonance of a source structure 

(for example, Chouet, 1996, or Neuberg and others, 2006) could 

also contribute. Examination of surface-wave propagation in a 

3D velocity model was beyond the scope of the current study. 

Unfortunately, we concluded that our simple 1D velocity model 

was inadequate for modeling waveforms for the purpose of 

determining the source time function and focal mechanism for 

these events, because the coda would map into the source.

Figure 11. Seismic data for 

earthquake on March 1, 2005. 

Stations are identified at right. A, 

Time series for record section from 

vertical component. Vertical traces 

have been bandpass filtered between 

0.1 and 1 Hz. First onset of the surface 

wave propagates at roughly 2 km/s. 

B, Synthetic record section assuming 

a one-dimensional velocity model 

(table 2) with a moderately slow 

surface layer (see text). Bandpassfiltered traces between 0.1 and 1 Hz 

are shown for an implosive, triangular 

source pulse. The surface wave 

propagates at roughly 2 km/s.

5. Broadband characteristics of earthquakes recorded during a dome-building eruption at Mount St. Helens
108 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Summary

A network of broadband seismometers was rapidly 

deployed at Mount St. Helens when it became restless in 

September 2004 and maintained through the winter of 2004–5 

without a significant loss of data. A principal reason for deploying broadband seismometers was to record any very long period 

(VLP) seismicity (5–30 s) associated with the eruption that 

would otherwise have been missed by the PNSN short-period 

stations. We found no evidence of VLP pulses at our stations in 

the first 2 months of recording. Our closest stations were more 

than 5 km distant from the centroid of seismic activity, and these 

pulses may attenuate too quickly to be observed at that distance.

More than 57,000 earthquakes were recorded during the 

7-month deployment. They had a normal magnitude distribution, with a mean magnitude of 1.78 and a standard deviation 

of 0.24 magnitude units. Both the mode and range of earthquake size and the rate of activity varied during the 7-month 

period. Some cases of this variation can be associated with 

spine growth and breakup. In general, larger earthquakes 

tended to occur when the earthquake rate was high.

The dataset consisted largely of hybrid and low frequency 

(LF) earthquakes. These lower-frequency events typically had 

a low amplitude onset with frequencies up to about 10 or 20 

Hz, followed by a strong, long-duration, and lower-frequency 

waveform. The shape of the Fourier spectrum at the lowfrequency end was similar for both types of events, having 

significant energy between 0.3 and 5 Hz with multiple spectral 

peaks. Hybrid events exhibited more energetic high-frequency 

onsets than LF events, but the rate of high-frequency decay 

with time was similar for both types. The difference in highfrequency energy was more likely related to the source than to 

attenuation near the source or along the path, inasmuch as the 

high-frequency decay rate was the same for both events.

Layer 

Depth 

to top of 

layer (km)

S-wave 

velocity 

(km/s)

P-wave 

velocity 

(km/s)

Density 

(g/cm3

) QS QP

1 0 2.0 3.4 2.1 100 200

2 2 3.18 5.4 2.2 100 200

3 4 3.75 6.38 2.2 100 200

4 9 3.88 6.59 2.3 100 200

5 16 3.96 6.73 2.8 250 500

6 20 4.04 6.86 2.8 250 500

7 25 4.09 6.95 2.8 250 500

8 32 4.06 6.90 2.8 250 500

9 41 4.59 7.85 3.3 300 600

Table 2. Velocity structure of Mount St. Helens used in this study.

[For a material cycled in stress, Q is the inverse ratio of the energy lost each 

cycle to the peak strain energy. QP and QS specify the anelasticity of a material 

to P- and S-wave motion, respectively.]

The spectra of large and small events that occurred close 

in time were quite similar in shape between 0.5 and 10 Hz. 

The similarity in spectral shape suggested that both the source 

and path were similar for the larger and smaller events. As 

opposed to the ideal tectonic earthquake spectra, in which the 

peak in the velocity spectrum changes to lower frequency as 

rupture zone area increases for larger earthquakes, the frequency range of the peak of spectral energy remained constant 

for different-size, lower-frequency earthquakes. This relative 

constancy is consistent with a resonating source, whereby the 

shape of the spectrum is constrained by the size and properties of the resonator. However, a search of spectra of the same 

events observed at multiple stations revealed no specific frequencies of source resonance, and so the spectra may reflect a 

complex, perhaps asymmetric, resonating structure.

The spectra of a lower-frequency earthquake and an 

episode of volcanic tremor had distinct peaks in common. 

However, the lower-frequency earthquake spectrum had 

a prominent peak at 0.6 Hz that was not observed in the 

spectrum of the volcanic tremor, and higher frequencies were 

more favorably excited during the tremor. This lower-frequency peak would be consistent with resonance of a larger 

structure during the earthquake or a change in the source such 

that excitation of the longer dimension of the same structure 

was favorable.

The high-frequency spectral difference between hybrid 

and LF events was quantified using the slope of the highfrequency falloff. The slope computed for several thousand 

lower-frequency earthquakes between October 4 and December 1, 2004, defined a continuum between the ideal hybrid 

and ideal LF events. The slope started relatively high on 

October 4 and transitioned smoothly to lower values until 

around October 14. The gradual decrease in the slope of the 

high-frequency falloff over time reflected a progressive loss 

of high-frequency seismic energy as magma approached and 

eventually reached the surface.

The first onset of the surface-wave train produced by 

low-frequency earthquakes propagated at roughly 2 km/s. This 

propagation velocity favors an areally extensive surface layer 

with an average S-wave velocity of about 2 km/s. A window of 

the vertical and radial components for a 2 km/s group velocity has retrograde elliptical particle motion when observed 

at 20 km distance. Signal duration ranges from 15 to 50 s. A 

one-dimensional velocity model was unsuitable for waveform 

modeling to determine the source time function and focal 

mechanism of the lower-frequency earthquakes.

Acknowledgments

Deployment and maintenance of the seismic network was 

facilitated by numerous individuals at the U.S. Forest Service, 

the Cascades Volcano Observatory, and the Pacific Northwest 

Seismic Network of the University of Washington. Federal 

Express shipped our equipment to Portland, Oregon, without 

cost. We are grateful to Thomas Brackman, Chris Watson, 
 109

and Gary Patterson of the Center for Earthquake Research and 

Information (CERI) at the University of Memphis for taking 

time to assist in the field. We obtained seismic data for stations 

LON, JRO, and STD from the Incorporated Research Institutions for Seismology Data Management Center. Greg Waite 

and Charles Langston provided thoughtful and helpful reviews 

of this paper. This is CERI contribution 511.

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 6

Seismicity and Infrasound Associated with Explosions at 

Mount St. Helens, 2004–2005

By Seth C. Moran1, Patrick J. McChesney2, and Andrew B. Lockhart1

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Abstract

Six explosions occurred during 2004–5 in association 

with renewed eruptive activity at Mount St. Helens, Washington. Of four explosions in October 2004, none had precursory 

seismicity and two had explosion-related seismic tremor that 

marked the end of the explosion. However, seismicity levels 

dropped following each of the October explosions, providing 

the primary instrumental means for explosion detection during the initial vent-clearing phase. In contrast, explosions on 

January 16 and March 8, 2005, produced noticeable seismicity in the form of explosion-related tremor, infrasonic signals, 

and, in the case of the March 8 explosion, an increase in event 

size ~2 hours before the explosion. In both 2005 cases seismic 

tremor appeared before any infrasonic signals and was best 

recorded on stations located within the crater. These explosions demonstrated that reliable explosion detection at volcanoes like Mount St. Helens requires seismic stations within 

1–2 km of the vent and stations with multiple acoustic sensors.

Introduction

On September 23, 2004, a swarm of volcano-tectonic 

earthquakes heralded the reawakening of Mount St. Helens 

after 18 years of quiescence (Moran and others, this volume, 

chap. 2). On October 1 the first small explosion occurred, 

with three others following over the next four days (table 1). 

Several of these happened during daylight hours and were 

broadcast on live television across the United States, perhaps 

creating an impression that Mount St. Helens was building 

towards a larger explosive eruption. Instead, these explosions 

were followed by the steady-state extrusion of a new dome in 

the southern part of the 1980 crater. The relatively steady-state 

extrusion was punctuated by just two explosions, one on January 16, 2005, and the other on March 8, 2005 (table 1). All 

six explosions were phreatic, with no evidence, such as fresh 

pumice, of any significant magmatic component (Scott and 

others, this volume, chap. 1; Rowe and others, this volume, 

chap. 29). Overall, the eruption in 2004–5 featured very little 

explosive activity, a result of the gas-poor nature of the erupted 

magma (Gerlach and others, this volume, chap. 26).

We use the term “explosion” in this paper to refer to an 

impulsive, sudden yet sustained emission of volcanic gas and 

pyroclasts. Rapid detection of ash-producing explosions is of 

paramount importance because of the demonstrated dangers 

posed to aircraft by airborne ash particles (for example, Neal 

and others, 1997). Despite the paucity of explosions, seismic 

and acoustic recordings of the six explosions at Mount St. 

Helens have yielded several important insights into the utility 

and placement of seismic and acoustic sensors for improved 

explosion-detection capabilities, as well as insights into the 

nature of minor explosive activity during dome-building eruptions. In this paper we present a chronology of observations 

and recordings of explosions from 2004–5, focusing on our 

ability to detect the onset and termination of each explosion 

and on insights gained from seismic and acoustic data into the 

evolution of individual explosions.

Explosions During the Vent-Clearing 

Phase, October 1–October 5, 2004

Explosion 1: October 1, 1202 PDT

The first explosion occurred without warning at 1202 

PDT on October 1 (fig. 1A, table 1), 8.5 days after the start of 

seismic unrest (Moran and others, this volume, chap. 2). The 

explosion was well documented, as it took place on a clear day 

2

 Pacific Northwest Seismic Network, Department of Earth and Space 

Sciences, University of Washington, Box 351310, Seattle, WA 98195
112 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

with a U.S. Geological Survey (USGS) crew flying over the 

crater at the start of the explosion (Schneider and others, this 

volume, chap. 17), and many television cameras were transmitting images live from a site near Johnston Ridge Observatory 

(JRO) (fig. 2) to a nationwide audience (Driedger and others, 

this volume, chap. 24). Infrared recordings of the base of the 

eruption column showed peak temperatures around 160°C, 

consistent with visual observations that the explosion was 

phreatic (Schneider and others, this volume, chap. 17). Eruption cloud tops reached ~4,500 m above sea level (asl), or 2,400 

m above the vent.

As described by Moran and others (this volume, chap. 2), 

seismicity had intensified in a series of steps through the first 

Explosion 

number, 

date, and 

time 

Duration 

(minutes)

Seismicity before, 

during, and after

Infrasound 

detected?

Plume height 

(approximate, 

m above vent)

Eruption

phase

Explosion 1: 

10/01/04, 

1202 PDT

19 No precursors; earthquakes 

stopped following onset; 

tremor associated with end 

of explosion; almost no 

earthquakes for several hours 

afterwards

N/A (no sensors) 2,400 Vent clearing

Explosion 2: 

10/03/04, 

2240 PDT

~25 No precursors; earthquake rates 

decreased after onset (drop not 

as pronounced as explosion 

1); tremor associated with end 

of explosion; seismicity low 

following explosion

Possible, but could 

also be coseismic 

shaking

400 Vent clearing

Explosion 3: 

10/04/04, 

0943 PDT

~32 No precursors; earthquake rates 

gradually declined after onset 

(drop not as pronounced as 

explosion 2); no tremor associated with end of explosion

No 1,500 Vent clearing

Explosion 4: 

10/05/04, 

0905 PDT

70 No precursors; gradual but 

significant decline in earthquake rate and size after onset; 

no tremor associated with end 

of explosion

No 2,400 Vent clearing

Explosion 5: 

01/16/05, 

0312 PST

~33 No precursors; tremor occurred 

during entire event; no change 

in earthquake rate observed 

after explosion

Yes Unknown Dome building

Explosion 6: 

03/08/05, 

1725 PST

~20 Earthquake magnitudes increased ~2 hours before onset; 

tremor associated with entire 

event; short-lived increase in 

earthquake rate after explosion, 

rate returned to normal within 

1–2 hours

Yes 9,000 Dome building

Table 1. Summary of eruption parameters and seismic and acoustic observations of the six explosions 

occurring at Mount St. Helens, Washington, during 2004–5.

[PDT, Pacific daylight (saving) time; PST, Pacific standard time.]

several days of the crisis that began on September 23, 2004 

(fig. 3). By September 29, earthquakes of magnitude 2 and 

greater were occurring at a rate of ~1 per minute. Although 

seismicity intensified gradually between 0800 and 1100 PDT 

on October 1, several more-rapid intensifications had occurred 

in the previous two days, and so this intensification was not 

recognized as a short-term warning sign. Because preexplosion 

seismicity was so intense, it was difficult to see any obvious 

seismicity associated with the explosion in the first several 

minutes after 1202 PDT (fig. 1A). The most obvious seismic 

signal was the sudden cessation of earthquakes ~1 minute after 

the explosion began. After the earthquakes stopped, explosionrelated tremor (defined here as seismic tremor accompanying 
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 113

Figure 1. Multistation 2-hour-long spectrograms for explosions 1–4 at Mount St. Helens, Washington, October 2004. Time 

series (top) and frequency spectrogram (bottom) are shown for each of four seismic stations (see fig. 2 for station locations).

Stations are shown in order of distance from vent, with closest station at top of the plot. Each time series is normalized to the 

maximum amplitude in each 2-hour window. Frequency spectrograms show spectral amplitudes for frequencies ranging from 

0 to 10 Hz using a rainbow color palette, with dark blue corresponding to low spectral amplitudes (< 30 dB) and red to high 

amplitudes (>100 dB). Explosion onset time indicated by an arrow at top of each spectrogram. A, October 1, 2004, 1100–1300 

PDT (explosion at 1202). B, October 3, 2004, 2200–2400 PDT (explosion at ~2240). C, October 4, 2004, 0900–1100 PDT (explosion at 

0943). D, October 5, 2004, 0800–1000 PDT (explosion at 0905).

an explosion) with frequencies of 1–3 Hz and a peak reduced 

displacement, or DR (Aki and Koyanagi, 1981), of ~6 cm2 could 

clearly be seen on many nearby stations (fig. 1A; see fig. 2 

for station locations). Tremor began increasing in amplitude 

at 1216, then abruptly stopped at 1221, when the explosion 

also stopped. Weak tremor occurred again starting at 1240, but 

earthquakes did not occur again until ~3 hours later (Moran and 

others, this volume, chap. 2).

If the earthquakes had not ceased when the explosion 

started and the explosion had not occurred during daylight 

hours in clear conditions, it is possible that staff at the Cascades Volcano Observatory (CVO) and the Pacific Northwest 

Seismic Network (PNSN) might not have realized that an 

explosion had occurred. The intense earthquake activity at the 

time was continuously saturating nearby seismic stations, all 

of which had short-period vertical-component seismometers, 

and the explosion-related tremor would have been almost 

impossible to distinguish had the earthquakes not ceased. In 

an attempt to improve our explosion detection capability, 1-Hz 

infrasound-sensitive acoustic sensors (see appendix 1) were 

1100 1124 1148 1212 1236 1300

5

10

5

10

5

10

5

10

CDF

0800 0824 0848 0912 0936 1000

5

10

STD

5

10

SHW

5

10

YEL

5

10

TIME (PDT), OCTOBER 5, 2004

TIME (PDT), OCTOBER 1, 2004 TIME (PDT), OCTOBER 3, 2004

2200 2224 2248 2312 2336 2400

CDF STD SHW YELCDF STD SHW YEL

0900 0924 0948 1012 1036 1100

CDF

5

10

STD

FREQUENCY, IN HERTZFREQUENCY, IN HERTZ

FREQUENCY, IN HERTZFREQUENCY, IN HERTZ

5

10

SHW

5

10

YEL

5

10

TIME (PDT), OCTOBER 4, 2004

5

10

5

10

5

10

5

10

Data Gap

Tremor Tremor

A B

C D
114 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Steam-and-ash explosion

Tremor, no explosion

0

200

400

600

800

1,000

1,200

1,400

1,600

09/22 09/23 09/24 09/25 09/26 09/27 09/28 09/29 09/30 10/01 10/02 10/03 10/04 10/05 10/06 10/07

DATE (PACIFIC DAYLIGHT TIME)

10-MINUTE RSAM, IN COUNTS

Swarm

onset

JRO

SHW

HSR

JUN

STD

SUG

Coldwater

46°15′

46°00′

122°22′30′′ 122°15′ 122°7′30′′ 122°00′

46°7′30′′

46°22′30′′

0 5 10 km

WASHINGTON

AREA OF MAP

BLIS

NED

MIDE

SEP

YEL

500 m

Figure 3. Plot of 10-minute RSAM values (Endo and Murray, 1991) for station SHW (see fig. 2) from September 22 to 

October 8, 2004. Black arrows indicate timing of individual explosions; white arrows, timing of two tremor episodes that did 

not correspond to any explosions (the only two such cases during entire eruption; Moran and others, this volume, chap. 2).

Figure 2. Station-location map showing 

all seismic and acoustic sensors jointly 

operated by the Pacific Northwest Seismic 

Network and Cascades Volcano Observatory 

at Mount St. Helens. Inset map shows 

stations plotted on a digital elevation model 

from April 19, 2005 (Schilling and others, this 

volume, chap. 8). Black triangles correspond 

to stations operating before 2004, white 

triangles to stations installed after the start 

of the eruption in 2004. All stations discussed 

herein are labeled. Stations SEP and NED 

are on the 1980–86 lava dome, MIDE and 

BLIS are near or on parts of the newly 

erupted lava dome.
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 115

installed at the Studebaker Ridge station (STD) on October 2 

and at Sugar Bowl (SUG) on October 3 (McChesney and others, this volume, chap. 7; see fig. 2 for station locations).

Explosion 2: October 3, ~2240 PDT

The second explosion occurred at ~2240 PDT on October 3 (fig. 1B). Although this explosion occurred at night, the 

weather was clear and U.S. Forest Service observers stationed 

13 km north-northwest of the crater at the Coldwater Ridge 

Visitor Center reported seeing an ash cloud that barely reached 

the crater rim (~2,500 m asl, or 400 m above the vent) starting at ~2240 PDT (M. Guffanti, written commun., 2004). 

Earthquake rates and real-time seismic amplitude measurement (RSAM; Endo and Murray, 1991) levels had increased 

several hours before the explosion (fig. 3), and by 2100 PDT 

events were occurring so close together that they were difficult to distinguish, forming a spasmodic, tremorlike signal 

(Moran and others, this volume, chap. 2). Because there had 

been much more energetic tremor earlier on October 2 and 

October 3 with no associated explosions, the spasmodic tremor 

was not regarded as a short-term warning sign. As with the 

October 1 explosion, earthquake rates dropped significantly 

at ~2305 PDT following a ~3-minute-long tremor episode 

(peak DR ~3 cm2) that presumably marked the end of the explosion (fig. 1B). A weak, continuous acoustic signal appeared on 

the newly deployed microphone at STD. However, with just 

that single sensor, we cannot distinguish between mechanical shaking of the microphone by passing seismic waves 

and explosion-related infrasound. The STD microphone did 

record many small (~0.5 Pa) infrasonic pulses associated with 

M >2 earthquakes, indicating that it was sufficiently close to 

record weak infrasonic signals from the crater.

Explosion 3: October 4, 0943 PDT

Earthquake rates gradually increased on the morning of 

October 4 until the third explosion occurred at 0943 PDT (maximum plume height ~3,650 m asl, or about 1,500 m above the 

vent), lasting until 1015 (fig. 1C). The explosion was recorded 

by a U.S. Forest Service Web camera at the JRO and was also 

noted by many observers. Earthquake rates decreased following 

the explosion (fig. 3), although not as markedly as either of the 

decreases following the previous two explosions. In contrast to 

the October 3 explosion, no obvious signal was apparent on the 

acoustic sensors and no explosion-related tremor was recorded 

on the nearby seismic stations (fig. 1C). The decline in seismicity was gradual and, unlike that accompanying the October 1 

explosion, did not correspond in any direct way to the onset or 

termination of the explosion. Given the lack of obvious associated seismic signals, it is likely that the explosion would not 

have been detected had it occurred during bad weather.

Dzurisin and others (2005) show in their figure 1 a second 

steam-and-ash explosion on October 4 at ~1400 PDT. This was 

based on a report of anomalous steaming in the crater at that 

time (M. Guffanti, written commun., 2004), a report that coincided with a small drop in RSAM values. However, subsequent 

review of images from the JRO Web camera (taken every five 

minutes; Poland and others, this volume, chap. 11) showed no 

obvious steam or ash plume at that time, in contrast to the 0943 

explosion. Given the absence of visible ash, we consider the 

1400 event, if there was an event at all, to be at most a small 

steam explosion that is not comparable to the confirmed explosions, and we do not consider it further in this paper.

Explosion 4: October 5, 0905 PDT

The fourth and final explosion of the vent-clearing phase 

occurred at 0905 PDT on October 5 and lasted until 1015 (fig. 

1D). This was the most vigorous and long-lasting explosion of 

the sequence, with the ash plume reaching ~4,500 m asl (2,400 

m above the vent) and depositing trace amounts of ash ~100 

km from the volcano (Scott and others, this volume, chap. 

1). Before the explosion, RSAM levels had increased over a 

6-hour period, reaching a peak level at 0600 that was maintained until the explosion (fig. 3). Earthquake sizes and rates 

began declining ~15 minutes after the explosion began (fig. 

1D), with RSAM levels falling below post-September 29 levels by the end of the explosion (fig. 3). As with the October 4 

explosion, no obvious explosion-related signals were apparent 

either on seismic or acoustic sensors (fig. 1D). However, the 

decline in seismicity was significant enough that an explosion 

could have been inferred if weather conditions had prohibited 

observation of the explosion. A delayed indicator of the explosion was the loss of the radio signal from station YEL from 

1045 to 1238 following the explosion as a result of attenuation 

of its radio signal by the ash cloud.

Discussion

Because there was no attempt to maintain a full-time 

official observer near the volcano, there is a remote possibility 

that other small explosions occurred between October 1 and 

5 that were not detected. We feel confident that no undetected 

explosions occurred during daylight hours, as clear weather 

provided excellent viewing conditions for the mass of people 

and media watching the volcano from various vantage points 

(Driedger and others, this volume, chap. 24). During nighttime 

hours clear viewing conditions still existed, aided by moonlight from an almost-full moon, and, as a result, members of 

the public and U.S. Forest Service staff were able to see one 

nighttime explosion (October 3). Nevertheless, we cannot rule 

out the possibility that other explosions occurred at night when 

the volcano was not watched.

Short-term declines in seismic energy following all four 

explosions were perhaps the most reliable indicator that an 

explosion had occurred. Seismicity declines following the 

explosions on October 1, 3, and 5 were particularly significant (fig. 3). However, similarly significant declines followed 

tremor episodes on October 2 and 3, which were not associ-
116 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ated with explosions (Moran and others, this volume, chap. 

2). In addition, smaller declines similar to that following the 

October 4 explosion (including several on October 4) did not 

correspond to known explosions. Thus short-term declines in 

seismic energy were not, by themselves, a reliable indicator 

that an explosion had occurred.

Explosions During the Dome-Building 

Phase

Several days after the October 5, 2004, explosion, a lava 

spine emerged from the vent (Vallance and others, this volume, chap. 9). The rest of 2004 was dominated by lava dome 

construction accompanied by low gas levels (Gerlach and others, this volume, chap. 26) and regularly spaced earthquakes 

(Moran and others, this volume, chap. 2). During October, 

November, and December of 2004 we installed several seismic 

stations within 500 m of the vent (fig. 2) and disabled the 

microphone at STD, because we needed the radio telemetry 

channel for data from another station (McChesney and others, 

this volume, chap. 7). One of the new stations on the 1980–86 

lava dome (SEP) had a 2-Hz three-component velocity sensor and two 1-Hz acoustic sensors spaced ~15 m apart in a 

north-south alignment (appendix 1), roughly radial to the vent. 

With two microphones we hoped to be able to use relative 

arrival times of signals between the two sensors to distinguish 

between wind gusts (which would not necessarily produce 

similar waveforms, but any similar waveforms would have 

separations of as much as several seconds between the two 

sensors), coseismic signals due to shaking of the microphones 

by passing seismic waves (which would vary depending on the 

coupling of each microphone to the ground), and infrasonic 

signals (which would produce very similar waveforms with no 

more than ~0.05 s difference between the two sensors).

Explosion 5: January 16, 2005, ~0312 PST

The dome-building eruption was punctuated by a relatively short-lived explosion at ~03:12:50 PST on January 16, 

2005. The explosion occurred at night during poor weather 

and was not visually observed. It was instead signaled by the 

sudden onset of a continuous broadband (1–10 Hz) tremor 

signal (fig. 4) accompanied by several larger-than-average 

low-frequency (dominant frequency <5 Hz) seismic events. 

The explosion-related tremor was relatively small (peak DR of 

~0.5 cm2) and did not show up well on stations outside the 

crater (fig. 4). As a result, preestablished amplitude-based 

thresholds for generating automated alarms were not exceeded 

(Qamar and others, this volume, chap. 3) and the explosion was 

not detected by CVO and PNSN staff until more than an hour 

later during a routine scan of seismic records.

The explosion signal initially was most obvious on station BLIS (fig. 5), located ~250 m east of the vent (fig. 2). The 

estimated start time for the explosion is based on the onset of 

tremor at BLIS. Tremor did not become obvious on other stations, including two stations located ~500 m from the vent, until 

it increased in amplitude starting at ~0318 (figs. 4, 5). As BLIS 

was close to the actively growing spine 4, the tremor signal 

could conceivably have been caused by rockfalls coming off 

the spine. Rockfalls at Mount St. Helens commonly produce 

spindle-shaped signals 1–2 minutes long that often only appear 

on nearby stations. The tremor signal on BLIS was continuous 

between 0312 and 0318, however, indicating that the signal was 

the result of a longer duration process. We speculate that this 

signal could have been caused by relatively weak jetting before 

the more significant explosion that presumably occurred in association with the increase in tremor amplitudes starting at 0318.CDF

0200 0212 0224 0236 0248 0300 0312 0324 0336 0348 0400

5

10

STD

5

10

SHW

5

10

YEL

5

10

NED

5

10

SEP

5

10

BLIS

5

10

PACIFIC STANDARD TIME, JANUARY 16, 2005

Tremor

FREQUENCY, IN HERTZ

Figure 4. Multistation 2-hour-long spectrogram from 0200 to 

0400 PST on January 16, 2005. See figure 2 for station locations 

and caption to figure 1 for description of spectrograms. Blacktipped arrow indicates estimated start time of explosion at 0312 

PST; duration of tremor as seen on station SEP is indicated. 

Note that stations BLIS and NED had accelerometer sensors, 

SEP had a 2-Hz velocity sensor, and all other stations had 1-Hz 

velocity sensors. For more details on configuration of seismic 

stations, see McChesney and others (this volume, chap. 7).
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 117

BLIS

SEP

SEP

mic 1

SEP

mic 2

SUG

mic 1

NED

YEL

0310 0315 0320 0325 0330 0335 0340

PACIFIC STANDARD TIME

1,000

-1,000

1,000

-1,000

-200

200

-500

500

2,000

-2,000

1,000

-1,000

2,000

-2,000

SEP

SEP

mic 1

SEP

mic 2

AMPLITUDE, IN COUNTS

AMPLITUDE, IN COUNTS

200 s

-400

400

-200

200

-500

500

03:21:40 03:24:40 22:40 23:40

20 s

-400

400

-50

50

-200

200

03:26:30 26:40 03:26:50 26:35 26:45

1 s

Tremor rapidly intensified between 0318 and 0320. At 

~0320 an emergent continuous signal became apparent on the 

SEP and SUG microphones (fig. 5). Given that these signals 

appeared in association with the seismic tremor, it is possible 

that some, if not most, of this signal reflects mechanical shaking of the microphones by seismic waves. We have no means 

of assessing this at the SUG site because it had just a single 

microphone and no collocated seismometer, but the near-simultaneous onset of acoustic signals at SEP and SUG suggest that 

at least a component of the SUG signal was due to shaking of 

the microphone housing. At 03:20:30, however, the two acoustic 

sensors at SEP began recording small-amplitude individual 

phases within the acoustic signals that were coherent between 

the two microphones and consistent with a source located at the 

vent. No correlative signals were apparent on the SUG acoustic 

sensor ~5 s later (the traveltime difference between SEP and 

SUG for acoustic waves traveling at 340 m/s). However, this is 

not surprising, given the fact that SUG is located on a ridgetop 

that was being buffeted by high winds from an oncoming storm 

system at the time. We infer that the coherent signals on the two 

acoustic sensors at SEP reflect the onset of explosion-related 

infrasound. Although the microphone recordings were still 

mostly incoherent at the onset of the continuous acoustic signal 

(fig. 5), individual phases in the signal became progressively 

more coherent with time. By ~0326, most of the signal was 

coherent and thus likely due to explosion-related infrasound (as 

opposed to explosion-related air currents or mechanical shaking of the sensors). The continuous acoustic signal peaked at 

~0327 with maximum peak-to-peak amplitudes of 2–4 Pa (see 

appendix 1 for discussion of microphone calibration uncertainties) and faded to background levels by ~0331. The timing, size, 

duration, and character of the SEP infrasonic signals are consistent with an emergent infrasonic signal starting at 03:21:22 

recorded on a microbarometer array located 13 km northwest of 

Figure 5. Multistation 30-minute plot showing unfiltered time-series data from 0310 to 0340 PST on 

January 16, 2005. See figure 2 for station locations. Insets at bottom of plot show 3 minutes (left) and 20 s 

(right) of data from vertical component of SEP seismometer (top) and two acoustic sensors (microphone 

1, middle, and microphone 2, bottom), bandpass filtered with corner frequencies of 0.1 and 15 Hz. Left 

inset shows two events, recorded by acoustic sensors at ~03:22:00 and ~03:23:50, that we infer to be 

wind gusts associated with the explosion. Right inset shows mostly coherent signals between the two 

microphones, indicating that signals are mostly infrasonic energy produced by the explosion.
118 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

the vent near the U.S. Forest Service’s Coldwater Ridge Visitor 

Center (Matoza and others, 2007).

Superimposed on the continuous infrasonic signal at SEP 

were several larger ~1 Hz pulses at 03:20:00, 03:22:05, and 

03:23:50 (fig. 5). Although these pulses were broadly coherent between the two microphones, in detail the waveforms 

have significant variability. In addition, the time separation 

between relative peaks (as much as 2 s) is too large for an 

infrasonic signal, given the 15-m spacing of the microphones. 

The variability and large time differences between these pulses 

on the two microphones indicate that they were not traveling 

at sound velocities (~340 m/s) and are thus not infrasound. 

Other such signals were recorded in the hours before and after 

the explosion on both the SEP and SUG acoustic sensors, as 

well as at other times during stormy weather. Because a storm 

system was moving into the area at the time of the explosion, 

these pulses are likely wind gusts unrelated to the explosion. 

The 03:22:05 and 03:23:50 pulses, however, had larger-thanaverage amplitudes and broadly similar waveforms (fig. 5). 

On the basis of the admittedly subtle contrasts to other wind 

gust-related signals, we speculate that these pulses may have 

been caused by air flow from the explosion that swept across 

the SEP site at speeds of ~20 m/s.

Evidence that ash was moving across the 1980–86 lava 

dome shortly after the tremor-amplitude increase at 0318 

comes from a several-second-long loss (or “dropout”) of the 

radio signal from NED at 03:22:20 (fig. 5; see fig. 2 for site 

location). The radio signal from NED was known to be very 

strong. At the radio receiver for NED we measured a 30 dB 

fade margin—a quality determined by adding impedance at 

the receiver until it stops receiving the radio signal from the 

transmitter. The NED signal had never before been lost, even 

during stormy weather. The most likely explanation for the 

NED dropout is that the radio signal was temporarily blocked 

by ash. Radio telemetry was lost ~60 s later at BLIS, which 

had a much weaker radio signal (~10 dB fade margin at the 

receive site) and commonly dropped out during winter storms. 

Signals from both stations returned in time to record the peak 

of the tremor signal at 0327 (fig. 5). Shortly after this peak 

there were two more short-lived radio dropouts at 0328 and 

0329 from station NED, followed by a dropout of several 

hours at station BLIS. The long duration of the BLIS dropout 

is consistent with a diffuse ash cloud lingering in the vent area 

following the explosion. Given that no dropouts occurred at 

SEP and that the first NED dropout occurred before the first 

BLIS dropout (despite BLIS being located much closer to the 

vent), we infer that ash was initially blown northeastwards 

across the dome towards NED.

Explosion-related seismic tremor gradually subsided 

after 0327, eventually fading to background levels at ~0345 

PST on the crater stations. The tremor signal lasted for 

roughly 32 minutes, similar to the duration of the October 

1, 2004, explosion. Subsequent geological reconnaissance 

on January 19 (3 days later but the first day of clear weather 

after the explosion) confirmed that an explosion had occurred 

(Scott and others, this volume, chap. 1), with the primary 

axis of ash deposition extending east-northeast from the vent 

toward both NED and BLIS (fig. 2). A field of impact craters 

as large as 1 m in diameter extended several hundred meters 

eastwards from the vent towards and beyond BLIS. Station 

BLIS survived in large part because it was mostly buried in 

snow. The distribution and thickness of deposits were similar 

in scope and size to those from the October 1, 2004, explosion (Scott and others, this volume, chap. 1), and it is reasonable to assume that the plume reached heights similar to the 

October 1 plume (~3,500–4,500 m asl).

In contrast to the October 1 explosion, nearby seismic 

stations were not saturated by large earthquakes before the 

January 16 explosion; as a result, the explosion-related tremor 

could clearly be seen on stations within the crater. Seismic 

signals from the explosion showed up poorly on stations outside the crater, however (fig. 4). If there had been no seismic 

stations inside the crater at the time of the explosion, it is 

conceivable that CVO and PNSN staff would not have known 

that an explosion had occurred until the next observation 

flight in the crater. This highlights the importance of having 

stations close to the vent for accurate and timely detection of 

explosions. In particular, the loss of telemetry on several crater 

stations, coupled with the tremor signal, made it clear that an 

ash-producing explosion had occurred.

As described above, the explosion produced infrasound 

that was recorded at SEP and at the Coldwater microbarometer 

array (Matoza and others, 2007). Had there been no crater stations, the Coldwater array might have provided the only definitive instrumental evidence that an explosion had occurred. 

The infrasonic signal was relatively subtle at both sites, 

however, and required multiple collocated acoustic sensors to 

distinguish between explosion-related infrasonic signals, wind 

noise, coseismic shaking of the acoustic sensors, and other 

acoustic noise sources. The subtlety of the infrasonic signals 

illustrates the important role that arrays of acoustic sensors can 

play in detecting explosions at volcanoes. Given the relatively 

small infrasonic signals, this explosion may represent an 

example of the type of ash-rich explosion that elsewhere has 

been found to be relatively inefficient at producing infrasound 

(Woulff and McGetchin, 1976; Johnson and Aster, 2005).

Explosion 6: March 8, 2005, 1725 PST

The dome-building eruption was not at all disturbed by 

the January 16 explosion, with earthquakes and steady-state 

lava-dome extrusion continuing unabated for the next seven 

weeks. The steady-state lava-dome extrusion was again punctuated by an explosion at ~17:25:20 PST on March 8, 2005. 

This was the largest of the six 2004–5 explosions; it was also 

the best documented and recorded of the six, as it occurred in 

the early evening of a cloudless day and as a result could be 

seen from the Portland, Oregon, metropolitan area, 85 km to 

the south. Tremor associated with the explosion was visible on 

all stations within 15 km (fig. 6), and CVO personnel who had 

seen the signal on seismic displays began contacting Federal, 

airline, and emergency officials within a minute of the start of 
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 119

the explosion. Explosion-related tremor appeared simultaneously at stations inside and outside the crater and achieved 

maximum amplitudes (peak DR ~1 cm2) within the first minute, indicating that the explosion rapidly reached maximum 

intensity, in contrast to the January 16 explosion. Visual observations confirmed the rapid intensification, with the resultant 

ash plume reaching heights of about 11,000 m (9,000 m above 

the vent) within 5 minutes (Scott and others, this volume, 

chap. 1). Explosion-related tremor began declining after ~8 

minutes and became indistinct from normal seismic background levels within 20 minutes. All three stations within 500 

m of the vent were destroyed by the explosion (fig. 7).

CDF

1500 1518 1536 1554 1612 1630 1648 1706 1724 1742 1800

5

10

STD

5

10

SHW

5

10

YEL

5

10

NED

5

10

SEP

5

10

MIDE

5

10

PACIFIC STANDARD TIME

Stations destroyed

FREQUENCY, IN HERTZ

Figure 6. Multistation 3-hour-long spectrogram from March 

8, 2005, between 1500 and 1800 PST. See figure 2 for station 

locations and caption to figure 1 for description of spectrograms. 

White-tipped arrow indicates start of an increase in number 

of larger events ~2 hour before the March 8 explosion; blacktipped arrow indicates start of explosion at 1725 PST. Note that 

stations MIDE and NED had accelerometer sensors, SEP had 

a 2-Hz velocity sensor, and all other stations had 1-Hz velocity 

sensors. For more details on configuration of seismic stations, 

see McChesney and others (this volume, chap. 7).

RSAM Increase

Roughly 2 hours before the March 8 explosion, RSAM 

levels began increasing at stations as far as 6 km from the 

vent (fig. 8). Although the increase was relatively small, 

particularly in comparison to increases seen in the first 

week of the eruption (Moran and others, this volume, chap. 

2), the synchronous rise in RSAM at stations within and 

outside of the crater was unusual enough for CVO and PNSN 

staff to take notice roughly an hour before the explosion. 

This increase was one of the factors that contributed to the 

timely detection of the explosion, and it may be a relatively 

rare example of a short-term seismic precursor to an explosion. The RSAM increase was caused by an increase in the 

percentage of larger (Md >1.5) earthquakes (fig. 9). Events 

in this magnitude range had been occurring at a rate of one 

every 3– 4 minutes before the explosion. Starting at ~1529 

the rate of larger events increased to one every 1–2 minutes, 

with the rate increasing through to the start of the explosion. No other attributes of the seismicity, including the total 

number of earthquakes per unit time (fig. 9), event frequency, 

event type, event location, and degree of similarity of waveforms between events, changed at this time.

Similar increases in event size occurred over short time 

intervals during the first week of seismic unrest in 2004. The 

fact that seismic energy levels dropped following the four 

explosions (particularly the October 1 and October 5, 2004, 

explosions) and two noneruptive tremor episodes indicates that 

the preexplosion elevated seismicity was a result of increased 

pressures in the shallow (<1 km) conduit system (Moran and 

others, this volume, chap. 2). We infer that the increase in seismic energy ~2 hours before the March 8 explosion similarly 

reflected an increase in pressure within the conduit. Given 

that an explosion occurred 2 hours after this increase, it is 

reasonable to assume that a pocket of steam and/or magmatic 

gas had accumulated within the conduit or along its margins 

at shallow depths shortly before the explosion, and that the 

gas pocket locally increased pressures. Given that there were 

no other significant changes in earthquakes before the explosion, we infer that the pressure increase merely perturbed the 

regular seismogenic process. Certainly, the increase in event 

size could reflect an increased pressurization in gas- or fluidfilled cracks (for example, Chouet, 1996). However, other 

factors, such as the ~1-m-thick layer of fault gouge found on 

most spines (Pallister and others, 2005; Cashman and others, this volume, chap. 19; Moore and others, this volume, 

chap. 20); the geologic evidence found for shearing within the 

extruded lava domes (Pallister and others, 2006); the low gas 

content (Gerlach and others, this volume, chap. 26); and the 

correlation of changes in earthquake character with changes in 

extrusive style at the surface (Moran and others, this volume, 

chap. 2), all combine to suggest that the regular earthquakes 

may have been a result of stick-slip motion (Iverson and others, 2006). If the earthquakes were the result of a stick-slip 

process, the pressure increase could have caused a seismic 

energy increase through (1) an increase in pressure at the base 
120 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

MIDE

SEP

SEP

mic 1

SEP

mic 2

NED

YEL

17:20:00 17:22:30 17:25:00 17:27:30 17:30:00

PACIFIC STANDARD TIME

1,000

-1,000

1,000

-1,000

-1,000

500

-1,000

1,000

2,000

-2,000

1,000

-1,000

AMPLITUDE, IN COUNTS

SEP

SEP

mic 1

SEP

mic 2

17:26:10 17:26:50 26:20 26:30 26:40

50 s

10 s

-1,000

1,000

-100

100

100

-100

17:27:15 17:27:55 27:25 27:35 27:45

-1,000

1,000

500

-500

500

AMPLITUDE, IN COUNTS

-500 10 s

Figure 7. Multistation 

10-minute plot showing 

time-series data from 1720 

to 1730 PST on March 

8, 2005. See figure 2 for 

station locations. Insets 

at bottom of plot show 

two 40-s windows of 

unfiltered data from the 

vertical component of the 

SEP seismometer (top) 

and two acoustic sensors 

(microphone 1, middle, and 

microphone 2, bottom). 

The left inset shows 

the onset of infrasonic 

signals at ~17:26:20, with 

amplitudes increasing 

significantly at ~17:26:40. 

Right inset shows several 

sharp spikes on acoustic 

sensors, which we infer 

to be caused by ballistic 

fragments landing close to 

each sensor. Conversion 

factor for counts to 

Pascals is nominally 44.5 

counts/Pa for microphones 

at SEP (see appendix 1).

of the plug (Iverson and others, 2006); (2) a localized increase 

in applied shear stress along the conduit margins; (3) a change 

in fault properties in the seismogenic region; or (4) some 

combination of these factors. Because there are no constraints 

on the rate of motion of the active spine or the location of the 

gas pocket in the hours before the explosion, there is no basis 

for favoring or discarding any of these explanations. Nevertheless, the increase in RSAM values shortly before the March 

8 explosion represents a relatively rare instance of a seismic 

precursor to a volcanic explosion and provides an example 

of a signal that could enable future short-term forecasts of 

explosive events during eruptions similar to the dome-building 

eruption at Mount St. Helens during 2004–5.

Acoustic Recordings

In contrast to the January 16 explosion, the March 8 

explosion produced significant infrasonic signals that were well 

recorded on the two SEP microphones (the SUG microphone 

had stopped working before the March 8 explosion) as well as 

on the Coldwater infrasound array (Matoza and others, 2007). 

No infrasonic signals were recorded, however, until a discrete 

low-frequency (1–2 Hz) pulse at 17:26:20, 

~60 s after the explosion signal first appeared on seismic stations (fig. 7). We note that the delay between seismic and infrasonic signals is similar to seismicity and infrasonic observations 

reported by Johnson and others (2003) in association with 

explosions at Guagua Pichnicha volcano, Ecuador, in 1998–99. 

These first infrasonic pulses were low frequency (1–2 Hz) and 

appeared on the SEP acoustic sensors ~2 s after a seismic event 

was recorded on the SEP seismometer (fig. 7), consistent with 

the time lapse expected between seismic and acoustic waves for 

a source ~500 m away. Matoza and others (2007) report seeing 

similar infrasonic signals at their Coldwater array starting at 

17:26:55, 35 s after they appeared on the SEP acoustic sensors 

(35 s is the expected time difference, given the 12.5-km separation between the two sites). Several similar pulses occurred 

over the next 20 s, followed by much larger infrasonic signals 
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 121

100

150 MIDE

25

75

SEP

50

100

NED

75

175 SHW

75

125

3/7/05

1200

3/8/05

0000

3/8/05

1200

3/9/05 

0000

JUN

10-MINUTE RSAM, IN COUNTS

DATE AND TIME (PACIFIC STANDARD TIME)

Explosion

 (1725)

0

4

8

12

16

1.8

2.2

2.6

3/8/05

0000

3/8/05

0300

3/8/05

0600

3/8/05

0900

3/8/05

1200

3/8/05

1500

3/8/05

1800

3/8/05

2100

3/9/05

0000

NUMBER OF EVENTS

 PER 10 MINUTES

CUMULATIVE MAGNITUDE

PER 10 MINUTES

Explosion

(1725)

DATE AND TIME (PACIFIC STANDARD TIME)

Figure 8. Plot of 10-minute 

RSAM values (Endo and 

Murray, 1991) on five stations 

over 36-hour time window 

extending from March 7 to 

March 9, 2005. Stations are 

ordered by distance from vent, 

nearest (MIDE, 0.25 km) to 

farthest (JUN, ~6.5 km). Vertical 

dashed line indicates start time 

for March 8 explosion.

Figure 9. Plot showing 

number of events detected 

every 10 minutes at station 

HSR (top) and 10-minute 

cumulative event 

magnitude (bottom). Event 

detection is done using a 

standard STA/LTA trigger 

algorithm; for details see 

Moran and others (this 

volume, chap. 2). Event 

magnitudes determined 

using peak amplitudes 

at station HSR and a 

local magnitude relation 

calibrated to match the 

Pacific Northwest Seismic 

Network’s coda-duration 

magnitude scale. Note 

that magnitudes began 

increasing slightly after 

1500 PST, whereas number 

of events stayed constant 

until after explosion.
122 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

starting at 17:26:40. These larger signals were highly correlative 

between the two SEP acoustic sensors and did not correspond to 

seismic events on the SEP seismometer (fig. 7). Thus over the 

space of ~80 s the first phase of the explosion transitioned from 

producing mostly seismic signals in the first 60 s to producing both acoustic and seismic signals, and finally, 20 s later, to 

producing mostly acoustic signals.

Such variability in the production of acoustic and 

seismic energy during explosions has been documented at 

other volcanoes (Mori and others, 1989; Garcés and others, 

1998; Johnson and others, 2003; Johnson and others, 2005; 

Johnson and Aster, 2005). Johnston and Aster (2005) demonstrated that this variability can be attributed to a variety of 

factors, including plume density, with dense plumes resulting in reduced acceleration of the atmosphere and reduced 

infrasonic amplitudes; magma/wall rock impedance contrast, 

with low impedances resulting in reduced seismic amplitudes; 

viscous flow in the conduit, with long and narrow conduits 

acting to reduce seismic and enhance infrasonic amplitudes; 

and source dimension, with large source regions yielding 

reduced infrasonic amplitudes. For the March 8 explosion, we 

favor an ash-rich plume as the primary reason for the apparent absence of acoustic signals in the first 60 s. Given that no 

open vent existed before the explosion, the initial explosion 

must have contained substantial amounts of ash and fractured 

rock mixed with steam and/or magmatic gas during the initial 

vent clearing. Evidence for this comes from a picture taken at 

17:27:42 from a remote camera located 2.5 km northeast of 

the vent at SUG (fig. 10, lower photo; see fig. 2 for site location). That photograph shows a dark ash-rich cloud to the left 

(south) of the vent, a convecting lighter-colored plume to the 

right (north) that is relatively close to the SEP site, and impact 

craters (white patches) in snowfields on the northern side of 

the 1980–86 lava dome. The extent of the dark ash cloud, 

which enveloped most of the southern half of the crater just 

140 s after the initial explosion signal (fig. 10), suggests that 

the initial plume contained a substantial amount of ash. The 

appearance of acoustic signals after 60 s could correspond to 

reduced ash content in the plume (and more efficient acoustic 

energy production) following vent clearing, with the subsequent increase in acoustic-to-seismic energy ratios over the 

next 20 s corresponding to progressively reduced amounts of 

ash entrained in the plume.

At 17:27:00 a high-frequency (>5 Hz) tremor signal that 

rapidly increased in amplitude appeared on the SEP seismometer (fig. 6), clipping the seismometer after 20 s. This increase 

was also registered at other stations within 15 km, although 

relative amplitude increases were much smaller, indicating 

that the source was close to the SEP seismometer and was 

likely shallow. There was no associated change in the nature or 

amplitude of the infrasonic signals recorded on the SEP microphones. We infer that this signal represents the generation of a 

second and perhaps more vigorous ash-rich plume. Although 

the explosion chronology is poorly detailed, photographs taken 

from Camas, Washington, (75 km south of Mount St. Helens), 

show a second plume starting to rise above the crater rim at 

1727; by 1729 it had risen to the first plume’s altitude at 1727 

(fig. 11). The appearance of this second plume matches well 

with the timing of the onset of the more energetic tremor.

Further evidence for a second, more vigorous explosive 

phase comes from ballistic impacts. Most evidence comes 

from events that occurred after the tremor increase at 1727. 

At 17:27:22 the first of several sharp high-frequency signals 

appeared on the SEP microphones (fig. 7). Most of these signals appeared only on a single microphone, indicating that the 

source was small and very close to the sensors. We infer these 

to be caused by small ballistic fragments landing near individual microphones. This inference is supported by the appearance of impact craters in snow fields on the 1980–86 dome in 

a photo from the Sugar Bowl camera taken at 17:27:40 (fig. 

10), indicating that ballistic fragments were falling at the time 

of the high-frequency spikes. Furthermore, the downward firstmotion of the spikes (fig. 7) is inconsistent with an explosive 

source. We speculate that the downward first-motions might 

have been related to a pressure drop associated with nearby 

ballistic impacts. At 17:27:46, 12 s before SEP was destroyed, 

a low-amplitude, high-frequency signal appeared superimposed on the lower frequency infrasonic signals on both SEP 

microphones. On the basis of the high frequency content and 

small amplitude, we infer that this signal was caused by a rain 

of many small ballistic fragments across the SEP site. Finally, 

SEP and NED were both destroyed by hot ballistic fragments 

at 17:27:58 and 17:28:18, respectively, as indicated by multiple sharp punctures in enclosures, melted nylon ropes, and 

melted circuit-board solder found in equipment that was later 

retrieved from each site. The evidence for the bulk of ballistic 

impacts occurring after the tremor increase at 1727, coupled 

with the photographic evidence for a second plume starting 

at 1727, leads us to conclude that the 1727 tremor increase 

heralded the onset of the second and most energetic phase of 

the March 8 explosion.

Discussion

The explosions of 2004–5 at Mount St. Helens, although 

few in number and small in size, yielded a number of important insights regarding the use of acoustic and seismic instruments for reliable explosion detection, in particular for small 

explosions. One such insight comes from the observation that 

the most reliable seismic indicator for each of the four explosions occurring during the vent-clearing phase of October 

1–5, 2004, was the seismicity decline following each explosion (table 1). None of the explosions had an obvious seismic 

precursor, and no seismic signals were recorded that heralded 

the start of any of the October 1–5 explosions. The lack of 

such obvious seismic signals may in part have been due to the 

intense seismicity associated with the vent-clearing phase, 

which would have obscured any moderate-amplitude seismic signals associated with the explosions. Despite the fact 

that seismicity declines followed each explosion, however, 
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 123

Figure 10. Photographs of the crater of Mount St. Helens taken by a remote camera installed 

by the Cascades Volcano Observatory at SUG (2 km northeast of vent; see fig. 2) 10 s before 

(top) and ~140 s after (bottom) the start of the March 8 explosion at 17:25:20 PST. Times for each 

photo are accurate to nearest second, because time stamps were based on a clock periodically 

synchronized to Internet time server. View is to south; the south crater rim is in background and 

the northeast crater wall in foreground. The new lava dome is mostly shrouded by steam in upper 

image and completely obscured by ash clouds in lower image. The 1980–86 lava dome is in the 

center of each image (station NED is located on northeast side of that dome). Small white circles 

in the snow, visible in center of lower image, formed as impact craters from ballistic fragments 

that fell on all sides of the 1980–86 lava dome during explosion. (Pink diagonal lines in upper right 

corner of lower image result from damage to the camera in previous months when sun at low 

zenith shone directly into lens.)

 March 8, 2005

17:27:42

 March 8, 2005

17:25:10

Impact craters

NED
124 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Plume 1

Plume 2

1729 PST

Plume 2

Plume 1

1727 PST

Figure 11. Views of March 8, 2005, ash cloud at 1727 and 1729 

PST as seen from Camas, Washington, 75 km south of Mount 

St. Helens. Time for each photo is precise to nearest minute 

(resolution of internal clock in camera used to take these photos). 

Two plumes are seen clearly in lower photograph, with second 

plume likely corresponding to intensification in tremor seen on 

SEP and other seismic stations starting at 17:27:00 (see fig. 7). 

Given that second plume can also be seen in the 1727 photo, we 

estimate that the photo was taken several tens of seconds after 

17:27:00. Photographs by Elisa Wells, used with permission.

seismicity declines also occurred during this phase that did not 

correspond to explosions. Clearly, real-time explosion detection cannot be based on seismic instruments alone.

Another insight comes from the fact that there were no 

obvious acoustic signals associated with the October 3–5 

explosions that occurred after two single-sensor acoustic sites 

were established. The lack of acoustic signals in part indicates 

that the explosions were not large. However, because each 

site had only a single acoustic sensor, it is impossible to use 

methods such as those employed by Matoza and others (2007) 

on arrays of collocated acoustic sensors to significantly reduce 

noise levels. As illustrated by the subsequent explosion signals 

recorded at SEP in 2005, multiple acoustic sensors make it 

much easier to distinguish between explosion-related infrasonic signals and noninfrasonic noise sources such as wind 

and coseismic shaking of the microphone from passing seismic 

waves, especially if the explosion-related signals are weak.

The intense seismicity during the first two weeks at 

Mount St. Helens highlighted an additional underappreciated 

attribute of acoustic sensors; because they are isolated from 

the ground, acoustic sensors are not nearly as susceptible as 

seismometers to saturation by energetic seismicity and thus 

may record explosion-related infrasound when seismic waves 

from frequent large earthquakes mask any explosion-related 

seismicity. Although acoustic sensors deployed 3–5 km from 

the vent did not record obvious infrasound associated with the 

October 3–5 explosions, at the same time the acoustic sensors 

were not overwhelmed by ground waves from the continuous 

large earthquakes and thus were more capable of detecting 

explosion-related signals than the seismic network. Because 

vigorous precursory seismicity can happen at any volcano, 

acoustic sensors with real-time telemetry, in particular arrays 

of sensors such as those deployed at SEP and Coldwater 

(Matoza and others, 2007), should be installed as soon as possible after the onset of volcanic unrest to improve explosiondetection capabilities. The destruction of SEP by the March 8 

explosion (and the resultant loss of acoustic information about 

the evolution of the explosion) also demonstrates the value 

in placing some acoustic arrays at safe distances, despite the 

greater complexity of path effects, the higher signal attenuation, and the increased time delay in explosion detection inherent in recording explosion-related signals at more distant sites.

The January 16 and March 8, 2005, explosions also 

demonstrated the importance of having seismic stations within 

1–2 km of a vent for reliable explosion detection. Explosionrelated tremor from the January 16 explosion first appeared 

only on a station located ~250 m from the vent and only 

became apparent (albeit marginally) on stations outside the 

crater ~8 minutes later, when infrasonic signals also became 

apparent on acoustic sensors. Without seismic stations operating within 1–2 km of the vent, the explosion might have gone 

unnoticed until later field work by CVO staff. The March 

8 explosion was also recorded first by the seismic network, 

with infrasound signals appearing ~60 s after the start of the 

explosion. Given the aviation sector’s stated need to be alerted 

within 5 minutes of the start of an eruption (Ewert and others, 2005), such delays are potentially problematic. Seismic 

signals were large enough to be seen on stations within 15 

km of the vent, but the explosion-related tremor was most 

clearly recorded on crater stations. An additional benefit of 

having stations in the crater came from the loss of telemetry 

experienced at one or more crater stations during the October 

1, October 5, January 16, and March 8 explosions. These dropouts provided independent confirmation that ash was in the air, 

and for this reason the loss of telemetry from individual sta-
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 125

tions has since become a key element in the automated alarm 

system employed by CVO and PNSN (Qamar and others, this 

volume, chap. 3). Thus, installation of real-time seismic stations within 1–2 km of a volcanic vent is vital for improving 

explosion-detection capabilities at erupting volcanoes.

Acknowledgments

We are indebted to Marvin Couchman for his contributions in assembling, testing, and deploying the acoustic 

sensors at BLIS and SEP. We also wish to acknowledge Jeff 

Johnson for his advice on deploying multiple acoustic sensors 

at SEP; Robin Matoza and Michael Hedlin for generously 

sharing data and interpretations of acoustic signals recorded 

at their Coldwater infrasonic site; Mike Poland for having 

the foresight to install the Sugar Bowl camera in October 

2004; Rick LaHusen, for maintaining the camera and associated telemetry system in the months after Mike’s move to the 

Hawaiian Volcano Observatory; and Elisa Wells (wife of S. 

Moran) for having the presence of mind to start taking photos 

of the March 8, 2005, explosion cloud and for allowing us 

to publish them. Finally, we gratefully acknowledge Jackie 

Caplan-Auerbach and Jeff Johnson for their thoughtful and 

helpful reviews of this paper.

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eruptive cycle at Mount St. Helens?: Eos (American Geophysical Union Transactions), v. 87, no. 48, p. 499–500, 

doi:10.1029/2005EO480006.

Pallister, J.S., Hoblitt, R., Denlinger, R., Sherrod, D., Cashman, K., Thornber, C., and Moran, S., 2006, Structural geology of the Mount St. Helens fault-gouge-zone field relations 

along the volcanic conduit-wallrock interface [abs.]: Eos 

(American Geophysical Union Transactions), v. 87, no. 52, 

Fall Meeting supplement, Abstract V41A-1703.

Poland, M.P., Dzurisin, D., LaHusen, R.G., Major, J.J., Lapcewich, D., Endo, E.T., Gooding, D.J., Schilling, S.P., and 

Janda, C.G., 2008, Remote camera observations of lava 

dome growth at Mount St. Helens, Washington, October 

2004 to February 2006, chap. 11 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Qamar, A.I., Malone, S.D., Moran, S.C., Steele, W.P., and Thelen, 

W.A., 2008, Near-real-time information products for Mount 

St. Helens—tracking the ongoing eruption, chap. 3 of Sherrod, 

D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Rowe, M.C., Thornber, C.R., and Kent, A.J.R., 2008, Identification and evolution of the juvenile component in 2004–

2005 Mount St. Helens ash, chap. 29 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Schilling, S.P., Thompson, R.A., Messerich, J.A., and Iwatsubo, E.Y., 2008, Use of digital aerophotogrammetry to 

determine rates of lava dome growth, Mount St. Helens, 

Washington, 2004–2005, chap. 8 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Schneider, D.J., Vallance, J.W., Wessels, R.L., Logan, M., 

and Ramsey, M.S., 2008, Use of thermal infrared imaging 

for monitoring renewed dome growth at Mount St. Helens, 

2004, chap. 17 of Sherrod, D.R., Scott, W.E., and Stauffer, 

P.H., eds., A volcano rekindled; the renewed eruption of 

Mount St. Helens, 2004–2006: U.S. Geological Survey 

Professional Paper 1750 (this volume).

Scott, W.E., Sherrod, D.R., and Gardner, C.A., 2008, Overview of the 2004 to 2006, and continuing, eruption of 

Mount St. Helens, Washington, chap. 1 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Vallance, J.W., Schneider, D.J., and Schilling, S.P., 2008, 

Growth of the 2004–2006 lava-dome complex at Mount St. 

Helens, Washington, chap. 9 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Woulff, G., and McGetchin, T.R., 1976, Acoustic noise from 

volcanoes; theory and experiment: Geophysical Journal of 

the Royal Astronomic Society, v. 45, p. 601–616.
6. Seismicity and Infrasound Associated with Explosions at Mount St. Helens, 2004–2005 127

Appendix 1. Acoustic Sensors Used at 

Mount St. Helens

The acoustic sensors installed by CVO and PNSN at 

Mount St. Helens were previously designed as additions to the 

McVCO voltage-controlled oscillator (McChesney, 1999) that 

is in wide use by the PNSN and the USGS in the Cascades 

and elsewhere. The sensor consists of either 9 or 18 electret 

microphones with a 1-Hz frequency response, the signals from 

all microphones being summed to reduce electronic noise. 

Including the responses of other components in the recording system (for example, radio, digitizers, system gain), each 

9-element microphone has a nominal response of 22.7 counts/

Pascal (44.5 counts/Pascal for an 18-element microphone). 

The acoustic sensors installed at stations SEP and SUG had 

18-element microphones, and the STD sensor had a 9-element 

microphone. An important caveat to the acoustic response of 

these sensors is that individual sensors were not calibrated 

before installation. Given that the sensitivity of individual 

electret sensors varies by ±3 dB and that their sensitivity is 

known to decline with time after installation (depending on 

environmental conditions), we estimate that the true response 

could vary by 30 percent or more for each nine-element 

microphone. For this reason we only give a range of possible 

pressure values in this paper.

The circuit board with electret sensors is typically placed 

in a PVC plastic tube with end caps and a hose connection on 

one end. The PVC tubing reduces bellows-type motion caused 

by wind or ground shaking. A soaker hose of variable length is 

attached to the hose connection to further reduce wind noise. 

At SUG and STD the sensors were placed inside an enclosure 

with the soaker hose strung in a line away from the enclosure. 

At SEP the two sensors were placed on the ground ~15 m 

apart, with rocks piled around the PVC tube to prevent movement and the soaker hose coiled next to each sensor.
128 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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 129

Chapter 7

Seismic-Monitoring Changes and the Remote 

Deployment of Seismic Stations (Seismic Spider) at 

Mount St. Helens, 2004–2005

By Patrick J. McChesney1, Marvin R. Couchman2, Seth C. Moran2, Andrew B. Lockhart2, Kelly J. Swinford2, 

and Richard G. LaHusen2

Abstract

The instruments in place at the start of volcanic unrest at 

Mount St. Helens in 2004 were inadequate to record the large 

earthquakes and monitor the explosions that occurred as the 

eruption developed. To remedy this, new instruments were 

deployed and the short-period seismic network was modified. 

A new method of establishing near-field seismic monitoring 

was developed, using remote deployment by helicopter. The 

remotely deployed seismic sensor was a piezoelectric accelerometer mounted on a surface-coupled platform. Remote 

deployment enabled placement of stations within 250 m of 

the active vent.

Introduction

The earthquake swarm that signaled the start of the eruption at Mount St. Helens on September 23, 2004 (Scott and 

others, this volume, chap. 1), was recorded by a dense network 

of short-period stations operated jointly by the Pacific Northwest Seismic Network (PNSN), based at the University of 

Washington, and the U.S. Geological Survey (USGS) Cascades 

Volcano Observatory (CVO). The network consisted of 13 stations within 20 km of the volcano; 6 of these were within 5 km 

(fig. 1). Many stations had been in place since the early 1980s. 

Although the established monitoring was sufficient to detect the 

onset of the unrest, it proved inadequate to record and monitor 

the intense seismicity of the developing 2004–2005 eruption.

This paper recounts the changes made to the Mount St. 

Helens seismic network during the first year of the eruption. 

These include changes to the existing short-period network, 

the first installations of telemetered broadband seismometers, 

the addition of infrasonic microphones, and the remote deployment of piezoelectric accelerometers. The method of remote 

deployment is new—it relies on an innovative instrument 

package, called a “spider,” that was developed during the first 

month of the eruption (LaHusen and others, this volume, chap. 

16). The package enabled us to safely deploy fully functional 

seismic stations, by helicopter, to sites within a few hundred 

meters of the active vent without setting a foot on the crater 

floor (fig. 2). Seismic spider development occurred during the 

response to volcanic unrest, and the following chronological 

sections complement the technical discussion by showing how 

the seismic spiders were used and how our field experiences 

drove the design process. Because the seismic spider uses 

both an uncommon sensor and unconventional deployment, 

we describe it at length in a subsequent section, “Technical 

Description of the Seismic Spiders.”

Deployment Chronology

The deployment chronology is divided into two sections. 

During the initial response we modified the existing shortperiod network and made new installations using traditional 

methods. Later, our focus was seismic spider deployment. 

To provide an overview, a summary of changes to the seismic monitoring system through 2005 is given in table 1. 

The instrument parameters for the spider deployments are 

described in a later section. The infrasonic microphones are 

described elsewhere (Moran and others, this volume, chap. 6).

1

 Pacific Northwest Seismic Network, Department of Earth and Space Sciences, University of Washington, Box 351310, Seattle, WA 98195

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008
130 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 1. Station locations 

in the area of Mount St. 

Helens, Washington, during 

the period from September 

2004 through December 

2005. Green triangles, spider 

accelerometers; black 

triangles, short-period velocity 

seismometers; red triangles, 

broadband seismometers. 

Some sites had more than 

one sensor type; see text and 

table 1 for details. Inset-map 

base from April 19, 2005, digital 

elevation model.

Initial Response to Seismic Unrest

Because Mount St. Helens was regarded as adequately 

monitored (Moran, 2004; Ewert and others, 2005), there was 

little urgency to install new instruments during the first several 

days of the seismic swarm that began on September 23, 2004. 

However, after seismicity intensified on September 26, the existing network proved inadequate in four ways (Moran and others, 

this volume, chap. 2): (1) The short-period network lacked 

dynamic range and repeatedly clipped on larger events. (2) There 

were no three-component instruments. (3) The network lacked 

the ability to record low-frequency signals. (4) There were no 

infrasonic sensors to record explosive activity. Consequently, 

CVO and the PNSN began mobilizing to put new stations in the 

field. Our priorities were the installation of broadband threecomponent seismometers and deployment of infrasonic microphones to complement the dense short-period network.

At the same time, other cooperating groups were mobilizing to deploy broadband and strong-motion sensors with 

onsite recording around the volcano (Horton and others, this 

volume, chap. 5). However, to overcome the limited dynamic 

range and frequency response of the short-period network, we 

felt it essential to have real-time data from several broadband 

stations transmitted to CVO and PNSN. Because of the added 

complication of providing telemetry, it took several days to 

organize the deployment.

JRO

SHW

HSR

JUN

STD

BOLM/SUG

Coldwater

maintenance

center

46°15'

46°00'

122°22'30" 122°15' 122°7'30" 122°00'

46°7'30"

46°22'30"

0 5 10 km

WASHINGTON

AREA OF MAP

SEND

RAFT

BLIS

NED

MIBL

MIDE

AHAB

SEP

YEL

WESG

500 m
7. Seismic-Monitoring Changes and the Remote Deployment of Seismic Stations (Seismic Spider) 131

Figure 2. Overhead view of seismic spider deployed at station 

RAFT. Sensor platform (Marv lander) is at upper left. Instrument box 

on spider is 70 cm long. USGS photo by S.C. Moran, June 20, 2006.

On October 1, 2004, we made our first attempt to install 

a broadband seismometer and an infrasonic microphone at 

Studebaker Ridge, the site of an existing short-period station 

(STD; fig. 1). Work at the site, about 5 km from the vent, 

stopped when the first explosion of the eruption (at 1202 

PDT) triggered a mandatory evacuation. As an alternative, 

we installed a broadband seismometer at the Johnston Ridge 

Observatory (JRO; fig. 1), 9 km north-northwest of the crater. 

Later, a broadband seismometer was successfully installed 

at Studebaker Ridge on October 5. Both seismometers were 

Guralp CMG-6TDs. These three-component instruments have 

a response of 30 s and were configured to measure each channel at 50 samples per second. Telemetry was not yet complete, 

so data were recorded initially in each instrument’s 2-GB flash 

memory and collected later with a portable hard drive.

The interrupted installation of an infrasonic microphone 

at Studebaker Ridge (STDM) was completed on October 

2, 2004, and another microphone (BOLM) was deployed at 

Sugar Bowl (fig. 1) on October 3 (Moran and others, this 

volume, chap. 6). The short-period seismic network was 

modified to provide telemetry for the microphone signals. 

Other changes to short-period stations were necessary in the 

first weeks of the eruption because signals were clipped for 

many earthquakes. To improve the dynamic range, a low-gain 

channel was added at several stations. Stations SHW and HSR 

were converted to dual gain on October 4 and JUN on October 

7 (fig. 1). The sensitivity on the high-gain channel from HSR 

was also reduced, because signals from even relatively small 

earthquakes were severely clipped.

Ideally the broadband instruments and microphones 

would have been located closer to the vent. Experience elsewhere has shown that very long period (VLP) earthquakes are 

best recorded on broadband instruments located within 4 km 

of the source region (B.A. Chouet, oral commun., 2004). By 

October 1, however, conditions inside the crater and on the 

crater rim were considered too dangerous for extended field 

work, so the instruments were installed more than 4 km from 

the source region. We will never know if VLP events occurred 

during the first several months of the eruption, a shortcoming 

that shows the importance of having at least one broadband 

instrument installed within 1–3 km of a potentially active vent 

before the beginning of unrest. None of the three explosions 

that occurred in the days after the October 1 event created an 

obvious signal on the microphones placed outside the crater 

(Moran and others, this volume, chap 6).

On October 20, an Earthworm data-processing node was 

established at the Gifford Pinchot National Forest’s Coldwater Ridge maintenance facility, ~13 km northwest of the 

crater (fig. 1). This achievement marked a turning point in 

our response to the eruption by allowing Coldwater to function as a radio-telemetry terminus near the volcano. Links for 

the data from JRO and STD broadband seismometers were 

made to Coldwater with Freewave spread-spectrum digital 

radios (2.4 GHz). The presence of an Earthworm digitizer 

vastly expanded our analog channel capacity and allowed 

further seismic deployments to occur without long and difficult radio paths to CVO. Internet connectivity at Coldwater 

was through VSAT (very small aperture terminal) linked to 

a commercial communications network by microwave radio 

and satellite.

The explosion on October 1, 2004, destroyed the shortperiod seismic station at September lobe on the 1980–86 lava 

dome (SEP; fig. 1). Positioned ~500 m from the vent, SEP 

was the closest station to the earthquake sources. Losing this 

station greatly hampered earthquake location quality (Moran 

and others, this volume, chap. 2). The only remaining station within 3 km of the vent was the short-period station YEL 

(fig. 1) on the crater floor. However, the signal from YEL was 

clipped on many events until the gain was reduced by 12 dB 

during a quick visit on October 21.

Chronology of Seismic Spider Deployment

The need to reestablish near-field seismometers in the 

still-hazardous working conditions in the crater turned our 

attention to the development of a seismic station that could 

be deployed remotely. Available to us was the spider instrument package, originally developed for GPS deformation 

instruments and telemetry (LaHusen and others, this volume, 

chap. 16). We adapted this platform to carry an accelerometer 

(described in the next section) and an infrasonic microphone. 

The first seismic spider (BLIS, fig. 1) was deployed on October 

12, 2004, at a location ~250 m east of the vent. The spider was 
132 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Station Date Change

SEP 10/1/04 Destroyed by explosion (L4-C seismometer)

JRO 10/1/04 Guralp CMG-6TD broadband installed

STDM 10/2/04 Nine-element infrasonic microphone installed

BOLM 10/3/04 Eighteen-element infrasonic microphone installed

SHW 10/4/04 Low gain added. High gain = 60 dB, low gain = 36.5 dB (L4-C seis.)

HSR 10/4/04 Dual gain. High gain = 28 dB, low gain = -1.32 dB (S-13 seis.)

STD 10/5/04 Guralp CMG-6TD broadband installed

JUN 10/7/04 Low gain added. High gain = 60 dB, Low gain = 36.5 dB (L4-C seis.)

STDM 10/12/04 Disconnected microphone telemetry for use at station BLIS

BLIS 10/12/04 Spider with accelerometer and 18-element infrasonic microphone

YEL 10/21/04 Gain reduced by 12 dB to 54 dB (L4-C seismometer)

SEP 11/4/04 Installed L22-3D and a pair of 18-element infrasonic microphones

NED 11/20/04 Spider with accelerometer

NED 12/20/04 Spider with accelerometer on separate platform (first Marv lander)

AHAB 2/8/05 Temporary spider deployment, accelerometer, and GPS

MIDE 2/16/05 Spider with accelerometer

SEP 3/8/05 Destroyed by explosion

NED 3/8/05 Destroyed by explosion

MIDE 3/8/05 Destroyed by explosion

SUG 3/9/05 L4-C seismometer at 42 dB added to station BOLM

SEP 3/14/05 Spider with accelerometer and two 18-element infrasonic microphones

YEL 3/15/05 Gain reduced by 6 dB to 48 dB

NED 4/6/05 Reinstalled spider with accelerometer

MIDE 4/6/05 Reinstalled spider with accelerometer, reinstalled again 04/14/05

SEND 6/30/05 Spider with accelerometer

WESG 7/12/05 Temporary spider deployment, accelerometer; removed 09/14/05

MIDE 7/19/05 Destroyed by rockfall

RAFT 7/28/05 Spider with accelerometer

MIBL 11/17/05 Spider with accelerometer

Table 1. Summary of changes to preeruption seismic network at Mount St. Helens, Washington, through 

2005, including infrasonic microphones.

remarkably successful, despite noise spikes in the seismic signal 

caused by radio transmissions from the co-housed GPS system. 

These noise spikes were eliminated by remotely turning off the 

GPS instrument, which had been damaged during deployment.

Volcanic and seismic activity declined substantially after 

October 5, 2004 (Moran and others, this volume, chap. 2). 

After several weeks of reduced activity, we decided it was safe 

to work at SEP for a short time to replace the destroyed station 

with a three-component seismometer and a pair of infrasonic 

microphones. Usually such installations take hours, but by 

using the spider package we thought the visit would require 

less than one hour. On November 4, a spider was deployed in 

advance of the crew. Everything required for the station was 

preassembled and housed in the spider except the seismometer 

and microphones. It took only 45 minutes to install the Sercel 

L22-3D three-component seismometer and two infrasonic 

microphones. Even so, the helicopter was forced to depart hurriedly when an ash cloud advanced on the site, just as the crew 

was loading the ship to leave.

The success of the spider at BLIS led to another spider 

deployment at the northeast corner of the old dome (NED; fig. 

1) on November 20, 2004. This site was selected because it 

improved the geometry of stations in the crater and included 

a patch of warm ground that was free of snow. Data became 

available on November 22, after an expansion of the analog 

telemetry channels at the Coldwater Earthworm node.

The first NED spider worked poorly. Unlike BLIS, the 

seismic signal from NED was often obscured by noise. At 

both stations the accelerometers were mounted in a leg of the 

spider platform. However, NED used a heavy radome antenna 

mounted on a mast, whereas BLIS had used a low-mass whip 

antenna. Wind or ground motion shook NED’s antenna and 

sent vibrations through the spider to the sensor. We considered using whip antennas in the future, instead of radomes, 

but the heavier antenna had two advantages—increased signal 

strength and protection from icing. In addition, we could not 

be certain that the antenna was the only noise source. Consequently we decided to isolate the sensor from spider vibrations 
7. Seismic-Monitoring Changes and the Remote Deployment of Seismic Stations (Seismic Spider) 133

by mounting it in a separate platform, which we dubbed the 

“Marv lander” (see technical description in the next section). 

On December 23, 2004, a new spider equipped with a Marv 

lander was exchanged for the old one, solving the noise problem at NED.

On February 8, 2005, the spider AHAB was deployed 

atop the actively moving spine 4 (fig. 1). The location on the 

spine was chosen because it appeared stable and its slope of 

~20o was not too steep for the spider. The objective of this 

short-term experiment was to record spine motion using GPS 

and to look for a correlation between “drumbeat” seismic 

events (Moran and others, this volume, chap. 2) and discrete 

movements of the spine. The spider housed an L-1 GPS 

receiver with an accelerometer and electronic thermometer 

mounted on a Marv lander. After 8 days of successful operation, AHAB was retrieved before the spine’s process of growth 

and collapse could destroy it.

By the end of January 2005, following a period of 

intermittent operation, the BLIS spider stopped operating. We 

were unable to find the spider because it was buried in snow. 

Consequently, when AHAB was removed from spine 4, it 

was modified in the field to become a replacement for BLIS. 

The new location, MIDE, was a snowfree warm spot near the 

BLIS site (fig. 1). Station MIDE began operation on February 

16, 2005.

A small but destructive explosion occurred on March 8, 

2005 (Scott and others, this volume, chap 1; Moran and others, 

this volume, chap. 6). Ballistic fragments from the explosion 

destroyed MIDE, SEP, and NED. The MIDE equipment was 

never found. Scattered parts of SEP were barely visible under 

a blanket of ash. Although NED was recovered, little could 

be salvaged—even the apparently undamaged accelerometer 

no longer functioned. The loss of these three stations severely 

reduced our ability to record and locate earthquakes, and the 

loss of the close-in microphones at SEP eliminated our ability 

to detect small explosions.

The first step in restoring lost monitoring capacity was 

the installation of a short-period seismometer just north of 

the crater at SUG (fig. 1) on March 9. We chose this site 

because of its relative safety and the fact that the Sercel L4-C 

seismometer could be easily connected to the existing infrasonic microphone (BOLM) telemetry. A new seismic spider, 

including two 1-Hz microphones, was rapidly constructed and 

deployed at SEP on March 14. The crater station YEL was 

visited on March 15 and the gain lowered by another 6 dB 

because of continuing clipping problems.

The weather then took a turn for the worse, delaying 

replacement of spiders at MIDE and NED until April 6, 2005. 

The NED deployment went smoothly, but the MIDE spider 

toppled when released by the helicopter. We retrieved and 

inspected it and, finding no obvious problems, returned the 

spider to MIDE. However, it failed several days later and was 

retrieved again on April 10. Several of the power system’s primary cell casings had melted, likely a result of internal damage caused by the spider’s tumble on April 6. After replacing 

the cells, the spider was returned to MIDE on April 14.

With the arrival of summer, more sites were free of 

snow and available for seismic spider deployment. On June 

30, 2005, in an effort to surround the source of the drumbeat 

earthquakes, a spider was placed at SEND, southeast of the 

former AHAB installation at the southeast end of spine 4, 

which by then was disintegrating (fig. 1). Another installation 

followed at WESG (western arm of the Crater Glacier) on July 

12 (fig. 1). Unfortunately, the sensitivity of WESG was poor. 

We suspected bad ice-rock coupling or an electronic fault. 

Examination after the station was retrieved showed an intermittent problem with the accelerometer interface circuit. On 

July 19, MIDE was destroyed by a rock fall. On July 28, 2005, 

a replacement spider was deployed at RAFT, the only available patch of stable ground close to the vent (figs. 1, 2).

Finally, with the approach of winter we took several 

steps to improve the robustness of the crater spiders. WESG 

was removed on September 14, 2005, before it could be 

buried by snow. We also retrieved all spiders except RAFT to 

replace batteries and, if necessary, to add directional antennas. 

Subsequently, ice accumulation from an early winter storm 

damaged the coaxial cable for the radio transmitter at SEND. 

The SEND spider was retrieved and repaired on November 17. 

Because earthquakes had become much smaller, we relocated 

this spider to a new site, MIBL, closer to the vent (fig. 1). The 

MIBL site was warm and free of snow and near the previous 

MIDE and BLIS sites.

By the end of 2005, the CVO and PNSN real-time seismic network at Mount St. Helens consisted of 2 broadband 

seismometers, 13 conventional short-period instruments (3 

with dual-gain channels), 4 seismic spiders, and 2 infrasonic 

microphones, all within 20 km of the volcano (table 1).

Technical Description of the Seismic 

Spiders

The loss of SEP during the explosion of October 1, 2004, 

forced us to improvise new techniques in order to restore 

seismic monitoring near the vent. Particularly for the first 

installations, we hurried to get something in the field, and each 

deployment was an experiment that led to changes in design. 

The evolving design complicates the description of the spider, 

because it is necessary to include the changes that were made. 

We trust that this section gives sufficient detail to satisfy 

both those who need to know the response parameters for a 

spider installation and those who are considering the design of 

remotely deployed seismic instruments.

Seismic Spider Overview

We designed the seismic spider for remote deployment 

close to an active volcanic vent where seismic signals were 

strong. Remote deployment ruled out the use of traditional 

seismic sensors, which require leveling by hand during instal-
134 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

lation. We knew of a sensor, used for measuring seismic activity of lava flowing through tubes in Hawai‘i (R. Hoblitt, oral 

commun., 2004), that did not require leveling, could measure 

strong signals, and could survive the rigors of remote deployment. This sensor, a piezoelectric accelerometer, made the 

seismic spider possible.

Analog telemetry was used for a variety of reasons. 

Given strong signals, the contribution of some noise by 

the analog telemetry system did not detract much from the 

overall signal-to-noise ratio. The limited dynamic range of 

the telemetry system was overcome by using high- and lowgain channels. Analog telemetry allowed the use of lower 

frequency analog radios that, unlike digital spread-spectrum 

radios, do not require line-of-sight transmission paths and 

use much less power. The radios operated on the 406 to 420 

MHz band, where antennas are fairly compact and their 

100-mW transmissions can cover tens of kilometers. Several 

antennas were tried, but high-gain directional antennas were 

preferred because the radio path was often obstructed by terFigure 3. Seismic spider slung by helicopter from staging area 

on March 14, 2005, for deployment at SEP. Included are two 

infrasonic microphones, one suspended immediately beneath the 

spider and another below it on the lander sensor platform (Marv 

lander), last item in the string. PNSN photo by P.J. McChesney, 

March 14, 2005.

rain or snow. To prevent breakage from snow and ice loading, 

radome-protected antennas were favored.

Power was supplied by Air-Alkaline primary cells (Celair 

Corp.). Nine series-connected cells provided 12.6 V at 1,200 

Ah. Total current drain was 110 mA, so 450 days of continuous operation were possible, but we planned for no more than 

a year because of capacity loss in cold weather.

The spider instrument package, developed early in this 

eruption to remotely deploy GPS instruments (LaHusen 

and others, this volume, chap. 16), provides an instrument 

compartment (~70×40×40 cm) optimized for helicopter 

transport and placement (fig. 3). We used the spider to house 

the power and telemetry systems and, initially, the sensor. 

Because of noise problems with the first two deployments, 

the sensor was moved to a separate platform, the Marv lander 

(described below).

Piezoelectric Accelerometer

Long used by industry for machine vibration measurement, piezoelectric accelerometers do not require leveling 

because they operate in any orientation. These devices convert 

dynamic forces to electrical energy through charge separation 

in a piezoelectric material. Acceleration (a) along the axis of 

the transducer acts on the seismic mass (M) to apply a force to 

the piezoelectric element. The force on the element produces 

a charge (q) such that q = dMa, where d is the piezoelectric 

constant. The stressed piezoelectric material acts as a capacitor (C), producing an open-circuit voltage (V) where V = q/C. 

Consequently the sensitivity of the accelerometer is V/g, where 

g is the unit of acceleration (Allocca and Stuart, 1984, p. 122).

Piezoelectric accelerometers respond to dynamic 

forces—to vibrations. The circuit model is a current generator in parallel with a capacitor, or its equivalent circuit, 

a voltage generator in series with a capacitor (Allocca and 

Stuart, 1984, p. 122–123). The model shows that there is an 

inherent low-frequency limit to a piezoelectric accelerometer. 

In addition, low-frequency vibrations produce weak acceleration signals, so piezoelectric accelerometers operating at low 

frequencies must use charge amplifiers to boost signal levels. 

The low-noise design of this amplifier is critical for lowfrequency measurements because noise from the high input 

resistance of the charge amplifier increases with decreasing 

frequency (Schloss, 1993, p. 2). Consequently, poor signalto-noise ratios at low frequency set the practical limit of 

piezoelectric response.

Sensor Parameters

We chose the Wilcoxon Model 731-207 Ultra Low Frequency Seismic Accelerometer for this application because of 

its relatively good low-frequency performance (see abbreviated specifications in table 2). There are other piezoelectric 

accelerometers with better low-frequency specifications than 

this model, but there is a tradeoff between increased low-
7. Seismic-Monitoring Changes and the Remote Deployment of Seismic Stations (Seismic Spider) 135

frequency response and increased fragility. Because the sensors are subject to significant forces during deployment, the 

250 g shock limit of the Model 731-207 was attractive. All 

units survived our handling and the helicopter deployment. 

However, one unit from NED, recovered after the explosion 

of March 8, 2005, was irreparably damaged, probably from 

ballistic impact.

The peak acceleration range of this model is more than 

sufficient for close-in monitoring. The low-gain telemetry 

channel was set for a maximum measurement of 0.1 g, a limit 

that was never exceeded despite recording ML 3.4 earthquakes 

within 250 m of the source. The high sensitivity of the device 

permitted conservative seismic amplifier gain settings. Two 

telemetry channels were used for high- and low-gain recordings of the signal, with the low-gain channel set at 2.5 V/V. 

The high-gain channel was initially set at 83.17 V/V but was 

reduced to 41.75 V/V for later deployments. In the absence of 

noise, these gain settings achieve 16-bit dynamic range using 

two analog telemetry channels digitized at 12 bits each.

Response information for individual sensors is unavailable from the manufacturer. The nominal frequency response 

at 0.2 Hz is a worst-case limit. The design value for the lowfrequency cutoff (−3 dB) is 0.1 Hz. This high-pass response 

has a pole at (−0.628319, 0) and a zero at (0, 0) set by charge 

amplifier components. Actual response is determined by 

component variation, and any particular unit may have a pole 

frequency from 0.1 to 0.15 Hz (Ron Denton, Application 

Engineer, Wilcoxon Research, written commun., 2006).

We attempted continuous electronic integration of the 

accelerometer output to produce a velocity response. Even 

though integration in the field was not satisfactory, it demonstrated the noise characteristics of the sensor. The data 

sheet for the Model 731-207 states that spectral noise, a 1-Hz 

bandwidth noise-density measurement at a particular frequency (Schloss, 1998, p. 1), increases as frequency decreases. 

At 2 Hz the noise density is 0.28 g/Hz. However, integration provides gain relative to the acceleration response; gain 

increases by a factor of two each time frequency is halved. 

In addition, below 2 Hz, sensor noise increases as 1/f (Ron 

Denton, Application Engineer, Wilcoxon Research, written 

commun., 2005). The combination of decreasing signal-tonoise ratio and increasing gain causes low-frequency noise to 

be prominent in the integrated response. Analog filtering of 

the low-frequency noise interfered with the integrator phase 

response. On the other hand, the acceleration response of the 

sensor has very good noise performance. Consequently the 

acceleration response was telemetered because we felt that 

digital postprocessing could do a better job if the velocity 

response was desired.

Piezoelectric Accelerometer Interfacing

The charge amplifier in the accelerometer can produce 

high-amplitude signals (5 V) and requires a power source. The 

voltage-controlled oscillators (VCOs) used to telemeter seisAccelerometer Specifications

Sensitivity 10 V/g, ±10 %, 25oC

Acceleration Range ±0.5 g

Frequency Response 0.2–1300 Hz, −3 dB

Resonance Frequency 2,400 Hz

Temperature Response −18% at 0oC, +8% at 80oC

Broadband Noise 2 μg, 2.5 Hz to 25 kHz

Input Voltage Range 18–30 VDC, in series with a 2–10 

mA current source diode

Shock Limit 250 g

Table 2. Abbreviated specifications for the Wilcoxon Model 731-

207 piezoelectric accelerometer.

[Additional information online at www.wilcoxon.com.]

mic signals are usually connected to passive seismometers that 

produce small-amplitude signals in the millivolt range. This 

mismatch required an interface circuit and VCO gain adjustments. We used McVCO, a microcontroller-based mimic of an 

analog VCO (McChesney, 1999). The interfacing problems are 

similar for other VCOs, but the details are best appreciated if 

the McVCO documentation is at hand.

The Model 731-207 accelerometer is a two-wire device. 

There is a ground connection and a combined signal and 

power connection. Power is provided through a current-source 

diode in series with a DC supply of 18–30 V. The signal rides 

on a 10-V bias. The most straightforward way to connect the 

accelerometer to the VCO is to raise the impedance of one 

VCO input and make a single-ended connection through a 

coupling capacitor. The nominal 24-V power required by 

the accelerometer can be generated from the 12-V telemetry 

power system with a voltage-doubler circuit.

This approach was used for the first deployments at BLIS 

and NED. The coupling capacitor at the VCO input produced a 

high-pass pole at 0.047 Hz, well below the nominal 0.2-Hz charge 

amplifier pole in the accelerometer. To achieve this low frequency, 

tantalum electrolytic capacitors were used. However, investigation 

of noise problems with the first NED deployment indicated that 

some sudden baseline shifts were due to the capacitor. The second 

deployment at NED used a polypropylene capacitor and higher 

input resistance to produce a pole at 0.04 Hz.

The polypropylene capacitor appeared to reduce some 

noise problems, but the design was still bothered by the spikes 

that occurred when GPS and seismic instruments cohabited 

the same spider. This and an interest in electronic integration 

of the accelerometer signal provoked the development of a 

different interface (fig. 4). The new interface used an input 

buffer to raise the impedance and a differential output stage 

that allowed connection directly to an unmodified VCO input. 

Rather than use a 24-V supply, the sensor ground was operated 

at −12 V, putting the output offset at −2 V and eliminating the 

need for a coupling capacitor at the sensor output. An optional 

integration stage was provided.
136 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A high-pass pole (R4, C16) is included in the interface 

after the buffer stage. This was initially set to 0.017 Hz, making the charge amplifier pole dominant at low frequencies and 

reducing interaction with the optional integrator. This was 

later raised to 0.17 Hz when some low-frequency noise problems were seen in the accelerometer response. It has not been 

possible to duplicate this noise; it may have been unique, but 

for the sake of uniformity the pole continues to be set at 0.17 

Hz. Table 3 lists deployment interface response characteristics 

and gains. Poles and zeros are from circuit simulation. All 

responses more than ten times the 30-Hz VCO low-pass filter 

are ignored.

This interface has been used for all our piezoelectric 

accelerometer deployments starting with the deployment 

of AHAB on February 8, 2005, and continuing through 

2005, but it could be improved. Cohabitation with GPS still 

produces small glitches in the seismic signal when the GPS 

R4

2 M

+

-

+

-

R5

10K

P1 100K

R15

1M

C16, 4.7 uF or 0.47 uF**

BATT+

BATT8

4

7

6

5

1

2

3

1 8

U1

7662

+

+

+12 V

R12

100K

R11, 100K

R10, 4.99K

R14

4.99K

R13

4.99K

C2

10 uF,16 V

C4,10 uF,16 V

C18

 0.1 uF

C17, 0.1 uF

OUTOUT+

D1, 1N5819

TVS5

7.5 V

TVS4

7.5 V

TVS1

18 V

+

-

D3, J508, 2.4 mA

Accelerometer, Wilcoxon, 731-207

-12 V

-12 V

+12 V

TVS2

12V

Optional

Integrator

R8

 1 M

R7, 1 M

7

C8

 56 pf

4

+12 V

-12 V

C6,10 uF

+ 25 V

+

C5,10 uF

16 V

C3, 0.1 uF

C1,10 uF

16 V

+12 V

-12 V

R9

 0 ohms

C7, 0.1 uF

C10, 0.1 uF

R6, 1 K

R17, 1 K

U4a

LT1013

U4b

LT1013

+12 V

-12 V

C12, 15 nF

U2

TL031

12 to

15 VDC

**C16 and R4 set the high pass pole.

 C16 = 4.7 uF, f = 0.017 Hz

 C16 = 0.47 uF, f = 0.17 Hz

data transmission occurs. We expect that GPS noise could 

be reduced further by locating the interface on the sensor 

platform (fig. 3). This would change the connection in the 

long cable between the lander and the spider from a high 

impedance single-ended connection to a lower impedance 

differential connection. Better supply regulation might also 

help eliminate noise.

Gain Setting

The gain model for the dual-gain mode we used for all 

seismic spiders is shown in figure 5. Maximum acceleration 

of the Model 731-207 produces a maximum signal level of 

5 V. This is greater than the input range of the McVCO, so 

the signal is reduced by half in all versions of the interface. 

The gain of two, created by the differential connection of the 

Figure 4. Piezoelectric accelerometer interface used from February 8, 2005, onward. Low-frequency 

response controlled by C16 and R4.
7. Seismic-Monitoring Changes and the Remote Deployment of Seismic Stations (Seismic Spider) 137

Station Date High gain

V/V

Low gain

V/V

Interface response

(poles), (zeros)

BLIS

lost

10/12/04

01/31/05

83.17

--

2.5

--

(−0.294685, 0), (0, 0)

--

NED(1)

exchanged

11/20/04

12/20/04

83.17

--

2.5

--

(−0.294685, 0), (0, 0)

--

NED(2)

destroyed

12/20/04

03/08/05

83.17

--

2.5

--

(−0.249969, 0), (0, 0)

--

AHAB

removed

02/08/05

02/16/05

41.75

--

2.5

--

(−0.106357, 0), (0, 0)

--

MIDE(1)

destroyed

02/16/05

03/08/05

41.75

--

2.5

--

(−0.106357, 0), (0, 0)

--

SEP

modified

03/14/05

10/18/05 

41.75

nc

2.5

nc

(−0.106357, 0), (0, 0)

( 1.06357, 0), (0, 0)

NED(3) modified 04/06/05

10/18/05

41.75

nc

2.5

nc

(−0.106357, 0), (0, 0)

(−1.06357, 0), (0, 0)

MIDE(2)

destroyed

04/06/05

07/19/05

41.75

--

2.5

--

(−0.106357, 0), (0, 0)

--

SEND

removed

06/30/05

11/17/05

41.75

--

2.5

--

(−1.06357, 0), (0, 0)

--

WESG

removed

07/12/05

09/14/05

41.75

--

2.5

--

(−1.06357, 0), (0, 0)

--

RAFT 07/28/05 41.75 2.5 (−1.06357, 0), (0, 0)

MIBL 11/17/05 41.75 2.5 (−1.06357, 0), (0, 0)

Table 3. Piezoelectric accelerometer channel gains and interface frequency responses.

[Different deployments to the same location are indicated by numerals in parentheses in station column. Last date 

for each station entry indicates modification or end of operation. nc, no change.]

Accelerometer

10 V/g

Interface

0.5 V/V

Preamp

10 V/V

Variable Gain

16.63 or 8.35 V/V

Fixed Gain

0.5 V/V

Total = 83.17

or 41.75 V/V

McVCO Total = 2.5 V/V

Filter

1 V/V

Modulator

50 Hz/V

Modulator

50 Hz/V

High Gain Channel

Low Gain Channel

Figure 5. Dual-gain model of spider accelerometer. Additional gain and response terms result from demodulation and digitization. 

McVCO, microcontroller-based mimic of analog voltage-controlled oscillator.
138 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Channel gain, 

V/V

Scale factor, 

counts/g

2.5 20,480

41.75 342,016

83.17 681,329

later interfaces, is cancelled by the attenuation caused by the 

impedance match between interface and VCO, so the net gain 

of all interface stages is 0.5 V/V.

The input amplifier in the McVCO usually has a gain 

of 100 V/V. This was reduced to 10 V/V with jumpers on the 

input stage (McChesney, 1999, p. 31). After filtering (4thorder Butterworth, low-pass, 30-Hz filter) the signal splits into 

low- and high-gain channels. The low-gain channel is modified to have a gain of 0.5 V/V after the filter; consequently the 

total gain is 2.5 V/V from the interface input to the low-gain 

channel output. The maximum input signal swing of the channel’s analog to digital converter is 2.5 V. A signal of 1 V from 

the sensor ( 0.1 g) produces a full-scale measurement on the 

low-gain channel.

The amplification of the high-gain channel after the filter 

stage depends on the gain switch setting and the microcontroller program version (McChesney, 1999, p. 15, p. 30–31). 

The high-gain channel after the filter stage was initially set 

for 16.63 V/V for a total gain of 83.17 V/V from the interface 

input to the high-gain channel output. As stated previously, 

this was reduced to 8.35 V/V for a total gain of 41.75 V/V for 

the deployments from AHAB onward (see table 3).

The McVCO modulator sensitivity is 50 Hz/V 

(McChesney, 1999, p. 30–31). The sensitivity of the discriminators used for demodulation was 0.02 V/Hz. The digitizer 

sensitivity was 819.2 counts/V. Where the total response to 

input excitation is the quotient of output counts divided by 

acceleration input, the scale factor for each channel voltage 

gain is shown in table 4.

Table 4. Scale factor for 

accelerometer channel gains.

The “Marv Lander” Platform

We discovered a significant noise problem caused by 

spider vibrations during the first NED deployment. Later 

testing showed that soft tapping anywhere on the spider 

generated high levels of noise as long as the accelerometer 

was mounted on it. Our solution was to put the sensor on a 

separate platform, the “Marv lander,” named for its developer, Marvin Couchman.

Previous experience with seismic installations at Mount 

St. Helens had shown that it was not always necessary to bury 

a seismometer to get good coupling and good noise performance. The installation at SEP before October 2004 was such an 

example. At the SEP site, warm temperatures a few centimeters 

below ground had caused rapid failures of several Sercel L4-C 

seismometers before a surface installation solved the problem. 

This approach was continued with the Sercel L22-3D installation at SEP on November 4, 2004. Other experiences with 

installations on rock, where digging was impossible, had also 

produced favorable results. In all these cases, isolation from 

surface noise was achieved by surrounding the seismometer 

with a rock pile, bags of sand, or concrete. We reasoned that a 

seismometer platform for remote deployment should provide 

significant mass around the sensor. Consequently the lander 

body is a 20-kg barbell weight (45 lb) with the piezoelectric 

sensor sealed with epoxy in the center hole (fig. 6).

The platform is coupled to the ground through three short 

legs formed by the eyebolts used to attach the rigging between 

it and the spider. Most of the spider sites had some ash cover, 

and, although we could not closely observe most of the platforms on the ground, we believe that the combination of tripod 

legs and weight helped root them in the ash and avoid tipping 

instability. The barbell weight also provided thermal mass, 

isolating the piezoelectric sensor from temperature changes 

rapid enough to produce an inband signal.

The lander rigging includes a 5-m length of 5-mm 

(3/16 in.) stainless steel cable attached to the spider’s instrument box with a ring and thimble that protect the cable from 

wear during the helicopter flight (fig. 7). At the other end, a 

10-mm (3/8 in.) rope harness attaches the lander to the cable. 

Soft rope is used to isolate the cable from the lander, eliminating cable vibrations as a source of seismic noise. Polypropylene rope was used for the first lander. When the instrument 

was retrieved after an explosion, one of the three leads was 

burned through, and hot rock was found embedded in the others. Subsequently, nylon rope was used because of its higher 

melting point and superior strength.

Separating the sensor from the spider body created some 

electrical problems. When the seismic spider is set down by 

the helicopter, the lander makes first contact with the ground. 

This can produce a strong static discharge between the lander 

  Marv lander under construction. Barbell weight stands 

on three legs formed by eyebolts that connect with rigging. 

Accelerometer is potted with epoxy inside the plastic pipe, which 

is cemented into center of the weight. USGS photo by M.R. 

Couchman, December 2, 2004.
7. Seismic-Monitoring Changes and the Remote Deployment of Seismic Stations (Seismic Spider) 139

2 s

Feb. 26, 2006

1207 UTC

Feb. 24, 2006 

0632 UTC

Figure 7. Seismic spider laid out at staging area. Instrument box is approximately 70x40x40 cm. USGS photo by S.C. Moran, June 20, 2006.

and the spider that can damage the sensor and electronics. To 

provide a low-impedance discharge path, a grounding wire 

connects one of the lander’s legs to the spider body. Discharge 

through the signal-cable shield is avoided by grounding it 

only at the spider end, and the sensor case is isolated from the 

lander body by potting it with epoxy inside a short length of 

PVC pipe (fig. 6). The signal cable is vulnerable to thermal 

and mechanical damage. Consequently, the two-pair shielded 

cable (Belden type 8723) is enclosed in thermal sleeving. 

Grounding wire and sensor cable are tied to the lander rigging.

The performance of an accelerometer mounted on the 

lander is illustrated in figure 8. The top trace is from the 

piezoelectric accelerometer installation at SEP shortly before 

it was replaced by a three-component velocity seismometer 

(Sercel L22-3D) on February 24, 2006. The bottom trace is 

from the L22 and shows an earthquake from February 26, 

2006. The two earthquakes were chosen because signals 

from other stations indicated that they were very similar, 

having approximately the same magnitude, location, and 

source mechanism. The accelerometer data were integrated 

to produce a velocity response and then filtered (2 Hz, Butterworth high-pass filter) to remove long-period noise and 

produce an instrument response comparable to the L22. 

Despite the fact that the two sensor locations differed by ~15 

m, their seismic traces share many similarities, particularly at 

the onsets. More to the point, the comparable signal-to-noise 

ratios show the effectiveness of the lander coupling.

Summary

Even though Mount St. Helens was regarded as a well-monitored volcano, the September 2004 unrest exposed weaknesses 

in the seismic monitoring network. The explosion of October 1, 

Figure 8. Comparison of numerically integrated 

response of piezoelectric accelerometer (top) and 

velocity seismometer (bottom) for similar earthquakes 

at station SEP. Accelerometer data were integrated 

to produce a velocity response and then filtered to 

remove long-period noise and produce an instrument 

response comparable to the velocity seismometer.

“Marv lander”

sensor platform

“Marv lander”

Rope harness Rope harness

“Spider” instrument package “Spider” instrument package

Cable assembly

Directional antenna

Sling line Sling line
140 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

2004, removed our near-field monitoring capacity when station 

SEP was lost, reducing the quality of hypocenter determinations. 

One or more additional close-in stations would have increased the 

chance that near-field monitoring could survive a small explosion. 

Deployment of broadband seismometers before the start of volcanic unrest would have supplemented the short-period network by 

making up for its limited dynamic range and, if positioned within 

4 km, could have detected any very long period events.

Once the eruption started, near-field sites quickly became 

too dangerous for fieldwork. However, with the seismic spider, 

we were able to improvise a solution that allowed us to establish 

close-in sites in comparative safety. The spider package not only 

enabled us to place a telemetered seismometer close to the vent 

but also to retrieve it for service or redeployment. This portability facilitated short-term studies such as AHAB on spine 4, and 

portability allowed us to adapt to changing field conditions during 

this prolonged eruption. A different choice of sensor and telemetry 

technology for remote deployment might have emerged from a 

less-hasty development process, but, given the advantage of strong 

signals in the near-field, both the sensor and analog telemetry were 

more than adequate for the task. We were particularly pleased with 

the performance of the Marv lander. It eliminated the spider platform noise and produced good coupling for the sensor. With the 

possible exception of the glacier site (WESG), there were no signal 

problems caused by poor sensor installation in 10 deployments.

Seismic spiders are not substitutes for well-established 

seismic stations because they lack the sensitivity to detect small 

earthquakes, unless they happen to be in the near field, or to detect 

the long-period earthquakes that may well mark the start of volcanic unrest. However, once unrest has begun, they can be used to 

supplement an existing network where additional near-field monitoring is desired. These portable, remotely deployed stations can 

be built in advance, allowing a very rapid response to events. The 

seismic spider is a new tool for monitoring erupting volcanoes, a 

tool that can be used in dangerous places with comparative safety.

Acknowledgments

The authors would like to thank Richard P. Hoblitt 

for suggesting the piezoelectric accelerometer used in the 

seismic spider and supplying them during the first deployments. His perceptive review of this paper greatly increased 

its coherence. Thomas L. Murray’s review gave us a new perspective on the seismic spider design; this resulted in a more 

balanced paper and allowed us to avoid several shortcomings. 

Finally, we would like to thank Anthony Qamar for his initial 

calculation of the peak acceleration that a near-field sensor would measure and his many discussions of instrument 

response. Because of Tony, seismic spider data were “on 

scale” and calibrated.

References Cited

Allocca, J.A., and Stuart, A., 1984, Transducers—theory and applications: Reston, Va., Reston Publishing Company Inc., 497 p.

Ewert, J.W., Guffanti, M., and Murray, T.L., 2005, An assessment of volcanic threat and monitoring capabilities in the 

United States; framework for a National Volcano Early 

Warning System: U.S. Geological Survey Open-File Report 

2005–1164, 62 p.

Horton, S.P., Norris, R.D., and Moran, S.C., 2008, Broadband 

characteristics of earthquakes recorded during a dome-building 

eruption at Mount St. Helens, Washington, between October 

2004 and May 2005, chap. 5 of Sherrod, D.R., Scott, W.E., and 

Stauffer, P.H., eds., A volcano rekindled; the renewed eruption of Mount St. Helens, 2004–2006: U.S. Geological Survey 

Professional Paper 1750 (this volume).

LaHusen, R.G., Swinford, K.J., Logan, M., and Lisowski, M., 

2008, Instrumentation in remote and dangerous settings; 

examples using data from GPS “spider” deployments during the 

2004–2005 eruption of Mount St. Helens, Washington, chap. 

16 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A 

volcano rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 1750 

(this volume).

McChesney, P.J., 1999, McVCO handbook 1999: U.S. Geological 

Survey Open-File Report 99–361, 48 p. [http://wrgis.wr.usgs.

gov/open-file/of99-361/, last accessed Dec. 26, 2006].

Moran, S.C., 2004, Seismic monitoring at Cascade volcanic 

centers, 2004; status and recommendations: U.S. Geological 

Survey Scientific Investigations Report 2004–5211, 28 p.

Moran, S.C., Malone, S.D., Qamar, A.I., Thelen, W.A., 

Wright, A.K., and Caplan-Auerbach, J., 2008a, Seismicity associated with renewed dome building at Mount St. 

Helens, 2004–2005, chap. 2 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Moran, S.C., McChesney, P.J., and Lockhart, A.B., 2008b, 

Seismicity and infrasound associated with explosions at 

Mount St. Helens, 2004–2005, chap. 6 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Schloss, F., 1993, Accelerometer noise: Gaithersburg, Md., 

Wilcoxon Research, Inc., Sound and Vibration instrumentation 

reference issue, March 1993, 2 p. [http://www.wilcoxon.com/

knowdesk/accelnoise.pdf, last accessed Dec. 26, 2006].

Schloss, F., 1998, Piezoelectric accelerometer specifications 

and specmanship: Gaithersburg, Md., Wilcoxon Research, 

Inc., Sound and Vibration, February 1998, 2 p. [http://www.

wilcoxon.com/knowdesk/piezo_spec.pdf, last accessed Jan. 

8, 2007].

Scott, W.E., Sherrod, D.R., and Gardner, C.A., 2008, Overview of 2004 to 2006, and continuing, eruption of Mount St. 

Helens, Washington, chap. 1 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).
Heading 1 141

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142 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 143

Geological Observations of Lava-Dome 

Growth

Seismic unrest at Mount St. Helens that started in September 2004 

evolved into a dome-building eruption whose first new lava erupted 18 days 

later, on October 11. The dome grew as a series of spines within the existing 

1980 crater. Each spine impinged upon the 1980–86 lava dome, on older parts 

of the new dome, and, in some cases, on the adjacent crater wall. Invariably, 

the stress of pushing caused the active spine to shear off, override previously 

extruded rock, and ultimately disintegrate. The numbered spines marked 

sequential growth episodes, whereas the vent itself never varied in location.

Crater Glacier, which was born in the deep shadows of the 1980 Mount 

St. Helens crater and grew in the years thereafter, was riven into two arms by 

the new dome, forced aside, doubled in thickness, and accelerated downslope. 

The glacier provided little resistance to the extruding lava. By being pushed 

away, the glacier was spared substantial melting—it lost only 10 percent of its 

volume despite years-long proximity to hot rock in its upslope reach.

On-the-ground field work within the crater was rarely undertaken in 

2004–5, because of the persistent threat of small explosions. Consequently, 

most geologic monitoring was done through photography and periodic dredging of samples by helicopter. Some photographs were taken during helicopter 

overflights, others by fixed cameras (some of which transmitted images in 

near-real time), and overhead aerial photographs were taken from airplanes.

Through photogrammetry, the overhead aerial photographs provided digital elevation models of the new dome. From these digital data came the popular hillshade relief maps, deformation maps, and interpretive geologic maps 

and cross sections that illustrated the sequence of spine growth and decay. 

Successive digital elevation models also allowed the calculation of changing 

dome volumes and extrusion rates.
144 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

View east on October 12, 2004, to the first spine of the 2004 dome sequence 

soon after it breached the Crater Glacier and crater-floor debris. Spine is 

light-gray feature mostly encased in steam. USGS photo by J.C. Wynn. 
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 8

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, Box 25046, Denver Federal Center, Denver, CO 

80225

Use of Digital Aerophotogrammetry to Determine Rates 

of Lava Dome Growth, Mount St. Helens, Washington, 

2004–2005

By Steve P. Schilling1, Ren A. Thompson2, James A. Messerich2, and Eugene Y. Iwatsubo1

Abstract

Beginning in October 2004, a new lava dome grew on the 

glacier-covered crater floor of Mount St. Helens, Washington, 

immediately south of the 1980s lava dome. Seventeen digital 

elevation models (DEMs) constructed from vertical aerial 

photographs have provided quantitative estimates of extruded 

lava volumes and total volume change. To extract volumetric 

changes and calculate volumetric extrusion rates (magma 

discharge rates), each DEM surface was compared to preeruption DEM reference surfaces from 1986 and 2003. Early in the 

2004–5 eruption, DEMs documented deforming glacier ice 

and crater floor that formed a prominent “welt” having a volume of 10×106

 m3 and a growth rate of 8.9 m3/s before dacite 

lava first appeared at the surface on October 11, 2004. Afterward, the rate was initially 5.9 m3

/s but slowed to 2.5 m3/s by 

the beginning of January 2005. During 2005, the extrusion rate 

declined gradually to about 0.7 m3

/s. By December 15, 2005, 

the new dome complex was about 900 m long and 625 m wide 

and reached 190 m above the 2003 surface. More than 73×106

m3 of dacite lava had extruded onto the crater floor.

Successful application of aerophotogrammetry was possible during the critical earliest parts of the eruption because we 

had baseline data and photogrammetric infrastructure in place 

before the eruption began. The vertical aerial photographs, 

including the DEMs and calculations derived from them, were 

one of the most widely used data sets collected during the 

2004–5 eruption, as evidenced in numerous contributions to 

this volume. These data were used to construct photogeologic 

maps, deformation vector fields, and profiles of the evolving dome and glacier. Extruded volumes and rates proved to 

be critical parameters to constrain models and hypotheses of 

eruption dynamics and thus helped to assess volcano hazards.

Introduction

The volume of a growing lava dome and its extrusion rate 

are primary measurements that may be compared with other 

traditional volcano-monitoring data from ground deformation, 

gas geochemistry, and seismicity for the purpose of monitoring and studying erupting volcanoes. Such comparisons have 

been made for some recent dome-building eruptions, including Santiaguito, Guatemala (Harris and others, 2003); Unzen, 

Japan (Nakada and others, 1999); Redoubt, Alaska (Miller, 

1994); and Soufrière Hills, Montserrat (Sparks and others, 

1998). In addition to their value as a fundamental dataset, volumetric data are needed to explore such linkages as extrusion 

rate thresholds for transition to explosive activity and volume 

thresholds for initiation of large-scale dome collapse. In this 

paper, we describe a new application of traditional photogrammetric techniques to track the growth of the 2004–5 Mount St. 

Helens lava dome.

In October 2004, a new period of dome growth began 

at Mount St. Helens that changed the topography of the 

1980 crater dramatically (fig. 1). Between October 2004 and 

December 2005, more than 73×106

 m3 of solid dacite lava 

extruded onto the crater floor immediately south of the lava 

dome that had formed in the 1980s. The new dome grew as a 

succession of large spines (Vallance and others, this volume, 

chap. 9). Dome growth intensely deformed and divided Crater 

Glacier, which had developed from a small debris-covered 

snowbank in 1985 into a glacier covering about 1 km2

 by 2001 

(Schilling and others, 2004). We use the informal names “west 
146 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Crater Glacier” and “east Crater Glacier” for the two ice bodies remaining within the crater. The growing dome compressed 

the east Crater Glacier against the east crater wall, resulting in 

spectacular thickening. Compression and thickening created 

crevasses and rapid advance of the terminus, about 185 m in 

two years, followed by similar effects on the west Crater Glacier (Walder and others, this volume, chap. 13).

The extruded volume and extrusion rate of lava associated with this protracted dome eruption are critical parameters 

used to constrain models of the magmatic system (Mastin and 

others, this volume, chap. 22) and eruption dynamics (Iverson, this volume, chap. 21), to determine how the magmatic 

system relates to surface measurements of magmatic gases 

(Gerlach and others, this volume, chap. 26) and seismicity (for 

example, to determine if extrusion rate correlates with size 

and rate of occurrence of volcano-related earthquakes; Moran 

and others, this volume, chap. 2), and to constrain calculations 

east

Crater 

Glacier

west

Crater 

Glacier

1980s lava

dome

2004–5

dome

A

B

that address loading effects of the growing dome on surface 

deformation (Lisowski and others, this volume, chap. 15).

Photogrammetry based on vertical aerial photographs 

has been used previously to monitor, model, map, and measure surface change and deformation at volcanoes (Achilli 

and others, 1998; Baldi and others, 2000, 2005; Zlotnicki and 

others, 1990). A recent photogrammetric study of the Mount 

St. Helens crater (Schilling and others, 2004) tied a block of 

overlapping vertical aerial photographs to a network of global 

positioning system (GPS) stations on the volcano’s flanks, 

dome, and crater floor (fig. 2A) resulting in a digital elevation 

model (DEM) of the volcanic edifice and entire crater configuration in October 2000.

The 2000 DEM has served as a baseline for comparison 

with past DEMs. Comparison of the 2000 surface with postMay 18, 1980, and 1990 DEMs, both derived from existing 

topographic contour maps, produced volume estimates of 

Figure 1. Oblique photographs of Mount St. Helens crater from east 

rim; views to southwest. A, Preeruption crater on August 30, 2004, 

showing 1980s lava dome and Crater Glacier (mostly coated with rock 

debris). B, Crater on October 12, 2005, showing new lava dome dividing 

and deforming glacier. USGS photographs by S.P. Schilling.
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 147

46 ̊12'

46 ̊10'

46 ̊14'

46 ̊16 '

122 ̊14' 122 ̊12' 122 ̊10' 122 ̊8'

JRO1

0 1 2 4 KILOMETERS

46 ̊12'

122 ̊12' 122 ̊10'

0 0.5 1 2 KILOMETERS

A B

Br

Br

glacier ice and talus accumulation, as well as crater-wall 

erosion (Thompson and Schilling, 2007). Near the beginning 

of the 2004–5 eruption, two other craterwide DEMs were 

constructed to estimate volume change. One was created from 

digital, photogrammetrically derived contours that depicted 

the crater in November 1986, and another was constructed 

from lidar data (Queija and others, 2005) obtained September 

20–22, 2003, using the same network of GPS sites for ground 

control that was used for photogrammetric monitoring.

Throughout the 2004–5 eruption, we used analytical 

photogrammetry and a softcopy (that is, digital image) system 

to provide stereo imaging and accurate measurement of the 

rapidly changing crater morphology. Seventeen DEMs have 

been constructed from vertical aerial photographs (Messerich 

and others, 2008), collected approximately every three weeks, 

in order to quantify volumetric changes associated with dome 

growth and collapse as well as deformation of Crater Glacier 

(Walder and others, this volume, chap. 13). Each new DEM 

was compared to preeruption reference surfaces of the crater 

in 1986 and 2003 to extract volumetric changes. For viewing 

purposes, each DEM was also converted to a hillshade-relief 

map, in the form of digital raster images (tagged image file 

format) and corresponding georeferenced world files (ASCII 

format). These raster maps are found in appendix 1, which 

is available in the DVD that accompanies this volume and in 

online versions of this chapter.

Methods

Acquiring low-altitude aerial photographs in inclement 

weather over mountainous terrain presents logistical challenges, and taking them over an erupting composite volcano 

such as Mount St. Helens, where rapid turnaround of data is 

needed to evaluate hazards, adds urgency. These problems 

were simplified by taking advantage of photogrammetric infrastructure set up during previous work at the volcano (Schilling 

and others, 2004). When the 2004–5 eruption began, however, some established flight-planning procedures required 

modification, such as selection of new ground control sites 

and design of flight lines for appropriate photograph scale 

Figure 2. Shaded-relief maps of preeruption surface of Mount St. Helens (2003 DEM). Red and blue boxes show approximate areas 

covered by two overlapping, 1:12,000-scale vertical aerial photographs. A, Broad GPS ground control network (black triangles) 

used for aerotriangulation of block of photographs to extract coordinates for construction of 2000 DEM and to serve as control for 

2003 lidar-based DEM. Br, Brutus control point on east rim. B, Ground control points for vertical aerial photographs obtained during 

2004–5 eruption. Black X’s, ground control points destroyed early in eruption. Red dots, photoidentified control points transferred from 

2000 aerotriangulation for use early in eruption. Black squares, new, and black triangles, existing ground control points measured 

in summer 2005. Yellow triangles, photoidentified control points transferred from August 2005 aerotriangulation solution to serve as 

ground control for December 2005 aerotriangulation.
148 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

and overlap. Early in the eruption, having procedures already 

established meant that DEMs were constructed rapidly, at 

times within a few days of obtaining the aerial photographs.

Digital photogrammetry refers to photogrammetric systems that have been adapted from hardcopy (film) 

aerial photographs mounted on optical-mechanical stereo 

plotters to softcopy (scanned, digital) aerial photographs 

imported into a desktop computer workstation. Hardware 

for our softcopy system includes a fast central processing 

unit, abundant memory, two display screens (one for stereo 

images and one for control menus), polarizing spectacles for 

viewing onscreen images in stereo, and a customized mouse 

to control three-dimensional (3D) cursor movement. Socet 

Set software includes spatial resection and spatial intersection algorithms to calculate X, Y, and Z ground coordinates 

of features observed in stereo aerial photographs; it requires 

precise camera calibration parameters, carefully measured 

photograph coordinates, and accurate GPS ground control 

(Wolf and Dewitt, 2000; Thompson and Schilling, 2007). 

The softcopy system automates many standard photogrammetric tasks, such as inner and relative orientations of stereo 

models. However, the critical task of exterior orientation, 

which ties photographs to ground control, requires a skilled 

operator to maintain precision. Assuming minimal operator 

error, the accuracy of the final model solution is constrained 

by flight height, which determines scale, and by scanning 

resolution, which determines the minimum feature size that 

the operator can discern in the aerial photographs.

After considering many factors, such as safety and size of 

features to be measured, we selected a flying height of 3,900 

m (12,800 ft), yielding a nominal photo scale of 1:12,000 with 

a 153.681-mm lens and 9×9-inch film format. Thus, 1 mm 

on the photograph at nominal scale represents 12 m on the 

ground. This nominal scale applies at about the altitude of the 

Crater Glacier surface immediately south of the 1980s lava 

dome, roughly midway through the range of relief within the 

Mount St. Helens crater—from about 2,540 m (8,330 ft) on 

the crater rim to about 1,815 m (5,955 ft) near the base of the 

1980s dome. The variation in altitude corresponds to a variation in scale, often within a single photograph, from 1:8,964 at 

the rim to 1:13,726 at the base of the 1980s lava dome.

Each aerial photographic negative was scanned at 12 

μm resolution, generating a graytone digital image roughly 

350 megabytes in size. Owing to the scale variation within a 

photograph, each cell in the image may represent a different 

distance on the ground, from 0.108 to 0.165 m.

Global Positioning System (GPS) Control 

Network

Accurate ground control is equally as important as highresolution scanning. About 30 GPS sites on the outer flanks, 

1980s lava dome, and crater floor of Mount St. Helens (fig. 

2A) form a network of ground control points for photogrammetric surveys. The sites were selected to accommodate the 

geometry needed for the exterior orientation (linking world 

and photograph coordinate systems) of blocks of overlapping 

photographs (multiple stereopairs along and between adjacent flight lines); each site was also accessible by helicopter. 

Most of the sites are part of a geodetic network established 

in 2000 (Dzurisin, 2003). Many are located near 3-m-high 

steel towers constructed after the 1980 eruption and used for 

electronic distance meter (EDM) surveys (Swanson and others, 1981). These towers or their shadows are relatively easy 

to find in the diapositive (film positives) rendering of aerial 

photographs used to help locate ground-control sites in the 

digital images. Using a helicopter and two crews, we were 

able to measure and place photo-targets at all of the stations 

in 1–2 days.

Measuring GPS locations in the field for ground control 

at Mount St. Helens is a straightforward task. We deploy a 

GPS receiver at each site, centering the antenna over a specific 

tower leg, piece of rebar, or benchmark. GPS receivers collect 

data at each site for about 1 hour. Before moving to the next 

site, three or four large plastic rectangular panels are placed 

radially about each GPS point to facilitate locating the control 

point in aerial photographs. The receiver data are downloaded 

and processed using Ashtech Office Suite for Survey ® 

(AOSS) software. The reference station for differential GPS 

processing is JRO1 (fig. 2A), which at the start of the eruption was the nearest continuous GPS station (approximately 

9 km north of the crater). Altitudes from AOSS are converted 

to orthometric heights using the National Geodetic Survey 

(NGS) program GEOID03. The U.S. Army Corps of Engineer 

program CORPSCON is used to calculate X, Y, and Z coordinates using a horizontal datum of NAD83 and a vertical datum 

of NAVD88. The GPS data yield position accuracies of 2 to 3 

cm in X and Y (horizontal) coordinates and 6 to 7 cm in the Z 

(vertical) coordinate.

Control points were remeasured annually from 2000 to 

2004 before taking aerial photographs of the entire volcano. 

Repeated measurement improved and validated the positional accuracy of most sites. However, the landscape evolves 

quickly at Mount St. Helens, and site positions change, at 

times dramatically. For example, in 2002 the photo panels and 

rebar of the Brutus site on the east crater rim (fig. 2, station 

Br) were buried beneath about 20 cm of wind-blown pumice. 

The following year, the site toppled into the crater and had to 

be replaced.

New Ground Control Based on GPS Data and 

Photo-Identified Points

When the 2004–5 eruption began, four established flight 

lines used in previous studies to capture aerial photography 

of most of the volcano were abandoned in favor of a single, 

south-to-north flight line centered over the crater. A single 

flight line offered advantages of safety, simplicity, speed, 

and economy. However, it eliminated use of ground control 

points on the east and west outer flanks of the volcano and 
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 149

forced reliance on control points on the crater rim, floor, 

outer south flank, and 1980s lava dome. Unfortunately, 

explosions early in the eruption eliminated three critical 

ground control sites on the 1980s lava dome (fig. 2B). Several of the remaining control points were outside the stereo 

coverage of the single flight line.

Rather than increase the number of flight lines and 

incorporate the ground control on the volcano flanks in an 

aerotriangulated block or replace the ground control in the 

crater’s potentially dangerous environment, we transferred 

control points from the stored orientation solution (aerotriangulation) of the 2000 block of aerial photographs. Photoidentifiable, measured points (fig. 2B), such as distinct boulders or 

topographic prominences, were carefully selected and passed 

from the aerotriangulated solution of 2000 models to the 2004 

aerial photographs as ground control. These points or a subset 

of them provided control for successive sets of photographs 

obtained throughout the winter of 2004–5. The resulting 

ground-control accuracy was reduced from a few centimeters, 

based solely on GPS data, to a few decimeters using photogrammetrically transferred control points.

In July 2005 we remeasured and repaneled the network 

of existing ground-control points on the crater rim and flanks. 

In addition, we reestablished one of the points on the west 

side of the 1980s lava dome and added points on the crater 

floor north of the 1980s lava dome and a site on the south 

crater rim (fig. 2B).

Aerotriangulation

Aerotriangulation provided a means of transferring 

accurate control from earlier photogrammetric work at 

Mount St. Helens, rather than establishing new control in 

hazardous areas as the eruption progressed. Aerotriangulation refers to solving relative orientation equations for 

overlapping aerial photographs (stereo model), identifying 

common points (pass points) between adjacent models to 

form continuous strips or blocks of stereomodels, and solving simultaneous equations to adjust mathematically the 

strips or block of photographs to ground control. A series 

of nonlinear expressions with many unknowns are truncated 

by Taylor’s theorem into a series of linear equations that are 

solved simultaneously by the method of least squares (Wolf 

and Dewitt, 2000). Our 2000 benchmark DEM was created 

from an aerotriangulation solution for a block of aerial photographs. The block comprised four adjacent, overlapping 

strips of photos; each strip was formed from seven or eight 

overlapping photographs. Point coordinates extracted from 

the 2000 stereomodels and aerotriangulation were used as 

control points for successive stereomodels and for aerotriangulation solutions during the early part of the 2004–5 

eruption. The resulting root-mean-squared (RMS) residual 

error for the early aerotriangulation solutions was a few decimeters. New and remeasured GPS ground-control points in 

the aerial photographs taken in the summer and fall of 2005 

resulted in an aerotriangulation RMS residual error of several 

millimeters to several centimeters. The average RMS error 

from all aerotriangulation solutions was 0.17 m in the X, Y, 

and Z coordinates. Extending that error over the area that 

has been affected by the eruption (about 1 km2

) gives rise to 

a volume of uncertainty of 1.7×105 m3. This error is about 4 

percent of the typical monthly extruded volume (4×106 m3) 

estimated from comparison of the series of DEMs (Iverson 

and others, 2006, Supplementary Notes).

Digital Elevation Model (DEM) Construction

We used spatial resection calculations to derive an 

aerotriangulation solution and then used spatial triangulation 

algorithms to extract three-dimensional coordinates and construct a DEM that defines the crater surface. Two techniques 

were used for obtaining X, Y, and Z coordinates from the 

stereomodels.

The first method was a manual technique that relies on 

the skill of the operator. The operator examined one or more 

stereomodels and identified the area being deformed and 

the features to be measured. For each feature, the operator 

carefully placed a floating mark on the feature, stored the Z 

(elevation) coordinate, and triangulated and stored the X and 

Y (planimetric) coordinates. Points were collected individually 

or as a stream as the operator moved the floating mark along 

the terrain.

The second method was an automated technique in 

which Socet Set software used coplanarity equations to 

derive a plane intersecting three points: (1) an object in 

one image; (2) the same object in the overlapping area of a 

second, adjacent image; and (3) the triangulated position of 

the object on the ground. The software used the line formed 

where the plane intersected the two adjacent photographs 

to guide its search for matching cell patterns. When the X 

and Y coordinates of matching cell patterns were identified, the software calculated the corresponding Z coordinate. The operator defined an area and density of points for 

the automated calculations. The automated method did not 

work well in steep terrain but did work in the relatively flatlying, glacier-covered parts of the crater, provided there was 

enough contrast (such as a dusting of ash) in the photographs 

to provide unique cell patterns. The automated technique 

identified, calculated, and stored X, Y, and Z coordinates for 

many locations relatively quickly. This technique was used 

sparingly, however, because it can be labor intensive. Automatically generated points must be checked by the operator, 

either individually or by generating contours of groups of 

collected points, to correct any errors.

Using these two techniques, individual points or streams 

of points marking breaks in slope (breaklines) were collected 

to better define topographic inflections, enabling intervening surfaces on the growing dome and deforming glacier 

to be represented by significantly fewer data points. The 

resulting three-dimensional surface is an array of triangular 
150 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

facets referred to as a triangulated irregular network (TIN), 

in which interpolated surfaces are triangles having measured points at each vertex. A TIN surface was checked for 

systematic errors or random operator error in a few seconds 

by generating digital contours for the surface. The operator 

easily located and repaired errors, such as points that caused 

contours to cross, and recalculated the TIN surface.

After error checking and visually determining that the 

TIN was an accurate representation of the three-dimensional 

surface depicted in the aerial photographs, the TIN data 

structure was stored and converted to a DEM. The DEM 

differs from the TIN in that the former is a regular array of 

square cells (or rasters), where each cell represents an area of 

the Earth’s surface (X, Y) with a specific altitude value (Z). 

During this conversion process, the cell size or resolution 

did not dictate the accuracy of an altitude value. Rather, the 

resolution dictated how closely the DEM represents the TIN 

model. The accuracy of altitudes and horizontal positions 

was determined by the combined photogrammetric orientations, operator’s skill, and in particular by the accuracy of the 

ground control.

Sources of error in the DEM construction included aerial 

camera calibration, film processing, flight parameters, ground 

control points, conversion of photos to digital form, stereomodel orientation (interior, relative, absolute), aerotriangulation, image-matching algorithms, operator bias, and random 

factors (Daniel and Tennant, 2001, p. 402–403). We estimated 

the uncertainty of any volume measurement to be a function 

of the area of the growing dome multiplied by the average 

RMS residual error (0.17 m). Thus, as the volume continued 

to increase, so did the estimated uncertainty of the volume calculation. However, the resulting uncertainty in volume-change 

calculations was about 4 percent—small compared to the 

ambiguity introduced by the unknown subsurface configuration of the lava dome (the lower part of the new dome masked 

by Crater Glacier).

Surface Depiction Using DEMs

The DEMs were imported into an ArcInfo Geographic 

Information System (GIS) for viewing and analysis after conversion to an ASCII text file in the softcopy system. This file, 

which was formatted for import to ArcInfo software, included 

a header containing the number of columns, number of rows, 

cell size, and X and Y coordinates of the lower left corner of 

the DEM, followed by a sequence in row-major order (top row 

first, bottom row last) of all the Z (elevation) values. The file 

was imported into the ArcInfo grid module as a high-resolution (2 m) grid (raster data structure).

Perhaps the most significant and primary use of any photogrammetrically constructed DEM is to examine the surface, 

either singly or in sequence with previous DEMs. Such a succession of DEMs for Mount St. Helens shows, qualitatively, 

the changing position, dimensions, and size of the growing 

dome and deforming glacier over time. The DEMs are difficult 

to examine or interpret directly. A hillshade algorithm was 

used to position an artificial “sun” at an arbitrary azimuth and 

altitude to render a gray-scale shaded-relief view of the changing dome and glacier. This rendition of the topography is free 

of potentially distracting details seen in the aerial photographs, 

such as steam, snow, or ash, and allows consistent and simplified viewing to study geomorphic change and compare cell 

alignment among DEMs.

The areal extent of the dome complex and deformed 

glacier increased throughout the 2004–5 eruption, forcing 

a corresponding increase in the extent of DEMs to document the change. To ensure proper alignment (registration) 

of the DEMs, they are cast within the same projection, same 

horizontal datum, and same vertical datum. We selected the 

Universal Transverse Mercator (UTM) projection’s zone 

10, which extends from 120° to 126° west longitude, the 

1988 North American Vertical Datum (NAVD 88), and the 

1983 North American horizontal datum (NAD 83). Previous DEMs derived from contours having a 1927 horizontal 

datum (NAD 27) were converted to NAD 83. In addition, a 

regional vertical correction of 1.25 m (Zilkoski and others, 

1992) was applied to convert the 1929 vertical datum (NAGD 

29) to NAVD 88. We confirmed this regional correction by 

calculating the difference between GPS control points stored 

in NAGD 29 and NAVD 88.

The extent of each DEM was varied in order to capture 

dome growth. Dome growth and the DEMs depicting the 

growth were bounded on the north by the relatively fixed 

1980s lava dome and eventually bounded on the south by 

the south crater wall. The DEM edges on the east and west 

(as well as the south early in the eruption) were delineated 

visually to capture observed deformation, using the softcopy 

system. These results were confirmed by using GIS software 

to compare any current DEM to earlier DEMs. If we found 

that any of the DEMs had inadequate extent to represent the 

entire deformation field, the boundary could be extended using 

the scanned aerial photographs and aerotriangulation stored 

within the softcopy system.

Initially, the 2000 DEM was used as the baseline data 

set for all comparisons in order to keep the construction 

method consistent. However, the craterwide 2000 DEM 

had a cell size of 10 m, whereas DEMs constructed for the 

2004–5 eruption had cell sizes of 2 m. The difference in cell 

size meant that each of the newer DEMs would have to be 

resampled to give a 10-m cell size. The error introduced by 

resampling outweighed the benefit of using consistent production methods; therefore, we later selected the craterwide 

2003 DEM derived from lidar, and a 1986 DEM derived 

from contours, each having a 2-m cell size, for comparison 

with all subsequent DEMs.

Volume Calculation Methods

Although a single DEM surface offers a quick, synoptic 

portrayal of ongoing surface deformation, the comparison of 

successive DEM surfaces can yield estimates of erupted vol-
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 151

ume and average extrusion rate. We wrote software that uses 

GIS functions to subtract one DEM surface from another 

and generates a third, isoline surface showing net elevation 

change. The isoline surface stores the calculated difference 

between its two parent surfaces, retaining their 2-m cell size. 

In this manner, each cell of the isoline grid stores a positive, zero, or negative value resulting from the subtraction 

of the two corresponding parent cells. After the subtraction, 

the software sums values of all cells having a negative value 

(melting, erosion, or subsidence) and of all cells having a 

positive value (extrusion, deposition, or uplift) separately, 

and it also multiplies each total positive or total negative value by the area of a single cell (4 m2

). The software 

writes the resulting volumes to a text file and generates two 

additional grids, one showing the location of the positive and 

another showing the location of the negative results.

The potentially straightforward task of determining 

extruded volume from total-volume changes was complicated 

by the presence of Crater Glacier, in some places more than 

150 m thick (greater than 200 m thick when including subglacial 1980–86 crater-floor deposits), which was displaced and 

severely disrupted as the eruption progressed. Two questions 

arose: (1) Does extrusion begin at the preexisting crater floor 

or at the glacier surface; and (2) how best to calculate erupted 

volume as a function of time? One approach (fig. 3A) is to difference each new DEM with the 2003 preeruption DEM and 

to use the total surface-volume change as a proxy for erupted 

volume. For example, 1 m2

 of glacier ice rising 2 m represents 

2 m3 of lava extruded beneath it or laterally displacing it. This 

approach has the advantage of accounting for all material that 

rose above the crater floor, some of which remains obscured 

by ice, and the method is simple and straightforward. Three 

disadvantages of this total volume-change technique are the 

unknown total amount of dilatation caused by ice deformation (though the largest crevasses are captured during DEM 

construction), volume gain by winter snow accumulation, and 

volume loss by melting.

Another approach (fig. 3B) is to assume that lava extends 

from the lava-ice contact at the surface vertically downward to 

the 1986 crater floor. This is a reasonable approach for three 

reasons: (1) field examples of ice-contact lava flows have 

steep sides that formed as they flowed against steep ice walls 

(Lescinsky and Fink, 2000); (2) where observed around the 

Mount St. Helens dome, ice-lava contacts are nearly vertical; and (3) this method is also simple and straightforward. 

One disadvantage is that this method ignores observed rising 

crater-floor material or glacier, especially early in the eruption, 

which probably deform in response to endogenous or subglacial lava emplacement.

Our solution was to use both approaches to track volume estimates. The first approach provided measurements of 

extruded rock independent of whether rock broke through the 

glacier surface. As the eruption proceeded, the dome grew 

larger and the impact of the glacier became proportionally 

smaller. The discrepancy between the two approaches therefore decreased over time.

Figure 3. Diagrams (A, B) illustrating DEM differencing for 

volume-change calculations. Semitransparent oblique cutaway 

of surface defined by April 19, 2005, isolines (colored bands) 

draped on DEM showing 1986 ground topography and 2003 

glacier surface. Yellow arrows, examples of local elevation 

changes. View to north-northwest. A, Total volume change. 

Note that differencing includes marginal areas, unlikely to be 

underlain by lava, that are composed of deformed glacier. B, 

Extruded-lava volume. Blue bars define inferred steep margins 

of extrusion. C, DEM and differencing surface of April 19, 2005. 

The DEM and differencing surface south of the line of section 

were removed to produce oblique views A and B.

d

A

B

d

C Apr. 19, 2005
152 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Results

A sequence of 17 pairs of images derived from DEMs 

of October 4, 2004, through December 15, 2005, document lava dome growth and glacier deformation during the 

2004–5 eruption (figs. 4, 6–8). For each pair, the left image 

is a shaded-relief map. Beginning with the October 13, 2004, 

map, a red line marks the approximate boundary of extruded 

lava on the surface that was extended vertically to the 1986 

crater floor to calculate the extruded-lava volume. The right 

image is an isoline map draped over the shaded-relief image. 

Each 10-m interval of the isoline map is assigned a unique 

color to illustrate better the magnitude and location of elevation change relative to the 2003 crater surface. All isoline 

cells having an elevation difference equal to or greater than 

10 m were summed to calculate total volume change. Some 

of the shaded-relief images, such as that for October 11, 2004 

(fig. 4E), have triangular facets that portray areas where a 

Figure 4. Shaded-relief images showing preeruption crater surface and growth of welt and initial lava extrusion (red outline 

in G). A, 1986 and B, 2003. Blue line, approximate boundary of 1980s lava dome. Red dots show altitude, in meters, of points 

on crater floor. Increases (in parentheses) caused by growth of Crater Glacier and deposits eroded from crater walls. Area 

labeled Opus is part of the 1980s lava dome involved in uplift in late September and early October 2004. C–H, sequence of 

three pairs of images from October 2004. Left image is shaded relief map; right image has 10-m isolines showing topographic 

changes relative to September 2003 (shown in B).
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 153

condensed steam plume in the aerial photographs prevented 

detailed coordinate collection.

Also shown for reference in figure 4 are shaded-relief 

images of the craterwide datum surfaces from 1986 (fig. 4A) and 

2003 (fig. 4B). The shaded-relief map of the 1986 crater shows 

the configuration of crater walls, talus fans and other deposits on 

the crater floor, and the 1980s lava dome. The 2003 shaded-relief 

map also shows the newly formed glacier, which buries some 

margins of the 1980s lava dome. Spot altitudes show the magnitude of surface-elevation change from 1986 to 2003.

As the new lava dome grew, questions about the volume 

of the 1980s lava dome arose. We differenced DEMs to estimate a volume of 92×106

 m3 for the 1980s lava dome (Thompson and Schilling, 2007), which is larger than the 74×106

 m3

estimate of Swanson and Holcomb (1990) but is in agreement 

with volume calculations of Mills (1992). The difference in 

volume estimates may result from (1) differences in measurement methods, (2) exclusion by Swanson and Holcomb (1990) 

of crater-wall debris that was incorporated into the dome as 

it grew (Mills, 1992), and (3) different configurations for the 

base of the lava dome as a horizontal versus sloping surface.

The following discussion of the DEMs focuses on four 

time intervals and highlights key events related to dome 

building and deformation of Crater Glacier: (1) Growth 

of the so-called welt (Dzurisin and others, 2005), south of 

the 1980s lava dome, and its migration southward in late 

September and early October 2004 involved uplift of a part 

of the 1980s lava dome (area labeled Opus on figs. 4A, 4B), 

Figure 4—Continued.
154 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

crater-floor debris, and glacier ice. Initial lava spines 1 and 

2 (see Vallance and others, this volume, chap. 9, for description of the lava-spine numbering scheme) rose through the 

northern part of the welt near the end of this interval. (2) 

From late October 2004 through mid-April 2005, extrusion, 

growth to the south and migration east, and breakup of curvilinear, smooth-surfaced spines 3 and 4, known as whalebacks, drove extraordinary deformation of the east Crater 

Glacier. The first evidence of lateral movement of the spines 

came from the DEMs. For example, the extrusion axis, a 

horizontal line oriented parallel to the direction of extrusion along the approximate center of a whaleback, initially 

moved southward, then shifted eastward from October 2004 

through April 2005 (approximately 4° horizontal rotation 

of the whaleback axis towards the east occurred between 

November and December 2004 and about 6° between February and March 2005), as successive whalebacks emerged 

from the vent area, displacing older parts of the growing 

dome. (3) Between mid-April and late July 2005, smoothsurfaced spine 5 grew at a much steeper angle than had the 

previous whalebacks and began a trend toward west-directed 

movement. (4) In August 2005 spine 6 began to grow and 

migrate westward, followed by penetration of its eastern 

part by spine 7 in early October. Both spines moved west, 

which compressed and thickened west Crater Glacier. For a 

complete description of the geologic interpretation of dome 

growth, see Vallance and others (this volume, chap. 9). For a 

complete description of glacier deformation, refer to Walder 

and others (this volume, chap. 13).

October 4 Through October 13, 2004

Recognition of the welt (Dzurisin and others, this 

volume, chap. 14) a few days after seismic unrest began on 

September 23, 2004 (Moran and others, this volume, chap. 

2), prompted efforts to obtain DEMs of the crater. The first 

group of images (figs. 4C–H) shows the remarkably rapid 

Figure 5. Photographs of 2004–5 dome. A, View of crater after 

explosion of October 5, 2004. Welt formed of uplifted glacier, 

crater-floor debris, and southeast part of 1980s lava dome. 

Light-gray ash covers most of crater. View to west. B, Recumbent 

spine 3 on November 29, 2004, after impinging upon southeast 

crater wall. Its smooth carapace and leading wave of deformed 

glacier ice inspired the descriptive term, whaleback. View to 

southwest. C, Crater on February 22, 2005, showing spine 4 flanked 

by talus and, to left, remains of spine 3. Highly deformed east 

Crater Glacier prominent in left part of photograph. View to south. 

D, Close-up view about 5 m across of gouge-covered spine 4 

extruding from vent, May 12, 2005. Striations parallel the direction 

of motion. Talus flanking spine is covered by dust from rockfalls 

and sloughing of gouge on right. View to south. 

USGS photographs by S.P. Schilling (A, C, D) and S. Konfal (B).
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 155

Date of 

photography

Total volume 

change

(x 106

 m3)

Total volume 

change rate 

(m3/s)

Extruded lava 

volume (x 106 

m3)

Lava 

extrusion 

rate (m3/s)

10/4/2004 5 1 -- 1 -- 2 -- 2

10/11/2004 10 8.9 -- 3 -- 3

10/13/2004 11 6.4 0.54 4 -- 4

11/4/2004 20 4.6 12 5.9

11/29/2004 27 3.0 21 4.4

12/11/2004 30 3.4 26 4.1

1/3/2005 35 2.4 31 2.5

2/1/2005 40 1.9 35 1.8

2/21/2005 43 1.8 39 2.4

3/10/2005 45 1.5 42 1.8

4/19/2005 55 3.0 47 1.5

6/15/2005 59 0.8 54 1.4

7/14/2005 59 -0.2 57 1.3

8/10/2005 60 0.4 62 2.0

9/20/2005 62 0.7 67 1.6

10/24/2005 66 1.3 70 0.9

12/15/2005 75 1.9 73 0.7

1

 Volume is the welt, obtained by method of differencing from topographic surface of September 2003. 

Start date for growth of welt is imprecisely known, so no rate offered.

2

 Eruption has not begun; no extruded lava.

3

 Too steamy to confidently discern limits of new extrusion.

4

 Steamy; extruded lava volume is crude estimate, and no rate is offered.

Table 1. Total volume change, extruded lava volume, and rates of change during 

2004–2005 eruption of Mount St. Helens.

[Extruded lava did not appear at the surface until October 11, 2004. Rates are calculated from the 

volume changes since the previous measurement.]

changes that occurred in 10 days as the welt grew rapidly 

upward and southward, deforming parts of the 1980s lava 

dome, crater-floor material, and Crater Glacier (fig. 5A). By 

October 11 the welt had grown to about 425 m in width, 475 

m in length, and 105 m in height (2,209 m altitude) above the 

2003 glacier surface. (Length of welt and dome is measured 

along the approximate longest dimension of each shape, and 

width is measured along a line roughly perpendicular to the 

length.) The welt attained a volume of 10×106

 m3 (table 1) as 

lava spine 1 emerged on October 11 (Vallance and others, this 

volume, chap. 9; Pallister and others, this volume, chap. 30). 

Presumably the volume of the welt approximated the volume 

of lava emplaced onto the subglacial crater floor.

November 4, 2004, Through April 19, 2005

Nine image pairs (fig. 6A–P) show the sequential growth 

of lava spines 3 and 4, which shared a similar history in terms 

of form, growth, and demise. These two striated, gouge-covered, curvilinear, whaleback-shaped spines (figs. 5B, 5C, 5D) 

grew south-southeast from the vent. The southern ends were 

pushed eastward over time and broke apart after impinging 

upon the south crater wall. Both whalebacks were surrounded 

by talus aprons on the west, south, and east. Initially the disrupted glacier adjacent to the talus aprons formed a crevassed 

ridge that, in map view, looked like the bow wave of a ship 

(figs. 6A, 6C). West of the growing dome, the talus apron, gently flexed ice, and intervening depression formed a relatively 

stable configuration throughout this period.

The November 29, 2004, image (fig. 6C) shows that lava 

spine 3 had a wide center with tapering ends. It was about 

145 m wide, 350 m long, and at its highest point 150 m above 

the 2003 surface (altitude 2,282 m). The December 11, 2004, 

image (fig. 6E) shows that a longitudinal fracture had formed 

and broken the whaleback into two parts. By January 3, 2005, 

prominent longitudinal and oblique transverse, northwest and 

north-northeast-striking fractures had broken spine 3 into many 

blocks, presumably as a result of the spine impinging on the 

south crater wall (Vallance and others, this volume, chap. 9). 

The resulting 31×106

 m3 dome was about 500 m long by 200 m 

wide and reached 184 m above the 2003 crater surface (altitude 

2,293 m). East Crater Glacier was squeezed between the grow-
156 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 6. Sequence of eight pairs of shaded-relief images (A–P) showing growth of lava dome complex (red outline) and 

deformation of Crater Glacier from November 2004 to April 2005. Arrows (panel I) shows prominent, arcuate step on east Crater 

Glacier. For full explanation, see figure 3 and text.

ing dome and the southeast crater wall (figs. 6C–H), becoming 

severely disrupted as it thickened more than 120 m into a conical form by December 11, 2004 (figs. 6E, 6F). The highest area 

of disrupted ice shifted about 100 m to the northeast by January 

3, 2005 (figs. 6G, 6H). The west Crater Glacier experienced 

broad uplift of about 10 m near the 1980s lava dome and near 

the south crater wall (figs. 6D, 6F, 6H).

A second whaleback (spine 4) began forming in early 

January 2005 from the severed stump of spine 3. By February 1 it was 145 m wide, 320 m long, and 210 m above the 

2003 surface (altitude 2,331 m) (elevation of corresponding 

point on 2003 surface varies slightly as location of high point 

of whalebacks changes) (figs. 6I, 6J). These dimensions are 

similar to those of spine 3 on November 29, 2004. Spine 4 

grew south-southeastward to about 450 m in length, 150 m in 

width, and 210 m (altitude 2,343 m) above the 2003 surface on 

March 10, 2005 (figs. 6M, 6N). As with spine 3, longitudinal 

fractures cut the southwest edge of spine 4 (fig. 6M). In April 

2005, longitudinal, northwest-striking fractures and oblique 

transverse, northeast-striking fractures broke spine 4 as it 

impinged on the south crater wall, forming rubble and several 

megablocks (fig. 6O).
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 157

Shoestring

notch

Shoestring

notch

Figure 6—Continued.

The corresponding isoline maps for this time period 

show the response of the glacier to the growing dome (figs. 

6J, 6L, 6N, 6P) as an increasing number of en echelon, 

east-trending, transverse crevasses cut east Crater Glacier. 

Apparent counterclockwise rotation and thickening of the 

glacier west of Shoestring notch to more than 110 m above 

its former surface in 2003 was in response to the domeglacier margin moving eastward. In March 2005, a second 

small area of glacier ice had reached a similar height (fig. 

6N) and, by April, an area about 100 by 50 m had attained a 

height of more than 120 m above the 2003 surface (fig. 6P ). 

Between late February and April, new crevasses oblique and 

perpendicular to the existing crevasses cut the southern part 

of the glacier into blocks (fig. 6O). A prominent east-west, 

arcuate, steplike feature (figs. 6I–P; highlighted by arrows 

on fig. 6I) about 30 m high formed at the glacier surface east 

of the 1980s lava dome (see discussion in Walder and others, 

this volume, chap. 13). This ice step was cut by short radial 

crevasses. The broad thickening of west Crater Glacier continued primarily along the southwest edge of the 1980s lava 

dome. For example, the 20-m isoline lay about 130 m farther 

northwest in April than in February (figs. 6J, 6P).
158 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Shoestring

notch

Shoestring

notch

Figure 6—Continued.

June 15 Through August 10, 2005

The third group of image pairs (fig. 7A–F) shows that 

dome growth began to migrate west from its previous locus 

in the southeastern part of the crater. Though it is obscured 

in the June image (fig. 7A), the July image (fig. 7C) shows 

spine 5 was shorter (145 m in length), narrower (100 m 

across at the base), and higher (altitude 2,352 m, 250 m 

above the 2003 glacier surface) than the previous two 

spines, and it sloped steeply (about 60°) in all directions 

(see Vallance and others, this volume, chap. 9). Talus filled 

the substantial trough between west Crater Glacier and the 

dome that had existed for more than six months, eventually spilling onto the glacier surface. The slope angle of 

the talus on the northwest part of the new dome was about 

32°. In July, spine 5 attained a height of 260 m (altitude 

2,368 m) above the 2003 surface—the highest measured 

so far and within 2 m of the altitude of Shoestring notch, 

the lowest point on the crater rim. In August the dome was 

about 700 m in length and 600 m in width, with a volume 

of 62×106

 m3. Throughout the summer, crevasses continued 

to increase in number, disrupting the east Crater Glacier 

surface south of the prominent step, and they appeared in 

the elongate terminus of the glacier for the first time (fig. 
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 159

Figure 6—Continued.

7A–F). The effect of the northward advance of east Crater 

Glacier’s terminus was to create an island of increased 

elevation shown by a series of closed isolines (for example, 

fig. 7B) (Walder and others, this volume, chap. 13). The 

isolines on west Crater Glacier indicate that thickening had 

slowed. However, subtracting the June 14 glacier surface 

from the July 15 glacier surface yields a negative change 

in volume (–1.5×106 m3), likely a result of glacier melting. 

The west Crater Glacier-dome contact migrated about 60 m 

west between March and August, and radial and circumferential crevasses formed along the southeastern edge of the 

west glacier.

September 20 Through December 15, 2005

The fourth group of image pairs, (fig. 8A–F) shows 

that the locus of new dome growth migrated about 200 m 

west between August 10 and September 20, 2005. Part of 

spine 5 had decreased in elevation by about 75 m, forming 

a depression. Between September and December, spines 

6 and 7 continued to grow westward, and the depression 

became a well-defined trough separating the newest growth 

from earlier spines. Spine 6 is difficult to distinguish within 

its flanking talus (fig. 8A), but it extruded on the northeast 

end of an elongate, arcuate ridge trending southwest along 
160 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 7. Sequence of three pairs of shaded-relief images (A–F ) showing growth of lava dome complex (red outline) 

and deformation of Crater Glacier from June to August 2005. For full explanation, see figure 3 and text.
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 161

the top of the talus. In September, this newest part of the dome 

was about 420 m long, 285 m wide, and about 170 m (2,274 m 

altitude) above the 2003 glacier surface. By October, the highest 

point of spine 6 had migrated about 85 m west. Between October and December, spine 7 penetrated the eastern side of spine 6 

and extruded upward and westward, bounded by a well-defined, 

nearly planar, gouge-covered surface striking approximately 

north-south. The nearly circular mass of talus and spines 6 and 7 

was 430 m in length, 425 m in width, and about 180 m (altitude 2,280 m) above the 2003 surface. By December 2005, the 

entire dome complex was about 900 m long, 625 m wide, and 

190 m (altitude 2,316 m) above the 2003 surface. Its volume, 

73×106 m3, was nearly equal to the 74×106 m3 volume estimate 

of erupted lava for the 1980s lava dome (Swanson and Holcomb, 1990). East Crater Glacier was severely disrupted by both 

crevasses and summer ablation. The west glacier-dome contact 

migrated about 100 m west between September and December. Large crevasses formed on the west glacier, most likely in 

response to westward dome growth (Walder and others, this volume, chap. 13). Radial and circumferential crevasses increased 

in both size and number between September and December 

until the part of the glacier adjacent to the new dome took on 

the appearance of the disrupted east glacier. A prominent linear 

step, similar to the one on the east glacier, had migrated about 

100 m north, and a broad area immediately south of the step had 

risen about 40 m. 

Discussion

DEMs constructed from vertical aerial photographs 

provided an effective means to track and quantify dome growth 

and to calculate average magma-discharge rates of the 2004–5 

lava dome at Mount St. Helens, as well as to gauge deformation 

of Crater Glacier. The DEMs met commonly identified needs 

in being (1) accurate—satisfying the appeal for precise, highquality measurements at growing domes (Newhall and Melson, 

1983); (2) safe—“acquired at low risk, and consequently * * * 

an important element in monitoring future activity of potentially 

explosive volcanoes” (Moore and Albee, 1981, p. 127); and 

(3) relatively inexpensive—acquisition costs of vertical aerial 

photographs and high-resolution scans are about 5 to 10 percent 

of those of other technologies such as lidar. Below we highlight 

suggestions for use of this technique at future eruptions, point 

out how DEMs and extruded-volume and extrusion-rate data 

were used in other studies, and draw a few brief comparisons 

between this and other dome-building eruptions.

Importance of Preparation and Suggestions for 

Improvements

Accurate measurements of the dome and glacier at Mount 

St. Helens were possible during even the critical earliest part 

of the 2004–5 eruption because we had baseline data and photogrammetric infrastructure in place before the eruption began. 

An earlier study of Crater Glacier (Schilling and others, 2004) 

had (1) identified an experienced photogrammetric contractor 

that ensured accuracy in camera calibration, film processing 

and scanning, and critical flight parameters such as altitude; 

(2) developed skills using aerotriangulation to orient blocks 

of photogrammetric stereomodels and to collect breaklines 

and points to create DEM surfaces with minimal operator bias 

and error; and (3) most importantly, established a network of 

ground control points on the flanks and crater floor of Mount 

St. Helens, validated by repeated measurements, with locational accuracy of a few centimeters. Hazards in the crater and 

uncertainty about the course of the eruption during late September and early October 2004 would have limited our ability 

to establish ground control. In retrospect, a greater number of 

control points in the crater area would have been useful, both 

to replace sites that were destroyed by explosions early in the 

eruption and to establish a set of points that could be remeasured to evaluate the accuracy of successive DEMs.

We spent days to weeks constructing each detailed DEM 

presented in this report, but, especially early in the eruption, a 

two-stage approach would have been advantageous to balance the need for rapid measurements for hazard-assessment 

purposes versus a greater level of detail for in-depth studies. 

Initial work would capture the minimal detail needed to obtain 

preliminary volume and extrusion-rate measurements. Later, as 

time allowed, we could add greater detail and improve accuracy. 

Average Extrusion Rates

Plots of volume through time, whether comparing new 

crater-surface topography to the relatively uniform 2003 

glacier surface (total volume change) or lava-dome volume 

as defined by the vertical projection of the dome outline to 

the 1986 crater floor (extruded volume), show that volume 

increased quickly early in the eruption and more slowly thereafter (fig. 9; table 1). Early in the eruption, measurements of 

total volume change document the deformation of glacier ice 

and existing crater floor to form a prominent welt (Dzurisin 

and others, 2005) having a volume of 10×106

 m3 and a growth 

rate of 8.9 m3/s before lava first appeared at the surface on 

October 11, 2004. This volume estimate may be a minimum, 

because steam obscured some of the vent area. After the 

appearance of dacite lava at the surface, lava extrusion rates 

were initially 5.9 m3

/s, slowing to 2.5 m3/s by the beginning of 

January 2005. After early 2005, the extrusion rate gradually 

declined to about 0.7 m3/s, with a minor increase in rate during 

late summer 2005, and the total extruded volume gradually 

increased to 73×106 m3 near the end of 2005.

For much of the eruption, the extruded-lava volume 

derived from the DEMs was the most appropriate measure of 

dome growth. However, especially early in the eruption, the 

unique situation of a lava dome erupting through a glacier 

required a second method, which included a measure of the 

total volume of surface deformation. This value was the sum 

of extruded volume and other surface volume change owing 
162 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 8. Sequence of three pairs of shaded-relief images (A –F ) showing growth of lava dome complex (red outline) 

and deformation of Crater Glacier from September to December 2005. For full explanation, see figure 3 and text.
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 163

Total volume-change rate

Lava extrusion rate

9

8

7

6

5

4

3

2

1

0

RATE, IN

CUBIC METERS PER SECOND

October

December

June

February

April

August

October

December

2004 2005

B

RATE

Total volume change

VOLUME, IN MILLION

CUBIC METERS

90

80

70

60

50

40

30

20

10

0

A

VOLUME

Extruded-lava volume

Figure 9. Time-series plots of dome growth at Mount St. Helens from start of eruption through December 

2005. A, Total volume change (blue line) and extruded-lava volume (orange line) versus time. B, Total volume 

change rate (blue line) and volumetric extrusion rate (orange line) versus time. Early in eruption, all surface 

deformation likely resulted from extrusion or shallow intrusion of lava into glacier, and total volume change 

therefore would approximate lava extrusion rate.

to deformation, snow and ice removal or accumulation, and 

addition of new talus from the crater walls. As the eruption 

continued, we used the extruded-lava volume to track dome 

growth and used the total volume change mainly to track 

changes in the glacier. Differences between the two rates 

resulted from factors unrelated to lava extrusion. For example, 

a prominent increase in the volume change rate in April 2005 

(fig. 9B) coincided with including a larger part of west Crater 

Glacier in the April DEM, which resulted in adding more 

thickened glacier and seasonal snowfall to the total volume. A 

notable decrease of total volume change during a period of low 

extrusion rate during the summer of July 2005 likely reflects 

summer snowmelt and glacier ablation. The apparent slight 

increase in the extruded-lava rate in August 2005 is likely real. 

Use of DEMs and Volume and Extrusion-Rate 

Data in Studies of the 2004–2005 Eruption

Vertical aerial photographs and the DEMs derived from 

them constitute one of the most widely used data sets collected 

during the 2004–5 eruption of Mount St. Helens. Scanned 

versions of the vertical aerial photographs provided a consistent base for interpretation and construction of photogeologic 

maps (Herriott and others, this volume, chap. 10). The DEMs 

provided numerical constraints to solve equations needed to 

estimate linear extrusion rates from a single remote camera 

(Major and others, this volume, chap. 12). Coupling the aerial 

photographic documentation with fields of deformation vectors and profiles of the dome derived from the DEMs yielded 

evidence to interpret the mechanics of dome growth (Vallance 

and others, this volume, chap. 9). Measurements and profiles 

from the DEMs helped to track the remarkable deformation of 

east and west Crater Glacier and showed that relatively little 

ice has been melted by dome growth (Walder and others, this 

volume, chap. 13).

Measurements of dome volume and extrusion rate 

were combined with other data sets such as seismology, gas 

geochemistry, and ground deformation to better understand 

eruptive processes. As the eruption continued, the extrusion 

rate remained relatively constant, even though the character of 

seismicity varied, suggesting that earthquakes were controlled 

more by changes in extrusion mechanics than by changes in 

extrusion rate (Moran and others, this volume, chap. 2). The 
164 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

nearly constant extrusion rate was also a key observation 

incorporated into a mechanistic model of extrusion driven by 

a nearly constant influx of magma from depth and resisted by 

a plug of solidified magma that slipped incrementally against 

the walls of the conduit (Iverson and others, 2006; Iverson, 

this volume, chap. 21). Models from GPS (Lisowski and others, this volume, chap. 15) and InSAR data (Poland and Lu, 

this volume, chap. 18) yield estimates of reservoir volume 

loss of 15–30×106 m3, compared with a volume of 70–80×106

m3 for the lava dome. This disparity most likely reflects a 

combination of expansion of magmatic volatiles and recharge 

(Mastin and others, this volume, chap. 22). However, erupted 

volumes and total volatile output for the 2004–5 eruption 

were used to estimate the volatile content of magma at 8.6-km 

depth. Gerlach and others (this volume, chap. 26) conclude 

that the magma was nearly depleted in excess volatiles, suggesting that new gas-rich magma has not been added into the 

reservoir during the months just before or during the eruption.

Volume measurements of the growing lava dome and 

derived extrusion rates help to place the 2004–5 eruption in 

context with those of other dome-building volcanoes (fig. 10) 

and within the growth history of Mount St. Helens since 1980. 

The average rate of growth of the current dome, about 2 m3

/s 

through the end of 2005, is an order of magnitude above the 

long-term (1980–present) eruption rate of Mount St. Helens 

(Iverson, this volume, chap. 21). However, when compared to 

other dome-growth episodes at Mount St. Helens and elsewhere the current growth rate is fairly typical (fig. 10). VOLUME, IN MILLION CUBIC METERS

MONTHS

Montserrat, 1995−97

Montagne Pelee

1902−5

Mount St. Helens cryptodome Unzen, 1990−95 (March 20−May 18, 1980)

Mount St. Helens

(1981−86)

Mount St. Helens 2004−5

0

50

150

200

350

0 20 30 40 50 60

100

10

Redoubt, 1989−90

St. Vincent, 1979

Figure 10. Mount St. Helens 2004–2005 growth rate curve compared with other historical lava 

domes (modified from Nakada and others, 1999; Tanguy, 2004).

Conclusions

Vertical aerial photographs, taken at time intervals ranging from successive days to every few weeks, and the digital 

elevation models (DEMs) constructed from them have been 

critical tools to document the remarkable growth history of the 

2004–5 lava dome at Mount St. Helens, especially in providing estimates of volume and volumetric extrusion rate. Single 

DEMs provided length and width measurements of the dome 

and glacier; sequential DEMs recorded temporal and spatial 

changes. Moreover, measurements of volume and volumetric 

extrusion rate proved to be critical parameters. When combined with other primary data sets, they helped to form models 

and hypotheses that illuminated the mechanics and dynamics 

of the eruption, and thus helped to assess volcano hazards. 

The successful application of aerophotogrammetry to monitor 

this eruption was possible largely because baseline data and 

photogrammetric infrastructure were in place before the onset 

of activity. The aerial photos and DEMs obtained in this study 

are an enduring resource for addressing basic questions about 

the eruption, including many raised in other papers in this collection and, we suspect, others yet to be asked.

Acknowledgments

We thank Richard Iverson, Richard Herd, and Dan 

Dzurisin for their insightful and helpful reviews. Dan Dzuri-
8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 165

sin also contributed insightful and helpful discussions. Bergman Photographic Services, including Larry and Bruce Bergman, Travis Marshall, Trevor Gray, Steve Tank, Kyla Dorsey, 

and John Hogl, have all contributed their expertise to ensure 

that we acquired excellent aerial photographs and scans. We 

are indebted to Jeff Linscott and the folks at JL Aviation for 

outstanding helicopter work, especially landing us gently 

in difficult places on the mountain for ground control work. 

Frank Trusdell, Mike Poland, Julie Griswold, Chris Harpel, 

Dave Ramsey, Joel Robinson, Stephanie Konfal, Angie Diefenbach, Sarah Thompson, and Susanne Ettinger helped with 

ground control work and offered insightful discussions. Last, 

but perhaps most important, we thank Bobbie Myers, who 

watched over us remotely, tracked our safety, and allowed us 

to focus on field tasks.

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8. Use of Digital Aerophotogrammetry to Determine Rates of Lava Dome Growth, Mount St. Helens 167

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1750 (this volume).

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doi:10.1016/0377-0273(90)90121-U.

Appendix 1. Hillslope Shaded-Relief Maps

[This appendix appears only in the digital versions of this work—in the DVD that accompanies the printed volume and as a 

separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750.] 

The appendix contains 17 raster images in Tagged Image File Format (filename extension is “tif”), one for each of the 

hillshade relief maps used in figures 4, 6, 7, and 8. The filename contains the date (year-month-day) of the map. Corresponding ASCII world files (filename extension “tfw”) for georeferencing the raster maps by GIS software are included, each using 

the Universal Transverse Mercator projection, zone 10, North American Datum 1983. A separate metadata file summarizes the 

pertinent details of image processing.
168 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 9

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, Alaska Volcano Observatory, 4200 University 

Drive, Anchorage, AK 99508

Growth of the 2004–2006 Lava-Dome Complex at Mount 

St. Helens, Washington

By James W. Vallance1, David J. Schneider2, and Steve P. Schilling1

Abstract

The eruption of Mount St. Helens from 2004 to 2006 

has comprised extrusion of solid lava spines whose growth 

patterns were shaped by a large space south of the 1980–86 

dome that was occupied by the unique combination of glacial 

ice, concealed subglacial slopes, the crater walls, and relics 

of previous spines. The eruption beginning September 2004 

can be divided (as of April 2006) into five phases: (1) predome deformation and phreatic activity, (2) initial extrusion 

of spines, (3) recumbent spine growth and repeated breakup, 

(4) southward extrusion across previous dome debris, and (5) 

normal faulting of the phase 4 dome to form a depression, a 

shift to westward extrusion and overthrusting of earlier phase 

5 products. Overall, steady spine extrusion gradually slowed 

from 6 m3

/s in November 2004 to 0.6 m3/s in February 2006.

Thermal camera data show that phase 1 activity included 

low-temperature thermal features, such as fumaroles, fractures, and ground warming related to rapid uplift, as well as 

deformation in the south moat of the crater. The relatively cold 

(<160°C) phreatic eruptions of early October heralded activity 

at a subglacial vent situated along the south-sloping margin of 

the 1980–86 dome. Thermal infrared imagery, documenting 

increased heat flow, presaged phase 2 extrusion of the October 

11–15, 2004, lava spine. The thermal images of the extruding 

spine revealed a hot basal margin and highest temperatures of 

600–730°C. 

During phase 3, a recumbent whaleback-shaped spine 

with a low-temperature shroud of fault gouge and a hot, 

U-shaped basal margin extruded. This spine pushed southward 

along the bed of the glacier until it encountered the south wall 

of the 1980 crater, whereupon it broke up, decoupled, and 

regrew. Continued southward growth of the recumbent spine 

pushed cold deformed rock, hot dome rubble, and glacier 

ice eastward at a rate of 2 m/d. In April 2005, breakup of the 

whaleback and growth of a lava spine across previous dome 

rubble heralded phase 4 spine thrusting over previous spine 

remnants. During phase 4, the active spine pushed southward with an increasingly vertical component and increasing 

incidence of large rockfalls. In late July, the spine decoupled 

from its source, the vent reorganized, and a new spine began 

to grow westward at right angles to the previous growth direction, defining phase 5. Dome migration again plowed glacier 

ice out of the way at a rate of about 2 m/d, this time westward. In early October, the spine buckled near the vent and 

thrust over the previous one. A massive spine monolith had 

been constructed by December 2005, and growth of spines 

with increasingly steep slopes characterized activity through 

April 2006.

The chief near-surface controls on spine extrusion during 

2004–6 have been vent location, relict topographic surfaces 

from the 1980s, and spine remnants emplaced previously 

during the present eruption. In contrast, glacier ice has had 

minimal influence on spine growth. Ice as thick as 150 m has 

prevented formation of marginal angle-of-repose talus fans 

but has not provided sufficient resistance to stop spine growth 

or slow it appreciably. Spines initially emerged along a relict 

south-facing slope as steep as 40° on the 1980s dome. The 

open space of the moat between that dome and the crater walls 

permitted initial southward migration of recumbent spines. 

An initial spine impinged on the opposing slopes of the crater 

and stopped; in contrast, recumbent whaleback spines of phase 

3 impinged on opposing walls of the crater at oblique angles 

and rotated eastward before breaking up. Once spine remnants 

occupied all available open space to the south, spines thrust 

over previous remnants. Finally, with south and east portions of the moat filled, spine growth proceeded westward. 

Although Crater Glacier had only a small influence on the 

growing spines, spine growth affected the glacier dramatically, 

initially dividing it into two arms and then bulldozing it hundreds of meters, first east and then west, and heaping it more 

than 100 m higher than its original altitude.
170 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Introduction

Continuous, steady extrusion of gas-poor, solidified dacitic magma through glacier ice has characterized the 2004–6 

eruption of Mount St. Helens. Dome emplacement has been 

influenced by the geometry of the 1980 crater, an amphitheater 

~2 km across, 500 m deep, and open to the north (fig. 1). In 

the exact middle of the amphitheater, the 1980–86 dome grew 

to a volume of about 77×106 m3, attaining dimensions of 860 

by 1,060 m in plan and reaching a height of 270 m above the 

flat-floored amphitheater (fig. 1A) (Swanson and Holcomb, 

1990). By fall 2004, the north-facing aspect and steep walls, 

prolific annual snowfall, and frequent winter avalanches in the 

1980 crater had given rise to a rapidly growing glacier, as thick 

as 150 m and with a volume of about 80×106 m3, that wrapped 

the 1980s dome like a U (fig. 1B) (Schilling and others, 2004).

During initial volcanic unrest between September 23 and 

October 10, 2004, uplift and deformation along the southern 

part of the 1980s dome and glaciated areas to the south formed 

a welt more than 100 m high. Deformation of the 1980–86 

dome, crater-floor debris, and glacier ice south of the 1980–86 

dome has continued throughout the eruption. However, during 

subsequent dome growth, the locus of deformation has shifted 

southward, then alternately eastward and westward, as actively 

growing spines plowed old rocks, recently emplaced but 

inactive spines, and glacier ice out of their way (Dzurisin and 

others, 2005).

Following unrest that began on September 23, 2004, and 

culminated with phreatic eruptions 8–12 days later, intrusion 

and extrusion of solid magma has typified the eruption. The 

magma is unusually gas poor (Gerlach and others, this volume, 

chap. 26) and crystal rich (Pallister and others, this volume, 

chap. 30). Several meters of pulverized, variably sintered rock 

(Cashman and others, this volume, chap. 19) has commonly 

coated emergent lava spines, lending them a smooth appearance. Other spines have broken apart to become surrounded by 

hot talus fans.

Terminology used in this paper is as follows. A single 

lava dome was extruded at Mount St, Helens from 1980 to 

1986 and a second from October 11, 2004, through the time 

of this writing. Because of their solid-state character, individual extrusions of the current dome-building eruption are 

termed “spines,” not lobes. The term “recumbent” implies 

that the horizontal component of extrusion is greater than the 

vertical component. The term “whaleback” describes a form 

of smooth-surfaced recumbent spine. Several spines include 

upthrusted deformed rock of previous crumbled spines and 

older rock from the 1980–86 dome. Such spines are termed 

“compound” following Blake (1990) and distinguished from 

individual growing spines, which are typically monolithic 

rather than rubbly.

Spine morphologies of the current eruption are variants 

of Blake’s (1990) upheaved plugs and peléean domes but do 

not include the more fluid, low lava domes and coulees. As 

of April 2006, the growing dome has included seven spines 

1980–86

dome

Mount

St Helens

1980

 Crater 

5117000

5115500

5116000

5116500

561500 562000 562500 563000 563500

A

C

B ́

B

A ́

D

C ́

D ́

E E ́

F

F ́

Mount

St Helens

Crater

Glacier 

5117000

5115500

5116000

5116500

561500 562000 562500 563000 563500

1980–86

dome

B

A

Opus,

1985

Opus

1985

1985

fault

0 200 400 600 800 1,000

METERS

0 200 400 600 800 1,000

METERS

Figure 1. Digital elevation models (DEMs) of Mount St. Helens 

showing 1980 crater, 1980–86 dome, Opus, and Crater Glacier. 

Coordinate system is North American Datum 1927 Universal 

Transverse Mercator, zone 10N, in this and subsequent map 

figures. A, 1986 DEM. B, 2003 DEM. Solid rectangle locates DEMs 

in subsequent figures, and dashed lines locate cross sections 

illustrated in subsequent figures.

(figs. 2, 3) but no surface flowage features or extruded silicic 

lava flows. Spine shapes have included steeply inclined fins, 

broken blocky forms, and whalebacks (fig. 2). Four of the 

spines have grown recumbently, and five of them have pushed 

through thick glacial ice.

Thermal infrared imagery, petrography, and seismology all suggest that extrusion of these solid spines occurred 

at temperatures below the rock’s solidus temperature. Rock 

samples that are porphyritic, microlite rich, and glass poor are 

consistent with subsolidus eruption (Pallister and others, this 

volume, chap. 30). Shallow seismic signals that locate within 

about 1 km of the surface and an absence of deeper signals 

(Moran and others, this volume, chap. 2; Thelen and others, this volume, chap. 4) suggest a possible viscous-to-solid 

transition of the magma at that depth. The hottest recorded 

temperatures of the extruding spines, culled from >10,000 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 171

images collected during 37 missions, cluster where cracks 

and avalanches expose fresh interior surfaces and fall between 

700°C and 730°C. These measurements provide a minimum 

limit for the temperature of extruded spines that is much lower 

than the solidus temperature of 920–960°C for the dacite of 

spines 1–7 (Pallister and others, this volume, chap. 30).

The purpose of this report is to document the characteristics of spine growth at Mount St. Helens, the dramatic 

near-field deformation that accompanied it, and the impact of 

dome growth on Crater Glacier. To achieve this we examined 

and analyzed oblique and vertical aerial photography, digital 

elevation models (DEMs), and thermal infrared imagery. In 

particular, we used DEMs and aerial photographs to document 

growth of spines and nearby deformation as a function of time. 

To illustrate how the dome has grown and how that growth has 

affected surrounding areas, we generated surface deformation 

maps and interpretive cross sections. As supporting evidence 

we considered results reported elsewhere in this volume, such 

as geologic mapping (Herriott and others, chap. 10), GPS 

instrumentation (LaHusen and others, chap. 16), and repeat 

photography from fixed sites (Major and others, chap. 12; 

Dzursin and others, chap. 14; Poland and others, chap. 11).

Deformation Within the Crater

Methods and Assumptions

During a period of about 18 months, repeated visual 

observations, oblique aerial photography, thermal infrared 

imagery, lidar, and high-resolution aerial photography delineated evolution of the 2004–6 dome at Mount St. Helens and 

deformation of nearby features in response to that growth. 

Frequent aerial reconnaissance allowed observations and 

oblique aerial photography as weather permitted. Cascades 

Volcano Observatory (CVO) staff collected such data almost 

daily in the period from September 27, 2004, until October 15, 

2004. Thereafter, observations were less frequent, with repeat 

intervals increasing from a few days to as long as eight weeks.

Thermal Infrared Imagery

Thermal infrared (TIR) images allowed estimation of 

pixel-integrated temperatures for exposed dome-rock surfaces, 

fumaroles, and other features. More generally TIR surveys 

showed how surface areas were heated before the appearance 

of spines at the surface, allowed differentiation of individual 

spines, showed thermal structures within spines, and revealed 

how spines evolved and cooled once extruded (fig. 3). We conducted 37 TIR surveys of the deformed area and the growing 

dome between October 1, 2004, and April 30, 2006.

The instrument used, a FLIR Systems ThermaCAMTM 

PM595 infrared camera, mounted on a helicopter, is a 

microbolometer that measures brightness in the 7.5–13 μm 

waveband to detect temperatures in the range from −40°C to 

1,500°C. It collects TIR images as frequently as once per second and can acquire both TIR and standard video (Schneider 

and others, this volume, chap. 17). Conversions to temperature 

depend on emissivity, atmospheric temperature, humidity, 

distance, viewing angle, steam, and gas (Ball and Pinkerton, 

2006; Harris and others, 2005). We can independently measure 

atmospheric temperature, humidity, and distance well enough 

that resultant errors are about ±10°C; if emissivity is 0.96±0.1, 

additional errors would be ±5 percent (Schneider and others, 

this volume, chap. 17). We can only minimize errors owing to 

the other parameters by repeating measurements at multiple 

viewing angles and reporting temperature values for conditions with minimal gas and steam. Images from a TIR survey 

at a distance of about 1 km yield a horizontal field of view of 

about 210 m and a pixel resolution of about 1.5 m. Integration of brightness within individual pixels means that hottest 

reported temperatures could be averaged across areas less than 

2 m2 (Schneider and others, this volume, chap. 17).

Repeat Aerial Photographs, Lidar, and DEMs

A sequence of aerial photographs and lidar converted to 

DEMs provided vertical and planimetric control at intervals 

of 1 to 55 days during the 18-month study period. Lidar data 

from November 2003 (Queija and others, 2005) provided 

initial datum control, and DEMs generated from topographic 

maps provided control for the 1980 and 1986 surfaces (fig. 

1). Three DEMs were derived from lidar surveys made early 

in the eruption between October 4 and November 20, 2004, 

(U.S. Geological Survey and National Aeronautics and Space 

Administration, unpub. data). In addition, Schilling and others 

(this volume, chap. 8) created 18 DEMs from vertical aerial 

photography taken between October 4, 2004, and February 9, 

2006. The DEMs of October 4 and 13, 2004, (Schilling and 

others, this volume, chap. 8) provided a check of lidar DEMs 

collected October 4 and 14.

Identification and Tracking of Features

During the study period, we used aerial photographs and 

DEMs to identify and track primary and secondary features. 

Primary features could be located three dimensionally in two 

or more DEMs and included points at distinctive topographic 

crests or, less commonly, troughs and intersections of linear 

features. Examples of point features include distinctive spine 

formations (for example, fig. 2, point c), megablocks, stranded 

ice blocks, the toes of avalanches from the 1980 crater walls 

(fig. 2, points a and i), and seracs. Intersection features include 

crack networks on the growing dome (fig. 2, point h) and 

crevasses on Crater Glacier. In many cases, primary features 

formed of ice and snow persisted only from October 2004 

through March 2005. Secondary features are those that we 

could track approximately in plan view but for which vertical 

control was difficult or impossible to obtain (fig. 2, features 

d, e, f, g). Secondary features included margins of actively 
172 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

a

b

12/11/04 d

f

e

11/20/04

a

c

11/29/04

a

b

c

i

i

C 

B 

A 

Opus

1980–86

 dome

Opus

Welt

Welt

Welt

Spine 3

145x440 m

1980–86

 dome

 East

 Crater

 Glacier

West 

Crater

 

Glacier

West 

 Crater

 Glacier

West 

Crater

 

Glacier

cw

Spine 3

125x420 m

Spine 3

120x460 m

East 

Crater

 

 

 Glacier

k

j

 East

 

Crater

Glacier

s1

s2

s1

cw

cw

j

s1

Welt

m

m

m

Opus

a, i - avalanche toes

b - avalanche margin

c - cracks on spine 3 (s3)

d, e, f, g - s3 fragments

j - 11/20–29 s3 fragment

k - 12/4–8 s3 fragment

m - block on glacier

s1, s2 - spines 1 and 2

cw - 1980 crater wall

Figure 2. Photographs of Mount St. Helens 

crater and whaleback-form spines 3 and 4 

taken looking south-southwest and illustrating 

primary and secondary features as they 

evolved during spine growth and deformation. 

Features denoted with squares are primary 

“point” features, and others are secondary 

features. A, November 20, 2004; USGS photo 

by J.N. Marso. B, November 29, 2004; USGS 

photo by M. Logan. C, December 11, 2004; 

USGS photo by J.S. Pallister. D, December 28, 

2004; USGS photo by S. Konfal. E, January 3, 

2005; USGS photo by M. Logan. F, January 14, 

2005; USGS photo by M. Logan.
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 173

a

b

d

e

f

g

1/14/05

h

a

b

d

f

e

g

1/3/05

E 

F 

Opus

1980–86

 dome

Welt

East

 Crater

 Glacier

Spine 4

Opus

1980–86

 dome

Spine 4

130x210 m

a - toe of avalanche

b - margin of avalanche

d, e, f, g - s3 fragments

h - cracks on spine 4

j - 11/20–11/29 s3 fragment

cw - 1980 crater wall

weltgl - welt ice lost to glacier 

cw

cw

East

 Crater

 Glacier

Welt weltgl

weltgl

a

b

d

f

e

g

 12/28/04

D 

Opus

Welt

East 

 Crater

 

 Glacier

 Transition

spine 3–spine 4

cw

j

weltgl

Figure 2—Continued.
4 21 3 3 4 eg 1/14/05 3

5

4

4

4

2

2

2

1

3

3

3

3

eg

4

4

4

5

2

3

wg

eg

eg

wg

7/13/05

3/15/05

2/16/05

1

1980s dome

4

4

5

5 5

6 4

5

3 4 6

wg

wg

3

eg

4

eg

5

5

7

6

wg

4

5 7

6

5

wg

eg

1980s dome

Opus

Opus

Opus

Welt

Welt

1980s dome

1980s dome

1980s dome

1980s dome

3

Opus

Opus

Opus

3

1

1

1

~300 m

Opus

~200 m

~200 m

~200 m

~200 m

~270 m

~300 m

~310 m

Figure 3. Photograph and thermal-infrared image pairs from Mount St. Helens, views from north-northeast (Jan. 14–Sep. 21, 2005), north (Oct. 1northwest (Nov. 18, 2005), illustrating evolution of the 2004–6 dome. Numbers designate spines 1–7; eg and wg designate east and west Crater Gin photograph of Aug. 10, 2005, is trace of fault developing between spines 5 and 6. USGS photos of Jan. 14, March 15, July 11, and Nov. 18, 2005, those of Feb. 16 and Sep. 21, 2005, by J.W. Ewert; those of Aug. 10 and Oct. 11, 2005, by C.A. Gardner and C. Fox-Lent. Thermal images of Feb. 16 aJ.W. Vallance; the remainder by M. Logan.
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 175

growing spines, the contact of the 1980–86 dome with the new 

dome, the contacts of the glacier with the new dome and crater 

wall, and glacier snouts. Such secondary features provided 

useful constraints on deformation but could not be used to 

measure that deformation directly.

Sources of error in locating primary features included 

accuracy of DEMs, accuracy of repeating a location in a single 

DEM, identification and location of features in consecutive 

DEMs, deformation of features with time, melting or addition of snow to features, and misidentification of features in 

successive DEMs. Precision of DEMs is a few centimeters to 

about a decimeter (Schilling and others, this volume, chap. 8). 

On individual DEMs, distinctive features could be relocated to 

within ±2 m horizontally and ±1 m vertically.

Features were identified and relocated on successive 

DEMs with some certainty unless the features were obscured 

by snow, shadows, steam, or clouds or were so deformed during the interval between DEMs as to become difficult to recognize. Fresh snow and ablation may have affected relocation 

and ultimately even recognition of primary features, but some 

areas on or near the active dome were windswept or remained 

warm year-round; features in such areas were not subject to 

relocation errors related to melting or snowfall. Comparison 

of relative motion of groups of neighboring objects provided a 

check on correlations. We eliminated correlations that yielded 

results greatly at variance with those of neighboring objects.

Generally, error in locating primary features in successive 

DEMs is between ±2 m and ±5 m in the horizontal and vertical 

dimensions. Features subject to the most deformation—those 

on or near the actively growing dome—tend to be those with 

the greatest errors in relocation. Features away from the locus 

of deformation may have small absolute location errors but 

may not move far enough to register significant motion, given 

the magnitude of error in locating them.

Generation of Surface-Deformation Vectors

Primary features that can be located in two or more successive DEMs allow estimation of vector components from 

one time to another. This allows calculation of average rates 

of deformation (for example, points a, b, c, and h in fig. 2). As 

we know of no systematic source of error in locating primary 

features, errors should not tend to accumulate for features that 

can be located in three or more successive DEMs. Repeated 

locations of primary features indicated in figures by solid 

arrows thus allow generation of surface-deformation vectors 

within our stated error limits.

Successive locations of secondary features give a sense of 

magnitude and direction of deformation but do not yield true 

vectors. In such cases, the vertical component of deformation 

may be poorly known or unknown. Despite poor constraints 

on vertical position, plan-view locations of some secondary 

features are as accurate as those of primary features. Examples 

include features located in rectified aerial photographs for 

which no DEM exists and features that crumble as they move 

laterally yet still can be identified. Crumbling features are common on or near active spines. Relocations of such features can 

be accurate in plan; but vertical changes, if given, are minimum 

values. Relocations of secondary features such as the contact 

between the glacier and active dome give minimum constraints 

on deformation in both horizontal and vertical directions.

Vector Fields

Vector fields were derived from simultaneous tracking of 

numerous primary features on successive DEMs. Vector fields 

were used to delineate growth of active spines, deformation 

of inactive parts of the dome and its surroundings, and deformation of the glacier in locations where motion exceeded a 

threshold of about 4 m in the time between successive DEMs. 

Because DEMs were produced at intervals of 9–55 days and 

vertical and horizontal precision were ~5 m, detection limits for 

time-averaged deformation rates range from 0.6 to 0.09 m/d.

Comparison with GPS Data

During certain intervals, portable GPS receivers that 

provided nearly continuous measurement of deformation 

(LaHusen and others, this volume, chap. 16) were located 

near features tracked during this study using DEMs, thereby 

providing a check on our results. The GPS deformation 

measurements compared well with those of this study. For 

example, a GPS receiver placed on the spine during November 21–29, 2004, gave a vector almost identical in magnitude 

and direction (10.3 m/d, S. 19° E., up 6°) to that of a nearby 

feature 30 m east that we tracked November 20–29 (10.4 m/d, 

S. 21° E., up 8°).

Our deformation measurements have advantages and 

disadvantages compared with those derived from GPS receivers. The chief advantage of our approach is that we can track 

numerous features simultaneously and thus obtain a complete 

picture of dome growth and nearby deformation patterns. 

GPS sensors are advantageous in that their data streams can 

be sampled frequently and transmitted back to the observatory. Such real-time acquisition permits the use of GPS data in 

monitoring. In contrast, our measurements are values averaged 

over intervals between successive DEMs and have no application in real-time monitoring because of the additional time 

required to prepare DEMs.

Volume and Flux Calculations

All reported volumes assume the 1986 topographic 

surface as a datum and subtract it from DEMs of various dates 

over pertinent areas (hot-rock volumes given in Schilling and 

others, this volume, chap. 8). Because thick glacial ice overlay 

the 1986 debris fill of the moat by 2004, we chose the 1986 

surface as a datum for volume calculations rather than the 

more recent ice-mantled surface of 2003–4. To facilitate calculations, we assumed that bounding surfaces between the datum 

and the areal extent of hot rock on any subsequent DEM were 
176 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

vertical. The assumption of vertical bounding surfaces dictates 

that calculated volumes are minimum values in cases where 

natural surfaces differ substantially from vertical. A steep, 

near-vertical contact between the extruding dome and glacier 

ice is probably a reasonable assumption for two reasons. First, 

visible upper parts of glacier contacts with hot rock were 

steep and nearly vertical. Second, relict ice-hot rock contacts 

exposed in other volcanic areas are commonly nearly vertical 

because the ice cools and buttresses the rock margin, preventing avalanches and rockfall that tend to form slopes more 

closely approaching the angle of repose.

Time-averaged volumetric extrusion rates for individual 

spines are derived by comparing volumes from one DEM 

to the next and dividing by the time between them. In many 

cases, volumetric rates are the same as the hot-rock extrusion 

rates of Schilling and others (this volume, chap. 8). Well-constrained growth intervals of certain spines allow more precise 

calculation of their volumetric extrusion rates.

Cross Sections

Cross sections were constructed from DEMs sampled 

at horizontal intervals of 10 m, are presented with no vertical exaggeration, and include both simple representations of 

successive surfaces and interpretive relations among units at 

depth. Our guiding philosophy in the construction of cross 

sections was not to extend geologic interpretation below levels 

for which we have no constraints. Therefore none is extended 

below our lowermost control surface, that of summer 1980.

Phases of Dome Growth at Mount St. 

Helens

Between the onset of unrest on September 23, 2004, 

and April 2006, the eruption developed in a manner that is 

divisible into five distinct phases, each with characteristic 

rate and pattern of eruption (table 1). An initial brief ventclearing phase included seismic unrest, spectacular deformation features, and phreatic explosions developed in the moat 

between the 1980s dome and the 1980 crater walls. Initial 

spine extrusion began October 11, 2004. As extrusion continued, the locus of spine growth shifted, spines grew and 

stagnated, and new ones formed in their stead (table 1). As of 

April 2006, a total of seven discrete spines have erupted that 

we have grouped on the basis of similar growth patterns into 

four additional phases (table 1).

Phase 1, Precursory Vent Clearing, September 

23–October 10, 2004: Phreatic Explosions and 

Deformation

The first indications of an impending eruption included 

a week of intensifying seismicity beginning September 23, 

2004, deformation-induced surficial cracks in glacier ice south 

of the 1980–86 dome that began to appear by September 29 

(Dzursin and others, this volume, chap. 14), and four phreatic 

explosions between October 1 and 5 (Moran and others, this 

volume, chap. 6). The phreatic explosions formed a vent at the 

west edge of deformed ice. Thermal IR images show that the 

explosions of early October had temperatures of no more than 

160°C (Schneider and others, this volume, chap. 17). On the 

basis of these low temperatures, we infer that the explosions 

were phreatic rather than magmatic. However, the explosions 

did indicate interaction of hot rock with the shallow hydrothermal system, thus suggesting rise of magma to near the surface.

During early October, a zone of highly fractured ice 

developed and expanded southward as subsurface intrusion 

fractured and thrust the part of the 1980–86 dome called Opus, 

which had formed in 1985, and adjacent crater-floor debris 

upward to form a feature named “the welt” (fig. 4). This shallow intrusion of magma caused surface uplift in excess of 70 

m (figs. 5, 6). Uplift was greatest along a north-south axis 

about 200 m east of the October vent (fig. 5) and diminished 

rapidly away from that axis.

The welt expanded southward, but motion of recognizable 

features through October 4 was upward and northward along 

a reverse fault with a strike of ~N. 80° E. and located between 

Opus and the remainder of the 1980–86 dome to the north 

(fig. 5). This faulting apparently reactivated a normal fault of 

1985 that bounded the north margin of the Opus feature (fig. 

1A). Surface deformation vectors south of the fault trace show 

motion of 25–30 m north and 50–70 m up. If deformation indicated motion along the fault, then its dip was ~60° south.

During October 4–14 the locus of deformation migrated 

south from Opus, and the sense of motion at the surface was 

radial, away from the most intense deformation. Motion on 

Opus was undetectable to barely detectable at ~5 m east and 

up (figs. 5, 6). On glacier surfaces, ballistic impact sites and 

distinctive avalanche toes near the periphery of the welt moved 

5–15 m away from the welt (southwest to east) and up (fig. 5). 

The few traceable points on the eastern and central parts of the 

actively deforming welt moved 30–80 m eastward away from 

the axis of the welt (fig. 5).

The DEMs of October 13 and 14 recorded displacement 

in the range of 5–30 m along the surface of the expanding welt 

(fig. 5). Points along the deformation axis, which coincided 

with the axis of spine 3 when it later emerged in November, 

moved 15–30 m S. 10° E. along the axis and rose 6–8 m in 

one day. Nearby features to the east moved laterally 12–15 m 

S. 20° E. to S. 45° E., with little vertical motion. Features on 

fractured ice to the east also moved away from the growing 

welt, whereas those on Opus showed no detectable motion.

Phase 2, October 11–24, 2004: Spines 1 and 2

Initial Spine Growth

Spines 1 and 2 each extruded rapidly within a few days 

and thereafter remained inactive, though each was affected 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 177

Phase Spine Time period Growth rate Nature of eruptive activity Deformation within 1980 crater Effect on glacier 1—Precursory vent clearing

Pre-dome

Sept. 23

–Oct. 10, 

2004

Not applicable

Vent clearing and phreatic 

explosions.

Ascent of solid spine to 

surface.

Uplift of 1980–86 dome, and 

crater floor to north of 1980–86 

dome to form welt. By Oct. 4, 

2004, total volume of deformed 

area (welt) is ~5×106

 m3.

Disruption and 

uplift of Crater 

Glacier.

 2—Initial spines 

Spine 1 Oct. 11–15, 

2004

2–3 m3

/s;

15–20 m/d

Near-vertical spine growth. Continuing uplift of crater floor to 

south of 1980s dome;

Volume of deformed welt 

increases, 5×106

–11×106 m3, Oct. 

4–13, 2004.

Continuing uplift 

and disruption of 

glacier but no appreciable melting.

Spine 2

Oct. 15– 

~Oct. 24, 

2004

3 m3

/s;

25 m/d

Advance of spine 2 to the 

south; probable subterranean and subglacial 

intrusion of spine 3.

Locus of deformation shifts southward and eastward; formation of 

roof pendant over intruding spine.

3—Recumbent growth of whaleback spines 

Spine 3

Oct. 25–

Dec. 18, 

2004

4–6 m3

/s;

8–11 m/d

Recumbent growth of spine 

3 toward south crater wall; 

spine 3 begins pushing 

against south crater wall 

~Nov. 12, 2004.

Emerging spine displaces older 

rocks to east and south; roof pendant is transported to south end of 

spine 3; spine 2 subsides.

Growth of spine 

3 divides Crater 

Glacier into east 

and west arms. 

Spines 3 and 4 

plow east Crater 

Glacier eastward 

and thereby 

thicken it as 

much as 100 m; 

crevasses form 

parallel to maximum principalstrain direction 

(~east–west) but 

no appreciable 

melting.

Transition

Dec. 18, 

2004–Jan. 

3, 2005

No data

Spine 3 deflects off south 

crater wall, fractures, 

breaks up, and decouples 

from source.

Spine 4 Jan. 3–Apr. 

9, 2005

1.5–2.5 

m3

/s;

5–8 m/d

Continuing extrusion forms 

spine 4, which continues 

pushing to the south.

Southward growth of spine 4 tilts 

and pushes spine 3 to east; deformed 1980–86 debris migrates 

eastward; remnants of spines 1–3 

to west are static

4—Spine thrusts over previous spines

Transition Apr. 10–19, 

2005 No data

Spine 4 encounters south 

crater wall, fractures, and 

decouples from source.

East Crater 

Glacier 

deformation 

slows, then stops; 

glacier responds 

to thickening 

caused by previous deformation 

by accelerating 

downslope to 

north. West Crater 

Glacier is pushed 

west.

Spine 5

Apr. 19-

July 18, 

2005 1–1.5 m3

/s;

3–6 m/d

Spine 5 thrusts over spine 

remnants west of spine 

4. Smooth surface forms 

at north end of spine and 

gradually steepens. In June, 

spine fractures to the south 

and disintegrates.

Deformation to east is greatest 

near vent and diminishes south. 

This deformation slows and stops. 

Deformation to west is moderate 

and continual. Rockfall from spine 5 

buries spine 2, then 1.

July 19–31, 

2005

Spine 5 crumbles to feed 

rockfall avalanches and 

slumping events.

5—Spines grow to west, 

 then thrust over one another

Transition Aug. 1–5, 

2005 No data

Spine 5 fractures near 

its source and begins to 

slump. West Crater Glacier thickens and 

cracks owing to 

westward migration of spines 6 

and 7; cracks 

radiate westward 

along maximum 

principal-strain 

axes.

Spine 6

Aug. 

6– Oct. 9, 

2005

1.5–2 m3

/s;

3–4 m/d

Sag depression grows 

owing to slumping of spine 

5 and westward migration 

of spine 6; spine growth 

is chiefly recumbent and 

endogenous.

Deformation to west; east part of 

2004–5 dome complex is stagnant.

Spine 7

Oct. 9, 

2005 –Apr. 

2006

0.5–1 m3

/s;

1–2 m/d

Endogenous growth followed by exogenous spine 

growth in depression.

Spine 7 pushes spine 6 to west and 

begins thrusting over elements both 

of itself and of spine 6.

Table 1. Timing, extrusion rates, character of dome growth, deformation within crater, and impact on glacier during 

each of five eruptive phases from September 2004 to April 2006, Mount St. Helens, Washington.

[Estimates of extrusion rates are derived from comparing DEMs from one date to the next (Schilling and others, this volume, chap. 8; and this 

chapter). Estimates of linear advance are obtained from tracking features in DEMs or aerial photographs (this chapter), from repeat photos from 

fixed positions (Major and others, this volume, chap. 12), and from portable GPS stations (LaHusen and others, this volume, chap. 16).]
178 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1980–86

 dome

Opus

Welt

~100 m

1980 crater wall

Fractured

uplifted ice

West 

 Crater 

 Glacier

East 

 Crater 

 Glacier

Figure 4. View of Mount St. Helens crater from north-northeast on October 10, 2004, illustrating welt, Opus, 

1980–86 dome, crater wall, and Crater Glacier. USGS photo by R. Wessels.

Figure 5. DEM of 

October 14, 2004, 

locating vent of 

October 4, 2004, and 

illustrating initial 

extrusion of spine 

1 and locations of 

welt, Opus, 1980–86 

dome, Crater Glacier, 

and cross section 

A–A ́ shown in figure 

6. Dots indicate 

features tracked, and 

solid lines indicate 

changes in fractured 

ice margin for dates 

identified by color 

in key. Arrows show 

surface deformation 

vectors; numbers 

show vertical 

component of vectors 

in meters.

+11

0

+95

+64

+72 +70 +54

+40 +23

+5

+3

+6

+5

+81

+21

+18

+68

+34

+32

+14

+5

+2

+12

+9

 0

wg1

?

+2

 +2 eg1

+34

10/4/2004

2003

A-A' section

10/13/2004

Fault trace

1980–86 dome

Opus

Crater Glacier

Welt

Axis 

Spine 3 

Nov. 4

Vent, Oct. 4

Spine 1

Crater Glacier

10/14/2004

0

0

+6

-1

+4 +8

A

A'

5116000

5115800

5116200

562400 562600 562800 563000

Crater 

Glacier

Vectors 

0 40 80 120 160 200

METERS
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 179

1,900

2,000

2,100

2,200

ALTITUDE, IN METERS

1980s 

 debris

 Crater

Opus, 1985–2003

 Opus, 

deformed

Crater Glacier, 2003

Crater Glacier, 2003

1980s 

 debris

 1980s 

 debris

 1980s 

 debris

Crater Glacier Crater 

Vent Glacier

Hot

rock

 Opus, 

deformed

 Deformed

ice and water

October 4, 2004

1980s 

 debris

 Opus, 

deformed

 1980s 

 debris

Crater Glacier

Crater 

Glacier

 Opus, 

deformed

November 4, 2004

0 100 200 300 400 500 600 700 800 900 1,000

DISTANCE, IN METERS

1,900

2,000

2,100

2,200

1,900

2,000

2,100

2,200

 Opus

Crater Glacier Crater 

Glacier

2003 2003

1986

1980

11/4/04

10/13/04

1986

1980

1,900

2,000

2,100

2,200

2,300

A A ́

Spine 1 

11/4/04

West East

 Crater

10/4/04 11/4/04

10/13/04

10/4/04

2003

1986

1980

Spine 1

Spine 2 Spine 3

10/4/04

2003

1986

1980

Figure 6. Cross section A–A ́ (location shown in figure 5) illustrating extrusion of spine 1, intrusion of 

spines 2 and 3, and deformation of Opus and glacial ice. Top panel shows known profiles for the dates 

given, and other panels illustrate geologic interpretations on given dates.
180 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

subsequently by deformation and eventual burial. Warming 

at the surface of the welt to as much as 50°C on October 10, 

2004, presaged the appearance of juvenile rock above ground on 

October 11 (Schneider and others, this volume, chap. 17). Spine 

1 extruded from October 11 to 15. Owing to poor weather, the 

timing of spine 2 emplacement is less well constrained. Observations and study of oblique photographs showed that its growth 

began after October 15 and finished by October 24.

Spine 1 grew at a steep angle as a solid blocky slab from 

a south-facing slope on the west part of Opus. By October 14, 

it was 85×60 m, ~60 m high, dipping 50° northwest, and had 

volume of ~0.6×106 m3 (table 2, fig. 6). Using the outline of 

spine 1 on the DEM of November 4, we calculate that the total 

volume of spine 1 was then ~0.9×106 m3. Spine 1 extruded 

rapidly, 15–20 m/d and 2–3 m3/s (table 2).

Although observations between October 15 and 24 

were insufficient to delineate the nature of spine 2’s growth 

directly, it emerged just south of spine 1 and appeared at the 

surface from north to south as though the deformed glacier 

ice from which it emerged had unzipped rapidly southward. 

It emanated from the October 4 vent (figs. 5, 7) and moved 

southward along subglacial slopes to form an elongate body 

oriented north–south at the surface (table 2). Its growth ceased 

when it encountered steep opposing slopes of the 1980 crater. 

Assuming vertical boundaries downward from the November 

4 surface to the 1986 surface allows a volume calculation of 

about 2×106

 m3. The spine’s estimated volume and interval of 

emplacement imply a time-averaged extrusion rate of ~3 m3/s.

Deformation Accompanying Spine Growth

Between October 14 and November 4, the locus of maximum deformation propagated south along an axis oriented ~S. 

20° W. Deformation diminished with distance normal to this 

axis. Opus and the 1980–86 dome were essentially static during this period (fig. 7). Severely deformed ice on the welt just 

east of its axis moved a few meters to the north; farther south, 

it moved as much as 60 m to the southeast and subsided (fig. 

7). On glacier surfaces to the east, distinctive features moved 

5–40 m eastward and rose 5–20 m. Farther north, glacial features moved a few meters northward. The few traceable points 

on ice west of the actively deforming welt rose a few meters, 

but only one feature nearest the northwest margin of the welt 

moved significantly westward, by ~10 m (fig. 7, point wg1).

Phase 3, Recumbent Growth of Spines 3 and 4: 

October 24, 2004–April 9, 2005

Growth of Whaleback Spine 3: October 

24–December 18, 2004

During mid-October, spine 3 began intrusive growth and 

pushed into pre-2004 rock, deforming the welt, disrupting 

glacier ice, and forming a cryptodome beneath the welt (fig. 

8). Evidence in support of intrusion included (1) deformation 

along an axis S. 19° E. that coincided with the axis of spine 

3 when it later emerged (fig. 5), (2) upward and southward 

motion of pre-2004 rock along a trend similar to that of the 

whaleback when it emerged (figs. 5, 7), and (3) warming of 

rock at the surface near the axis of deformation (Schneider and 

others, this volume, chap. 17).

Spine 3 breached the deformed surface of the welt and 

advanced rapidly to the south-southeast between late October 

and mid-December to form a smooth-surfaced whaleback 

feature 300–460 m long and 120–145 m wide (tables 1, 2). 

Between October 24 and 27, spine 3 emerged from an area 

about 50 m southeast of spine 1, through older dome and 

crater-floor rock along the crest of the deforming welt (fig. 9). 

By early November, it was 320×125 m, with a long axis oriented S. 18° E. (fig. 7). As the spine pushed southward from 

mid-November through December, its long axis pivoted 9° 

eastward (to S. 27° E.) about its origin at the vent (table 2).

The surface of the whaleback had a cool and smooth, but 

striated, surface except on the west, where it was broken and 

blocky. The striations at the surface of the whaleback were 

interpreted as slickensides (fig. 10). Growth of the whaleback 

also lifted a partial roof composed of fractured 1980–86 dome 

rock and crater-floor debris and transported it southward 

during November 4–29 (fig. 10). Cashman and others (this 

volume, chap. 19) show that the smooth outer carapace of 

the whaleback comprised powdered, partially sintered 2004 

dacite plus small amounts of 1980–86 dacite; they interpret 

this material as fault gouge formed through comminution as 

the solid spine rubbed and ground against older rock during its 

ascent in the conduit.

The stable crust of spine 3 insulated the hot rock within 

so that surface temperatures were low. Thermal images 

commonly showed a ~200°C zone around the base of the 

emerging spine (fig. 11A), and hot cracks showed rock 

temperatures as high as 730°C. The temperature of smooth, 

uncracked parts of the surface diminished exponentially 

with distance from the source at the base of the spine and 

approached ambient within 50 m (fig. 11B). Because the 

temperature also decreased exponentially with time (each 10 

m from source represented about a day), the moving surface 

of the spine showed a classic Fourier’s Law decline in temperature (Turcotte and Schubert, 1982). Such a temperature 

profile resembles those observed for blocky lavas at Santiaguito, Guatemala (Harris and others, 2002, 2004), though 

the lava core temperatures of Mount St. Helens spines differ 

from those of lavas at Santiaguito in likely being 100°C or 

more below solidus temperature.

Subtracting the spine 1 and 2 volumes from the total volume of hot rock emplaced by spines 1, 2, and 3 (Schilling and 

others, this volume, chap. 8) yields values for spine 3 volume 

and for its extrusion rate between late October 2004 and early 

January 2005 (table 2). Extrusion rate (flux through the 1986 

surface) was ~5 m3

/s until November 4. Extrusion rates then 

declined from 4.4 to 2.5 m3/s between late November and early 

January (table 2).
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 181

Phase Date of DEM Spine/

Welt

Dimensions

(m)

Orientation of 

welt or spine

Linear growth rate, direction, period of 

interest, and GPS station name where 

applicable

Volume

(×106

 m3 )

Volumetric 

growth rate

(m3

/s)

 1—Precursory 

vent clearing

Oct. 4, 2004 Welt 370×370 Equant ~70 m/d, South, Sept. 29–Oct. 4 5 a

12

Sept. 

29–Oct. 4

Oct. 14, 2004 Welt 470×380 Long axis northsouth ~10 m/d, South, Oct. 4–14 11 a 7

 2—Initial spines

Oct. 14, 2004 Spine 1 85×60 Strikes S55W; 

dips 50° NW 15-20 m/d, S35E, Oct. 11–14 0.6 ~2

Oct. 11–14

Nov. 4, 2004

Spine 1 150×45 Strikes S58W; 

dips 80° NW 0.9 ~3

Oct. 14–15

Spine 2 240×50 Long axis, Due S ~25 m/d, South, Oct. 15–24 2 3

Oct. 15–24

 3—Recumbent growth of whaleback spines

Nov. 4, 2004 Spine 3 320×125 Long axis S18E;

dip at vent 30° 11.4 m/d, ~ S5W, Nov. 4–7 9 4.7

Oct. 13–Nov. 4

Nov. 20, 2004 Spine 3 420×125 Long axis S19E 11 m/d, S22E, Nov. 4–20

Nov. 29, 2004 Spine 3 440×145 Long axis S23E 10.3 m/d, S19E, Nov. 21–29, ELEA b

10.5 m/d, S21E, Nov. 20–29 18 4.4 a

Nov. 4–29

Dec. 11, 2004 Spine 3 460×120 Long axis S27E 7–8 m/d, ~S20E c 23 4.1 a

Jan. 3, 2005 Spine 4 210×130 Long axis S27E 4–7 m/d, ~S20E c 5.5 2.5 a

Feb. 1, 2005 Spine 4 340×145 Long axis S28E 8.3 m/d, S19E, Jan. 15–16, CDAN b 10 1.8 a

Feb. 21, 2005 Spine 4 400×150 Long axis S31E 5.8 m/d, S30E, Feb. 8–14, AHAD b

4.5 m/d, S32E, Feb. 1–21 14 2.4 a

Mar. 10, 2005 Spine 4 440×140 Long axis S35E 3.9 m/d, S39E, Feb. 21–Mar. 10 17 1.8 a

Apr. 19, 2005 Spine 4 490×140 Long axis S40E

2.6 m/d, S71E, Mar. 10–Apr. 19

Spine 4 decouples from vent during 

this interval.

21 1.5 a

4—Spine thrusts over 

previous spines

Apr. 19, 2005 Spine 5 100×90 Long axis S25E;

dip at vent 40° 3.5–4.5 m/d, S10E c 1 1.5 a

June 15, 2005 Spine 5 340×170 Long axis S9E 3–4 m/d, S along axis c 8 1.4 a

July 14, 2005 Spine 5 285×105 Long axis; S5E;

dip at vent 54° 2–3.5 m/d, S along axis c 11 1.3 a

Aug. 10, 2005 Spine 5 15

5—Spines grow to west then thrust over one another

Aug. 10, 2005 Spine 6 100×90 Long axis S43W 3–4 m/d, ~West c 1; compound 

spine 6 is ~12 2.0 a

Sept. 20, 2005 Spine 6 350×280 Long axis S47W 3–4 m/d, ~West c 6; compound 

spine 6 is ~17

1.6 a

Oct. 24, 2005 Spine 7 110×40 Long axis S44W 3–4 m/d, ~West c New lava, 

spines 6+7, ~9

0.9 a

Dec. 15, 2005 Spine 7 260×~250 Long axis S74W ~3 m/d, ~West c New lava, 

spines 6+7, ~12 0.7 a

Feb. 9, 2006 Spine 7 310×~300 Long axis S83W;

dip at vent 50° 2.2 m/d, S80W, up 50° New lava, 

spines 6+7, ~15 0.6

a Schilling and others (this volume, chap. 8).

b

 LaHusen and others (this volume, chap. 16).

c

 Major and others (this volume, chap. 12).

Table 2. Dimensions, orientation, volume, and growth rate of the welt and spines of Mount St. Helens, Washington, at various dates.

[Applicable interval for linear and volumetric growth rates and direction is from date of previous DEM to date of DEM given for row, unless specifically indicated.]
182 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

562200 562400 562600 562800 563000 563200

5116200

5116000

5115800

5115400

5115600

0 100 200 300 400 500

METERS

10/4/04 vent

11/4/04 features and glacier margin

11/20/04 features and glacier margin

11/29/04 features and glacier margin

12/11/04 features and glacier margin

10/14/04 features

Vectors

Squares indicate immobile features 

Circles indicate moving features

Triangles indicate GPS stations

+11,+10,+6,+9

 eg1

_,_,+24,+12

-9,-6,0,-4

94 m, +14

S21E

9 days

0,+3,+4,+7

 wg1

+7,+5,+3,+4

+3,0,+2,+4

 _,+1,+5,+2

_,+1,+3,+3

_,-10,-10,-4

_,_,-4,-5

0,-10, 0, 0

+1,0,+2, 0

0,-2,0,0

0,-10,0,0

+5,-2,0,+2

+3,0,+2,+4

ELEA

82 m,+8

S19E

8 DAYS

CLF4, 

20 m,-8

S30W

19 days

_,-3,+3, 0 166 m, +61

S23E

16 days

94 m, +14

S21E

9 days

ELEA

82 m,+8

S19E

8 DAYS

166 m, +61

S23E

16 days

Opus

West 

Crater

Glacier

East

Crater

Glacier

1980–86 dome

Spine

3

Deformed rock and ice

+19,+11,+5,+6

E

E ́

B

B ́

MID9

-13,-11, 0, 0

-4,-24,-2,-3

_, +10 Vertical component, in meters 

Underscore, no data

Welt

Welt

Spine

2

Spine 1

Figure 7. Deformation vectors for the period October 14–December 11, 2004, shown on DEM of November 20. Squares 

and dots indicate features tracked for dates identified by color in key. These colored symbols delimit deformation 

vectors of each of the four time intervals between relocation of features. For quantitative data (example: _, -3, +3, 0), 

numbers are vertical components of vectors, in meters, for each interval. Underscores indicate no data for that interval. 

Dotted lines indicate dome-glacier margins for dates indicated by color in key. Lines B–B ́and E–E ́ are traces of cross 

sections shown in figures 8 and 14, respectively.

Growth of Spine 3: Onset of Recumbent Growth

Although the spine in early November 2004 appeared 

to grow more or less vertically from the axis of the welt, it in 

fact originated from a source beneath spine 1 and extended 

southward two to three times as fast as it pushed upward. 

Between November 4 and 7, identifiable features on the 

spine’s surface moved ~32 m S. 5° W. (fig. 10). Vertical 

movement of the spine was poorly constrained but appeared 

to be ~12 m up. The south end of the spine extended almost 

30 m S. 19° E., an orientation matching that of its long axis. 

A possible explanation for the contrast in surface-vector 

directions and overall spine extension is that, at this early 

stage of its growth, the spine rotated slightly westward while 

pushing south-southeast along its axis.

Deformation Adjacent to Spine 3 in Early 

November

Between November 4 and 7, areas to the east of spine 3 

moved tens of meters parallel to the spine or were pushed in 

easterly directions, and areas to the west and north of spine 

3 moved less than 10 m or remained static. Rock debris adjacent to the eastern margin of spine 3 moved parallel to the 

east margin by almost 30 m (fig. 10A). Rock debris farther 

east moved smaller distances. Motion of the rock debris had 

no detectable vertical component. Fractured and previously 

uplifted glacial ice less than 200 m east of the spine moved 

10–15 m along trends ranging from southeast to east (fig. 

10A). During November 4–7, a GPS unit west of the whaleback, CLF4, moved 8 m south-southwest and subsided 2 m 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 183

1,900

2,100

2,300

2,500

11/29/04

11/20/04

11/4/04

10/13/04

10/4/04

2003

1986

1980

ALTITUDE, IN METERS 1,900

2,100

2,300

1,900

2,100

2,300

1,900

2,100

2,300

500 700 900 1,100 1,300 1,500

DISTANCE, IN METERS

Spine 1

Spine 3

Spine 1

Spine 1

Spine 3

spine 3

Ice

Ice

Ice

Deformed rock

1980s debris

Deformed

 rock

1980s debris

Deformed rock

1980s debris

North South

1980s

dome

1980s

dome

1980s

dome

Deformed

 rock

Deformed

 rock

Deformed

 rock

Deformed

 rock

Deformed

 rock

Spine 1

Spine 3, whaleback

Vent

Oct. 4, 2004

1980s

dome

1980s dome

1980s dome

1980s dome

Welt

Faults

faults

Motion of ice

is chiefly to east

away from view

October 13, 2004

November 4, 2004

November 20, 2004

Motion of welt

is chiefly to east

away from view

B B ́

Crater

 wall

Crater

 wall

Crater

 wall

Figure 8. North-south cross section B–B ́(location shown in figure 7) 

illustrating intrusion and recumbent growth of spine 3. Top panel shows 

known profiles for the dates given, and other panels illustrate geologic 

interpretations on given dates.
184 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 9. Mount St. Helens crater floor on October 27, 2004; view 

from northwest, illustrating spines 1–3. A, Photograph showing 

1980–86 dome and ongoing extrusion, with fresh ash on Crater 

Glacier west of new dome. B, Mosaic of thermal-infrared images 

located in A from same date and time. Numbers indicate spines 

1–3. Maximum temperature in field of view is 600°C (scale is set 

such that all temperatures greater than 300°C are white). Spine 3 

has emerged since previous TIR survey on October 24. Thermal 

imagery shows hot debris fan and hot blocks on ice and firn near 

spine 1 and between spines 1 and 2. USGS photo by J.W. Vallance; 

thermal images by D.J. Schneider.

(LaHusen and others, this volume, chap. 16). Nearly static 

areas during this time included spines 1 and 2, Opus, and the 

west arm of Crater Glacier.

Striations and Relative Motion of Spine 3 in Early 

November

Growth of spine 3 and nearby deformation patterns explain 

otherwise enigmatic bidirectional striation patterns on the spine 

in early November 2004. On November 4, 7, and 10, striations 

plunging 31° N. 66° E. on the east face of spine 3 were superimposed on fainter striations plunging 20° N. 5° E. (fig. 10B). 

At the north end of the spine near the vent, only fainter N. 5° E. 

striations existed. A cursory analysis in early November suggested that the striations recorded a change in direction of spine 

growth, with the fresh striations indicating the most recent direction. However, our photogrammetric analysis indicates instead 

that welt debris adjacent to the whaleback was being dragged at a 

rate of almost 10 m/d along the base of the emerging spine along 

a ~S. 32° E. trend (fig. 10). This vector minus the true-growth 

vector yields a vector whose direction (N. 66° E.) matches the 

direction of the freshest east-flank striations (fig. 10B). Deformation patterns east of the spine thus suggest a simple explanation 

in which the fresh striations recorded a growth direction relative to debris being dragged southward, and the faint striations 

recorded the true growth direction of the spine.

Continuing Recumbent Growth of Spine 3

Through November, spine 3 continued to advance to 

the south-southeast (S. 19° E. to S. 23° E.) at a linear rate of 

10–11 m/d (fig. 7). Judging from the slope of the northern 

face, initial vertical components of motion were ~30°. As 

features on the spine moved to the south, their vertical motions 

gradually diminished to zero by ~300 m from the vent, and 

thereafter they began to subside (LaHusen and others, this volume, chap. 16; this study). The southerly (S. 5° W.) motion of 

features on the spine turned to south-southeast in mid-November (fig. 7). As the spine axis and growth direction converged, 

its whaleback form began to take on a smooth gouge-covered 

appearance on the east and west flanks.

Superposition of spine 3 outlines from November to 

December 2004 upon the 1986 surface suggests that the growing spine encountered opposing slopes of the 1980 crater wall 

in mid-November (fig. 8). The surface outline of the spine first 

overlapped the steep north-sloping crater wall at the bed of 

the glacier sometime between November 4 and 20 (fig. 12A). 

Northeastward acceleration of GPS spider MID9 north of 

spine 3 and south of the 1980–86 dome beginning November 

12 probably corresponded to the time at which spine 3 began 

to push against the opposing slope of the 1980 crater beneath 

the glacier (fig. 13).

As spine 3 continued to impinge on steep, opposing 

subglacial slopes of the crater wall in late November and 

December, the spine axis began to rotate eastward (table 

2, fig. 12A) and the spine began to break. Axis orientation changed 8° eastward from November 20 to December 

11 (table 2). Between November 20 and 29, fracturing and 

separation of the first piece of spine 3 was apparent (figs. 

2A, 2B), consistent with northeastward acceleration of GPS 

station MID9 during November 23–27 (“1st breakup” in fig. 

13A). By November 29, this small spine fragment had completely separated, and by December 11, the rotating spine 

had pushed it into a steaming heap along its margin with east 

Crater Glacier (fig. 2C). A second, more substantial fragment separated and decoupled in early December. Northeastward acceleration and deceleration of MID9 suggests that 

the period of the fracture and decoupling of the resultant spine 

fragment spanned December 6–12 (“2nd breakup” in fig. 13A). 

TEMP.

(oC)

300

-10

2

1

3

A

B

1980–86 ~100 m

 dome
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 185

Spine 1

Spine 2

Spine 3

Roof 

pendant

1980s

debris

Resultant

parallels

slickensides

 20 m

 N66E

 32 m

 S5W

+12 m

+12

 28 m

S32W

 CLF4

8 m S23W

 -2 m

10/4/04 vent

100 150 200

METERS

Vectors, spine 3

Crater Glacier

Crater

Glacier

Moving features

Fixed features

Welt

N

 32 m S5W

+12 m vertical

 28 m S32E

 -horizontal

Resultant

20 m N66E

-12 m vertical

Roof pendant

1980s debris

B Sense of

 motion

A

Fractured

glacial ice

of the welt

0 50

Vectors; motion 

 of welt features

Figure 10. Mount St. Helens crater floor in early November 2004. A, Rectified vertical aerial photograph of dome taken November 

7, 2004, with handheld camera from an airplane. The photograph was subsequently rectified using features on 1980–86 dome, Opus,

and Crater Glacier for both November 4 and 7 so that its central part is tied to geographic coordinates of the November 4 DEM. The 

aerial photograph shows deformation vectors in plan view for November 4–7. B, Oblique photograph looking west at spine 3 on 

November 4, 2004, showing faint striations parallel to deformation vector, 

Vblack , and young pronounced striations parallel to vector 

Vorange , where subscripts indicate vector colors in photograph. Debris adjacent to whaleback moves in the near-horizontal plane 

parallel to 

Vyellow such that yellow black VVV orange

 =− . USGS photos by J.S. Pallister.
186 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

50

100

150

200

250

300

TEMPERATURE, °C

A

A'

A ~100 m A'

Source Cracks

Warm

debris

PIXEL DISTANCE

Spine 3 source, hot basal zone

Sense of

motion

Crack 

~100 m

from 

source

100 200

A

B

TEMP.

(oC)

350

-15

Figure 11. Thermal aspect of spine 3 surface on November 29, 

2004. A, Thermal-infrared image from helicopter, view to southeast, 

by J.W. Vallance. B, Temperature profile from north to south.

There was no big north-south crack in a December 4 image 

(R.L. Helz, written commun., 2004). However, a December 8 

image revealed a well-developed crack, confirming that during 

that interval a substantial part of the spine had begun to separate from its west flank along a longitudinal fracture oriented 

parallel to the spine’s axis. December 11 photographs (fig. 2C), 

DEM (Schilling and others, this volume, chap. 8), and geologic 

map (Herriott and others, this volume, chap. 10) revealed that 

this fragment had largely decoupled and had begun to subside 

westward by that date. During November to mid-December 

numerous other small spine fragments separated from spine 3, 

mostly to the west. Generally, fragments separating to the west 

slowly subsided and stagnated, whereas the few separating to 

the east were bulldozed eastward.

Late October–Mid-December Deformation of Surrounding 

Areas during Spine 3 Growth

In response to recumbent dome growth from late October 

to mid-December, areas east of spine 3 rotated eastward about 

a pivot axis near Opus, showing horizontal displacements as 

great as 100 m. Areas west of spine 3 subsided or stagnated 

and moved less than 20 m (figs. 7, 14A). All areas east of spine 

3 rotated eastward in such a fashion that displacement per unit 

time increased from north to south. Ice masses on the welt subsided and then migrated eastward across the boundary between 

stagnant ice and the Crater Glacier, whereupon they began to 

rise. During this interval all ice on the east Crater Glacier rose 

between 10 and 50 m, with largest vertical displacements occurring to the south and the smallest to the north (figs. 7, 14A). 

In effect, rotating dome and welt rock bulldozed the glacier 

eastward, compressing and lifting it such that its cross-sectional 

profile changed from concave to convex (figs. 2, 14A).

Spine 1 moved in response to the emerging spine 3, then 

stabilized as that new spine continued to grow. Spine 1 was 

displaced about 40 m north-northwest and tilted from a dip of 

50° to 80° northwest between mid-October and early November (fig. 7). Comparison of the October 14 DEM with those of 

November (figs. 5, 7) shows the change in position of spine 1 

with respect to the October 4 vent position. This displacement 

and rotation coincided with southward growth and emergence 

of spines 2 and 3. The displacement and rotation were to the 

northwest, directly away from the origin of spine 3’s southward extrusion, and are likely to have been a response to spine 

3 emergence rather than growth of spine 2.

Spine 2 and the area between spines 2 and 3 rotated 

slightly and subsided between October and December. In late 

October and early November, spine 2 moved as much as 20 m 

to the northwest at its south end but remained relatively stable 

to the north. In late October, the top of spine 2 was as much 

as 20 m higher than the adjacent glacier surface (fig. 9). As 

spine 3 grew, spine 2 subsided by 20–25 m (fig. 7). By late 

November, spine 2 had become a nondescript entity hidden 

below the level of west Crater Glacier, with talus and rockfall encroaching upon it from the more prominent spine 3 to 

the east (fig. 2B). The GPS spider CLF4, placed on the welt 

October 27, was perched between spines 2 and 3 as spine 3 

emerged. The spider had moved ~44 m S. 24° W. and subsided 

13 m by November 19, when it was buried by rockfall from 

spine 3 (LaHusen and others, this volume, chap. 16). As spine 

3 advanced, it pushed parts of the dome south and west of it 

to the southwest, but, after its leading edge passed by and the 

spine began rotation in the opposite direction, subsidence of 

areas west of spine 3 began—possibly caused by removal of 

buttressing from the east.

West Crater Glacier, south and west of the spine complex, 

responded during a brief interval as spine 3 approached but 

otherwise remained relatively static from late October to midDecember (fig. 7). Dome growth after November 4 caused no 

detectable deflection of west Crater Glacier. A slow increase 

in altitude of the west glacier, caused in part by accumulation 

of snow, occurred as spine 3 grew. Areas of the glacier immediately south of the dome complex moved tens of meters as the 

whaleback approached (fig. 7), but then they stopped.

Breakup of Spine 3 and Formation of Spine 4: 

December 18, 2004–April 8, 2005

Spine 3 broke up and decoupled as spine 4 formed 

between mid-December 2004 and early January 2005. Photograph sequences showed fractures oriented diagonally across 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 187

A

B

0 100 200 300 400 500

10/4/04 vent METERS

3/10/05 spine 1

1/3/05 spine 4, glacier margin

3/10/05 spine 2

2/1/05 spine 4, glacier margin

2/21/05 spine 4, glacier margin

3/10/05 spine 4, glacier margin

4/19/05 spine 5, glacier margin

4/19/05 spine 4 remnants

Opus

Spine 3 

remnant,

April 19

Spine

4 

Spine

5 

4/19/05 spine 4 is decoupled

 and spine 5 is active

Spine 3 

remnant,

April 19

Spine 4 

remnant

that was

active in

early April

Spine

2 

562200

5116400

562400 562600 562800 563000 563200

5115600

5115800

5116000

5116200

562200

5116200

562400 562600 562800 563000 563200

5115400

5115600

5115800

5116000

563400

Space east 

of vent in

mid-April

Spine 4 

remnant,

April 19

0 100 200 300 400 500

METERS 10/4/04 vent

11/4/04 spine 1

11/4/04 spine 3

11/4/04 spine 2

11/20/04 spine 3, glacier margin

11/29/04 spine 3, glacier margin,

 spine 3 remnants

12/11/04 spine 3, glacier margin,

 spine 3 remnants

1980–86 dome

Opus

Spine

3 

Spine

2 

Spine

1 

1980–86 dome

Spine

1 

Figure 12. Outlines of active spines and dome-glacier margins at Mount St. Helens, Washington, 

superimposed on 1986 DEM. A, November 4–December 11, 2004. B, January 3–April 19, 2005.
188 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 13. Plots of displacement versus time for GPS stations MID9 (A), located between 2004–5 and 1980–

86 domes, and TOP7 (B ), located at apex of 1980–86 dome, with intervals of spine growth and key events 

during spine growth, as well as linear extrusion rates of GPS receivers and features tracked in this study 

superimposed on GPS time series. Initial locations shown in figures 7 and 15. Because of their locations just 

northeast and north of spines 3 and 4, MID9 and TOP7 stations commonly accelerated opposite spine growth 

when spine met resistance to its growth and decelerated when that resistance was relieved. GPS data from 

LaHusen and others (this volume, chap. 16).

-0.2

-0.1

0

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

10/2/04 10/22/04 11/11/04 12/1/04 12/21/04 1/10/05 1/30/05 2/19/05 3/11/05 3/31/05 4/20/05

DATE

0

2

4

6

8

10

 DISPLACEMENT RATE, IN METERS PER DAY

CDAN

S19E +34o

1 day

AHAD

S30E +14o

7 days

 HNY0

S76E -9o

24 days

 S32E +32o

 S79E -3o

 S71E +6o

 S39E +23o

 S85E -4o

 S81E -9o

Spine 3

Spine 1

Spine 2

Spine 4

Thrusting events Transition

-0.5

0

0.5

1

1.5

ELEA

S19E +6o

8 days

 S21E +9o

Spine 3

Spine 1

Spine 2

Spine 4

Transition

MID9

10/22/04 11/11/04 12/1/04 12/21/04 1/10/05 1/30/05 2/19/05 3/11/05 3/31/05

Spine 5 transition Spine 5 transition

Whaleback impinges onopposing crater slope

1st breakup

2nd breakup

Transverse

crack forms Spine 4 decouples

from spine 3

 S5W 

~+20o

14

12

+50o

S30E

Whaleback impinges onopposing crater slope

1st breakup

2nd breakup

Transverse

crack forms

Steady recumbent growth

S5W growth spine 3 S5W growth spine-3 whaleback

Vallance, ch. 9 Figure 13

DISPLACEMENT, IN METERS DISPLACEMENT, IN METERS

A

B

Northeast

of 2004–5

dome

TOP7

Top of

1980–86

dome

Spine 1

Spine 3

Spine 4

Displacement

 rates—Thick bars,

 GPS; thin bars,

 this study

Easting

Northing 

Vertical

Easting

Northing 

Vertical
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 189

Figure 14. East-west cross section E–E ́ (location shown in figures 1 and 7) illustrating response of glacier to recumbent 

growth of spine 3. A, Profiles for October 13–December 11, 2004. B, Profiles for January 3–April 19, 2005.

spine 3 that formed and became more prominent beginning 

about December 18 (Major and others, this volume, chap. 12). 

Station MID9 reversed its direction of motion from northeast to southwest on December 20 (LaHusen and others, this 

volume, chap. 16), an event correlated approximately to the 

formation of the transverse fractures (fig. 13). Southward 

thrusting of spine 4 over the slowing bulk of spine 3 had 

become evident by December 24–28 (figs. 2D, 2E; supplementary movie 2 in Iverson and others, 2006). A second 

fracture south of the first formed by January 3, 2005, and the 

decoupling of spine 4 from spine 3 could be considered complete by that time (fig. 2E). We infer that cessation of MID9’s 

southwestward motion on January 3 corresponded to complete 

decoupling of spines 3 and 4.

Between January 3 and April 2005, spine 4 pushed 

south-southeast from a source located ~70 m east of spine 1. 

Spine 4 had whaleback morphology similar to that of spine 3 

but underwent several cycles of thrusting over spine 3 before 

it established steady near-uniform growth. Photographs (figs. 

2E, 2F) and time-lapse video (Iverson and others, 2006; Major 

and others, this volume, chap. 12, appendix 1) reveal thrusting 

events during January 6–12 and January 14–February 2. The 

TOP7 GPS spider accelerated northward during each of these 

events (fig. 13B). Time-lapse photography and GPS records 

for the period of February 2–April 9, 2005, suggest nearly uniform, steady southward growth and no further thrusting events.

By early January spine 4 had a crest oriented S. 27° E. 

and dimensions of 210×130 m (table 2, fig. 15). Spine 4 grew 

to ~440 m in length by early March before beginning to break 

up in April (fig. 16). As with spine 3 during November and 

December, the long axis of spine 4 began to pivot eastward 

about its origin at the vent. This rotation began once the spine 

started to impinge on slopes of the crater wall (figs. 12, 17); 

its long-axis orientation swung 12° eastward between midFebruary and its breakup in mid-April (table 2, fig. 12B). This 

response suggests that the moving spine extended deep enough 

to be influenced by the slopes of the 1980 crater at depth, as 

modeled in cross sections (figs. 16, 17).

Linear extrusion rates of spine 4 diminished during January 3–April 10, as shown by analysis of photographs taken 

at hourly to daily intervals from fixed sites at the eastern 

crater mouth (Sugar Bowl) (Major and others, this volume, 

chap. 12), the motions of GPS spiders on the spine (LaHusen 

and others, this volume, chap. 16), and time-averaged results 

from this study (fig. 13, table 2). A GPS station, CDAN, 

moved 8 m/day along a path of S. 19° E. and 34° upward 

between January 15 and 16. A second station, AHAD, moved 

5.9 m/d along a path of S. 30° E. and 14° upward between 

February 8 and 15 (fig. 13) (LaHusen and others, this volume, chap. 16). A rate and direction estimate using features 

on successive DEMs for the longer interval between February 1 and 21 was slower: 4.5 m/d along a path of S. 32° E. 

2,000

2,100

2,200

 2004 dacite,

fractured 1980–86 rock,

and crater-floor debris

 East 

 Crater

Glacier

 West 

 Crater

Glacier

ALTITUDE, IN METERS

Crater-floor debris

West East

1,900

2,000

2,100

2,200

0 100 200 300 400 500 600 700 800 900 1,000 1,100 1,200 1,300

 East 

 Crater

Glacier

 West 

 Crater

Glacier

DISTANCE, IN METERS

Crater-floor debris

Spine 2

Spine 2

Spine 3

Spine 4 Spine 3

Spine 4

A

B

Spine 4

Light dash 

is trace of 

12/11/04 

from A

 2004–5 dacite,

fractured 1980–86 rock,

and crater-floor debris

E E ́

12/11/04

11/29/04

11/20/04

11/4/04

10/13/04

10/4/04

4/19/05

3/10/05

2/21/05

2/1/05

1/3/05

12/11/04

2003

1986

1980

2003

1986

1980
190 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0 100 200 300 400 500

METERS

10/4/04 vent

1/3/05 features, glacier margin

2/1/05 features, glacier margin,

 and spine 3 remnant

2/21/05 features, glacier margin

3/10/05 features, glacier margin

Squares, immobile features

Circles, moving features

Numbers, vertical changes, in meters

Triangles, GPS stations

4/19/05 features, glacier margin

West

Crater 

Glacier

East

Crater Glacier

1980–86 dome

Opus

Spine 4

on 3/1

C

C ́

E

E ́

CDAN

1 day

AHAD

7 days

HNY0

20 days

ICY5

TOP7

MID9

Vectors: red, GPS; yellow, this study

Spine 2

Spine 1

12/11/04 features

562200 562400 562600 562800 563000 563200

5116400

5116000

5115600

5115800

5116200

+15,+12,+5,-2,+8

-11,-20,+7,-10,-8

+3,+3,+2,+2,_

+11,+11,+1, 0,+7

+4,-1,+5,-1,+2

-1,-5,-6,-2,-5

_,-17,-9,-4,-8

_,_,+47,+30,+11

-4,-4,0,0,+16

 

0,+4, 0,-1,+6

-4,+4, 0,-3,+2

-4, 0,-1,-3,_

-1,-1, 0,+3, 0

-12,-1,-1,-4,-12

-6,-12, 0,-4,-1

+8,-2, 0,-4,+2

-11,+4,-1,-3,+2

no change

no change

_,-5, -5,-5,_

_,-18, -2,_,_

+15,+12,+5,-2,+8

 eg1

-11,-20,+7,-10,-8

+3,+3,+2,+2,_

+11,+11,+1, 0,+7

+4,-1,+5,-1,+2

-1,-5,-6,-2,-5

_,-17,-9,-4,-8

_,_,+47,+30,+11

-4,-4,0,0,+16

 wg1

0,+4, 0,-1,+6

-4,+4, 0,-3,+2

-4, 0,-1,-3,_

-1,-1, 0,+3, 0

-12,-1,-1,-4,-12

-6,-12, 0,-4,-1

+8,-2, 0,-4,+2

-11,+4,-1,-3,+2

no change

no change

_,-5, -5,-5,_

_,-18, -2,_,_

Spine 3 

remnant

Figure 15. Deformation vectors for the period from January 3 to April 19, 2005, shown on DEM of February 1, 2005. 

Squares and dots indicate features tracked for dates identified by color in key. These colored symbols delimit 

deformation vectors of each of the five time intervals between relocation of features. Numbers are vertical components 

of vectors, in meters, for each interval. Underscores indicate no data for that interval. Dotted lines indicate domeglacier margins for dates indicated by color in key. Dashed lines C–C ́, and E–E ́ are traces of cross sections shown in 

figures 16 and 17.

and 32° upward (figs. 13, 15). This slower rate is about the 

same as spine 4 extrusion rates reported by Iverson and others (2006) and Major and others (this volume, chap. 12), who 

compared daily Sugar Bowl photographs. With time, extrusion rates diminished and directions became more easterly: 

3.9 m/d along a path of S. 39° E. and 23° upward between 

February 21 and March 10. This decreased to 2.6 m/d along a 

path of S. 71° E. and 6° upward between March 10 and April 

19. We estimated a time-averaged rate of advance for the 

period of January through mid-March by measuring advance 

of the leading edge of the whaleback in cross section (fig. 

16). Between January 3 and March 10, the advance of ~300 

m gave a rate between 4 and 5 m/d. Rates were as high as 8 

m/d during some briefer intervals (table 2).

The initial volume of spine 4 on January 3, 2005, as it 

splintered from spine 3, was 5.5×106 m3, and total volumes 

of hot rock reported for different times during the spine 4 

extrusion (Schilling and others, this volume, chap. 8) allow 

calculation of spine 4 volumes and time-averaged, volumetric extrusion rates between January and April 2005 (table 

2). Time-averaged extrusion rates for these intervals suggest 

a spurt in growth from February 1 to 21. The growth spurt 

occurred during the same interval in which spine 4 transitioned 

from intermittent thrusting to steady recumbent growth.

Deformation of Areas Surrounding Spine 4 

Growth during January–April 2005

Spine 4 growth during January through April caused 

areas to the east to rotate eastward about a pivot near Opus. 

Displacements were as great as 200 m (fig. 15). In response to 

motion of the welt to the south of Opus, ice masses stranded 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 191

1,900

2,100

2,300

2,500

4/19/05

3/10/05

2/21/05

2/1/05

1/3/05

12/11/04

2003

1986

1980

1,900

2,100

2,300

1,900

2,100

2,300

1,900

2,100

2,300

0 200 400 600 800 1,000

DISTANCE, IN METERS

ALTITUDE, IN METERS

January 3, 2005

February 21, 2005

April 19, 2005

North-northwest South-southeast

C C ́

Spine 4

Spine 1

Spine 4

Spine 4

Ice

Spine 5

Spine 3

Spine

 3

Deformed

1980–86

dome

1980s debris

Ice

Ice

1980–86

dome

1980–86

 dome

1980–86

 dome

1980s debris

1980s debris

Deformed

2004–5 dome

and older rock

spine 1

Deformed

1980–86

dome

Deformed

2004-5 dome

and older rock

Spine 1

Deformed

1980–86

dome

Deformed

2004–5 dome

and older rock

Deformed

2004–5 dome

and older rock

Spine 4, 2nd whaleback

Spine 1

Motion of spine 3 and ice

is chiefly to east

away from view

Motion of spine 3 and 

ice is chiefly to east

away from view

Motion of spine 4

is chiefly to east

and slows to south 

1980–86

dome

Deformed1980–86 dome

Deformed

2004–5 dome

and older rock

Deformed1980–86 dome

Deformed

2004–5 dome

and older rock

Deformed1980–86 dome

1,900

2,000

2,100

2,200

1,900

2,000

2,100

2,200

2,300

1,900

2,000

2,100

2,200

1,900

2,000

2,100

2,200

0 100 200 300 400 500 600 700

Ice

Ice Ice

Ice Ice

Ice

Ice

Ice

1980s debris

1980s debris

1980s debris

1980s debris

 Welt,

2004–5

debris

 Welt,

2004–5

debris

 Spine 3

whaleback

Spine 3

Spine 3

Spine 3

 Spine 4

whaleback

Spine

 2

Spine 4

Spine

 2

Spine

 2

Spine

 2

Spine

 4

Spine 3

remnant

Spine 3

remnant

Spine 3

remnant

Talus

 Welt, 

2004 debris

West East

 Welt,

 2004–5 

 debris

November 20, 2004

January 3, 2005

February 21, 2005

April 19, 2005

1986 surface

1980 surface

1986 surface

1980 surface

1986 surface

1980 surface

1986 surface

1980 surface

ALTITUDE, IN METERS

DISTANCE, IN METERS

E E ́

 Welt,

2004–5

debris

 Welt,

2004–5

 debris

 Welt,

 2004–5 

 debris

 Welt,

 2004–5 

 debris

Figure 16. North-northwest to south-southeast cross section 

C–C ́ (location shown in figure 15) illustrating extrusion and 

recumbent growth of whaleback, spine 4. Top panel shows known 

profiles for dates given, and other panels illustrate geologic 

interpretations on given dates. Dashed lines indicate inferred faults.

Figure 17. East-west cross section E–E ́ (location shown in 

figure 15) illustrating extrusion and recumbent growth of spines 

3 and 4. Panels illustrate geologic interpretations on given dates. 

Dashed lines indicate inferred faults; heavy and light line weights 

indicate contacts between spines and within spines, respectively.
192 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

on the welt moved eastward until they crossed into the realm 

of the active east Crater Glacier, whereupon they flowed 

northward as part of the glacier. Recognizable remnants of 

the disintegrating spine 3 shifted eastward 100 m or more by 

February 1. Areas closest to spine 4 moved fastest. Between 

January 3 and February 1, an area on spine 3 south of spine 

4 moved 107 m S. 79° E. at a rate of 3.7 m/d and subsided 

slightly (fig. 15). The GPS station HNY0, located ~40 m to 

the east, moved 3.2 m/d, S. 76° E., and also subsided (fig. 

13) (LaHusen and others, this volume, chap. 16). Continuing 

deformation at nearby sites was S. 85° E., 3.6 m/d, and S. 81° 

E., 1.9 m/d, during February 1–March 10 (fig. 15).

Spine 4 bulldozed and tilted spine 3 eastward, rapidly 

fracturing it and causing it to disintegrate (figs. 2, 3, 17). By 

April 19, spine 3 had been reduced to a rubbly ridge adjacent to the east Crater Glacier, and its surface area had been 

reduced by a factor of five (Herriott and others, this volume, 

chap. 10). Overthrusting caused by eastward rotation of spine 

4 as it deflected off the crater wall caused the reduction in area 

of spine 3 (fig. 17). Except for one small area, spine 3 remnants to the west were buried by spine 4 talus.

Areas on the 2004–5 dome west of spine 4 continued to 

subside but moved laterally no more than about 10 m. Spine 1 

subsided in December but showed no significant motion thereafter. An area near the south end of spine 2 and another on a remnant of spine 3 west of spine 4 moved less than 10 m southwest 

between January and February and subsided 10–20 m. Thereafter, subsidence continued at a slower pace, and no further 

translation of these spine 2 and 3 fragments was detectable.

As the east Crater Glacier continued to be bulldozed 

eastward at rates as high as 1.5 m/d, it rose tens of meters and 

accelerated downstream. Between mid-December 2004 and 

mid-April 2005, the glacier profile bulged as much as 90 m 

and became markedly convex (fig. 14B). During this interval, individual surface features on east Crater Glacier rose as 

much as 50 m, with the largest vertical displacements in areas 

east and southeast of Opus. In contrast, farther upslope to the 

south, features lost tens of meters in altitude as they flowed 

north (fig. 15). Features with eastward components of displacement in the autumn of 2004 shifted to due north displacement during winter 2005 as the glacier accelerated away from 

the area of constriction between the rotating dome complex 

and the east crater wall. The glacier had thickened so much 

since the onset of the eruption (as much as 130 m) that its 

slope had increased dramatically, and it responded by flowing 

north to correct the imbalance.

The Crater Glacier west and north of the spine complex 

responded by moving a few meters northwest during late 

December to April, in places rising by a few meters. Although 

dome growth may have had a minor effect on west Crater 

Glacier (fig. 15) during the winter months of 2005, accumulation of snow and normal glacier flow downslope probably 

accounted for most of the observed change. Areas of the 

glacier immediately southwest of the 2004–5 dome complex 

moved about 30 m northwest in response to the approach of 

spine 4 (figs. 12, 15).

Phase 4, Extrusion of Spine 5 across Previous 

Spine Debris: April 10–July 31, 2005

Spine 4 broke up, decoupled, and changed direction in 

mid-April 2005. Continual pressure caused by spine 4 pushing 

against the opposing crater wall disrupted its steady southward 

propagation and caused it to break apart. Repeat photographs 

(Major and others, this volume, chap. 12) and time-lapse photography (Iverson and others, 2006) showed development of 

northeast- to southwest-striking fractures cutting spine 4 at this 

time. These fractures became progressively more prominent 

from April 10 to 19. Between April 19 and 24, as the fractures 

grew, spine 5 began to thrust over the top of spine 4 remnants. 

Like spine 4, the spine 5 source was ~50 m southeast of spine 

1 (fig. 15).

Spine 5 had decoupled from spine 4 by April 19, though 

it continued to drag the southern parts of spine 4 southward 

until mid-May and to displace it to the east through June. 

Spine 5 displayed a smooth, gouge-covered surface near the 

vent. It became progressively steeper with time, the slope 

increasing from 40° on April 19 to 54° on July 14 (fig. 18). 

The spine tended to fracture and crumble as it grew higher, 

leading to a substantial breakup and decoupling of the 

southern section of the spine between mid-June and mid-July. 

During this interval spine growth began to resemble that of 

the Mont Pelée spine of 1903, with a steeply thrusting lithic 

core surrounded by an apron of debris at the angle of repose 

(Blake, 1990). The thrusting spine 5 acted as the driving force 

for a conveyer that transported at least half of the volume of 

the crumbling spine southward to form a ridge of disaggregated rock with a trend of S. 10° E. in mid-June and S. 5° E. in 

mid-July.

From April to August 2005, we were unable to use surface deformation vectors to make independent measurements 

of spine growth rate because of plumes that obscured key parts 

of the spine in aerial photos on two of four dates. A limiting, 

average, lineal growth rate of 4.3 m/d can be calculated from 

knowing that spine 5 was 100 m long on April 19 and 340 m 

long on June 15. Data from Major and others (this volume, 

chap. 12) suggest that extrusion rates of the smooth northern surface diminished from ~4 m/d in late June to ~2 m/d 

by early July. Time-averaged volume flux decreased slowly 

between April and July (table 2).

Deformation of Areas Surrounding Spine 5 from 

April to July 2005

Deformation of areas surrounding spine 5 showed that the 

spine was deep seated only near its source and was thrusting 

upward over previous parts of the dome complex to the south 

(fig. 19). Spine 4 remnants east of spine 5 rotated northeast or 

east, with the pivot point about 300 m southeast of the vent in 

an area of stagnant spine 4 rock (near point -7,-2,-3 in fig. 19). 

The areas of maximum displacement were in the north adjacent to the vent. Maximum displacement east of the vent on 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 193

8/10/05

7/14/05

6/15/05

4/19/05

2003

1986

1980

1,900

2,100

2,300

0 200 400 600 800 1,000

1,900

2,100

2,300

1,900

2,100

2,300

1,900

2,100

2,300

2,500

Deformed

2004–5 dome

and older rock

Deformed

2004–5 dome

and older rock

Deformed

2004–5 dome

and older rock

Spine 5

Spine 5

Spine 5

Ice

Ice

Ice

Spine 5

Debris

Debris

Debris

1980–

1986 

dome

1980–

1986 

dome

1980–

1986

dome

North South

ALTITUDE, IN METERS

DISTANCE, IN METERS

Deformed

1980–86 

dome

Deformed

1980–86 

dome

Deformed

1980–86 

dome

April 19, 2005

June 15, 2005

July 14, 2005

F F ́

Deformed 1980–86 dome

Deformed 1980–86 dome

Deformed 1980–86 dome

Crater 

wall

Crater 

wall

Crater 

wall

Figure 18. North-south cross section F–F ́(location shown 

in figure 19) illustrating extrusion and thrusting growth of 

spine 5. Top panel shows known profiles for dates given, and 

other panels illustrate geologic interpretations on given dates. 

Southern part of spine 5 crumbled and broke up between June 

15 and July 14, 2005.

spine 4 was 40–50 m, generally to the northeast. Most of the 

displacement occurred in April and May; very little occurred 

thereafter. A feature near GPS station SEV7 moved about 40 

m to the northeast and subsided 16 m (point -7, -3, -6, fig. 

19). All but 2 m of translation had occurred by June 15, but 

subsidence continued through August. The GPS station SEV7, 

located ~30 m west of this feature, was deployed on May 24. 

By June 15 it had moved only 7 m S. 70° E., implying that 

about 80 percent of the translation in this area had occurred 

between the time of the April 19 DEM and May 24. All remnants of spine 4, including those that did not translate significantly, subsided 10–20 m during spring and early summer of 

2005. Farther east, spine 3 remnants also subsided between 

April and June but did not translate more than ~5 m.

Areas on the dome west and north of spine 5 moved 

northwestward 40–60 m (fig. 19). Spine 1 was pushed ~50 m 

northwest (point +10, fig. 19). An isolated remnant of spine 

3 to the southwest of spine 5 (point -1, fig. 19) moved about 

40 m westward before it was buried. Spine 2 was buried by 

encroaching spine 5 talus in June and then could not be tracked, 

but its contact with the west Crater Glacier receded as much as 

100 m westward between April 19 and August 10 (fig. 19).

Nearby parts of the west Crater Glacier began to be 

pushed westward and were uplifted (fig. 19). Traceable features 

on the west glacier moved west as much as 80 m and rose as 

much as 20 m near the new dome’s southwest margin, but areas 

farther to the north and west merely flowed downslope. Station 

WES6, a GPS receiver located on the west glacier ~200 m west 

of the dome, moved 10 m N. 51° W. and rose 0.8 m between 

July 14 and August 10. Meanwhile, the east Crater Glacier 

flowed passively to the north. East Crater Glacier fractured 

and became greatly crevassed as it descended rapidly to the 

north in response to the 100 m of excess thickness it had gained 

between November 2004 and April 2005. No further deformation of the east glacier occurred after mid-May, and its boundary with the dome became fairly stable (fig. 19).

Phase 5, Spines 6 and 7 Extrude Westward: 

August 1, 2005–April 2006 (Ongoing)

Normal Faulting and Westward Growth of Spine 

6: August 1–October 9, 2005

Crumbling of spine 5 presaged reorganization of the vent 

and growth of spine 6 in late July and early August 2005. 

At least seven substantial rock avalanches and two slumping 

events between July 18 and 31 (table 3) reduced the smooth 

steep (54°) slabs of spine 5 of July 14 (fig. 3) to a rubbly ridge 

by August 10 (fig. 3). During this period, southward motion of 

segments of spine 5 slowed successively from south to north 

in such a way that, by the end of July, only the northernmost 

segment remained active.

During a transitional period August 1–5 (table 1), spine 

motion seen in time-lapse photography (Iverson and others, 

2006) became localized to the vent area, where the sense of 
194 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0 100 200 300 400 50

METERS

10/4/04 vent

8/10/05 features, glacier margin

7/14/05 features, glacier margin,

spines 3 and 4 remnants

6/15/05 features, glacier margin

Squares, immobile features

Circles, moving features

Numbers, vertical changes, in meters

Triangles, GPS stations

4/19/05 features, glacier margin

West

Crater Glacier

East

Crater Glacier

1980–86 dome

Opus

Spine

 5

+10

D

D'

D

F

F'

MID9.1

SEV7

Vectors: yellow, this study; 

black, GPS

HIE5.0

WES6

ELE4.1

ELE4.3

ELE4.2

562200 562400 562600 562800 563000 563200

5116400

5116000

5115600

5115800

5116200

-8,-9,-7

-13,-8,-6

-4,-2,-4

-7,-5,-5

-6,-1,-5

-7,-2,-3

-5,-2,-4

+15,-3,-3

-4,-2,0

+8, -1,+1

+2, -4,-3

-7, -2,-3

+6,+1, +5

+14,-9,+7

-1

-12

0, 0,+5 no change

-3, -1,-3

-8, -5,-2

-7, -3,-6

-8,-9,-7

 eg1

-13,-8,-6

-4,-2,-4

-7,-5,-5

-6,-1,-5

-7,-2,-3

-5,-2,-4

+15,-3,-3

-4,-2,0

 wg1

+8, -1,+1

+2, -4,-3

-7, -2,-3

+6,+1, +5

+14,-9,+7

-1

-12

0, 0,+5 no change

-3, -1,-3

-8, -5,-2

-7, -3,-6

Spine 3 

remnant Spine 4 

remnant

Spine 3 

remnant remnant

Figure 19. Deformation vectors for the period from April 19 to August 10, 2005, shown on DEM of July 14, 2005. Squares 

and dots indicate features tracked for dates identified by color in key. These colored symbols delimit deformation vectors 

of each of the three time intervals between relocation of features. Numbers are vertical components of vectors, in 

meters, for each interval. Underscores indicate no data for that interval. Dotted lines indicate dome-glacier margins for 

dates indicated by color in key. Line F–F ́ is trace of cross section shown in figure 18; line D–D ́ is trace of cross section 

shown in part in figure 22.

motion was nearly vertical, as the remainder of spine 5 began 

to slump slowly. A fracture system, somewhat concealed by 

rubble, developed along a S. 10° W. trend and divided the 

stagnating spine from the active extrusion as spine 6 developed 

and became the dominant feature, evident in images of August 

10, 2005 (fig. 3).

From early to mid-August, spine 6 began moving westward, slowly at first but then more rapidly. This westward 

motion of the extruding mass beginning on about August 6 

marked the completion of the transition from spine 5 to spine 

6 (table 1). Time-lapse photography of August 6–12 (Iverson 

and others, 2006) showed that spine 6 disintegrated continuously as it extended to the west.

The spine 6 mass comprised buried spines 1 and 2, a 

substantial part of spine 5 that had slumped to the west, debris 

shed from spines 3, 4, and 5, and welt rock caught between 

spines 1, 2, and 3 (fig. 20). This mass thus included massive lava, as well as deformed and disintegrating blocks and 

debris. On August 10, the extruded lava volume in spine 6 was 

~1×106 m3, but the total volume of the deformed mass was 

~12×106 m3 (table 2). A spurt in lava extrusion occurred during the transition from spine 5 to spine 6 and continued into 

September (table 2).

No specific feature on spine 6 could be tracked in DEMs 

and vertical aerial photos between August 10 and September 

20. However, the trace of the active spine and the shift of the 

west glacier-spine margin constrain the magnitude of translation during this period (fig. 21). The most active part of the 

spine migrated ~140 m N. 75° W. during the 41 days at an 

average rate of 3.4 m/d. Overall the most active part of the 

spine subsided a net ~40 m, but this value ignores vertical 

growth—active disintegration removed tens of meters from 

the apex of spine 6. The center of most active extrusion moved 

westward away from the original (October 4, 2004) vent area 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 195

Interval during which 

event occurred Phenomenon Direction

July 18–19 Avalanche East

July 19–20 Avalanche West

July 20–21 Avalanche West from near apex 

of smooth slab

July 22–23 Avalanche West

July 23–25 Slump West

July 26–27 Avalanche East

July 27–29 Slump West from rubbly area 

of south part of spine

July 28–29 Avalanche East

July 30–31 Avalanche West

Table 3. Timing, style, and direction of mass-wasting events from 

spine 5 at Mount St. Helens, Washington, July 19–31, 2005.

[Timing and style of phenomena are inferred from time-lapse video (Iverson 

and others, 2006).]

(fig. 21). Along an arc west to southwest of this new vent, the 

glacier-spine 6 contact receded 60–90 m westward at an average rate of 1.5–2.2 m/d (figs. 20–22).

Spine 5 Subsidence and Deformation of 

Surrounding Areas during August–September 

2005

All parts of spine 5 translated westward and subsided as 

spine 6 moved away and removed its westward buttress (figs. 

21, 22). Areas near spine 6 experienced maximum displacements of 40–60 m to the west and subsided by as much as 65 

m. Areas close to the crest of spine 5 dropped by as much as 

50 m while translating ~20 m westward. Such a pattern implies 

normal faulting along a northerly strike with westward dip as 

steep as 70°. The principal fault surface appears to coincide with 

a subsurface boundary sloping steeply to the west and demarcated by the contact between the subsurface remnants of spines 

4 and 5 (fig. 22). This boundary could have formed a weak 

surface that was susceptible to subsequent subsidence. Features 

on spine 5 farther west of the fault trace translated more and 

subsided less than those near it, a pattern that implies rotational 

motion along a flattening fault (fig. 21). We infer that the fault is 

listric, dipping steeply westward along its near-surface trace and 

flattening as it extends deeper to the west (fig. 22).

The perspective provided by south-rim time-lapse photography (Poland and others, this volume, chap. 11) suggests 

that southwestern parts of spine 5 and southeastern parts of 

spine 6 subsided and migrated westward in tandem. Spine 

5, near its eastern margin, subsided very little, apparently 

because the underlying spine 4–5 contact dipped gently compared to the steeper contact to the west. The most significant 

translation had occurred by September 20, but some areas 

continued to subside through October. Subsidence of spine 5, 

coupled with westward extension of spine 6, resulted in formation of a sag between spines 5 and 6 (figs. 21, 22).

To the east, Opus, spine 3, and spine 4 were relatively 

immobile, moving less than 4 m. In August and September, 

the GPS station SEV7, situated on spine 4 about 30 m to the 

east of spine 5, translated 2 m westward and subsided 3 m as 

the spine 5 buttress gave way. Motion on other parts of Opus, 

spine 3, and spine 4 was too small to be detectable (fig. 21).

While west Crater Glacier accelerated westward to 

northwestward and thickened, east Crater Glacier continued to flow passively northward (fig. 21). Between August 

10 and October 24, traceable features on the west glacier 

moved west by 100–120 m and rose 20–35 m in response to 

the bulldozing caused by spine 6 advance (fig. 22). Ensuing 

crevasses radiated westward along the principal strain axis. 

Three GPS stations were located on west Crater Glacier 

for various time intervals (Walder and others, this volume, 

chap. 13). An example, WES6, originally about 150 m west 

of the dome, moved 50 m N. 51 W. and rose 13 m during 

August 10–September 14. In contrast, its motion in the 27 

days before August 10 was only 10 m northwest and 1 m up. 

Advance of spine 6 pushed west Crater Glacier westward and 

heaped it as high as 30 m above its previous surface.

Westward Extrusion and Overthrusting of Spine 

7: October 9, 2005–April 2006

Extrusion of spine 7 began in mid-October with subsurface spine intrusion centered east of spine 6 and near the trace 

of the October 4, 2004, vent. Subsurface intrusion gave way 

by November to spine extrusion, which continued through 

April 2006 (table 1, fig. 21B).

An increase in high-frequency earthquakes, beginning 

on October 9, 2005, marked the beginning of spine 7 growth 

(Moran and others, this volume, chap. 2). Fuming and heating of the surface area above the October 2004 vent observed 

in thermal images of October 11, 2005, revealed the first 

surface manifestation of the new intrusion (fig. 3). Time-lapse 

photographs from a camera on the south crater rim (Poland 

and others, this volume, chap. 11) showed general bulging of a 

rubbly area between spines 5 and 6 by October 13.

Uplift and westward motion that became increasingly 

evident between October 14 and 21 indicated extrusion of 

spine 7. From its origin in the depression between spines 

5 and 6, spine 7 pushed upward and outward to the west, 

steepening on the east as it grew and overthrusting spine 6 

to the west (fig. 20). By mid-November, a broken-up slabby 

spine had begun to emerge from the rubble-strewn slopes of 

the bulge. By mid-December, this slab of rock had become 

more coherent and prominent. As it continued to grow, the 

slab steepened eastward progressively, attaining a slope of 

50° by April 2006. The extruding slab also pushed spine 6 

and part of spine 7 across a sector extending from southwest 
196 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1,900

2,100

2,300

1,900

2,100

2,300

2,500

1,900

2,100

2,300

0 200 400 600 800 1,000 1,200 1,400 1,600 1,800 2,000

2,100

2,300

1,900

2,100

2,300

1,900

2,100

2,300

1,900

Spine 7

Opus Ice

Ice

Ice

Ice

Ice

Ice

Opus

Opus

Spine 6

Spine 6

Spine 6

Spine 7

Spine 7

Debris

Debris Debris

Debris

Debris Debris

DISTANCE, IN METERS

ALTITUDE, IN METERS

Ice Ice

Opus

Spine 6

Debris Debris

Spine 6

Ice Ice

Opus

Debris Debris

Deformed

dome rock

Deformed

dome rock

Deformed

dome rock

Deformed

dome rock

Deformed

dome rock

Deformed

dome rock

Deformed

dome rock

2/9/06

12/15/05

10/24/05

1986

1980

9/20/05

8/10/05

2003

West East

August 10, 2005

Crater

wall

Crater

wall

September 20, 2005

October 24, 2005

December 15, 2005

February 9, 2006

Deformed

1985 lobe

 Opus

1985 lobe

Deformed

spine 3-4

remnant

Spines 1, 2, and 3

inferred 11/4/04

A A ́

Deformed

 rock

Crater

wall

Crater

wall

Crater

wall

Crater

wall

Crater

wall

Crater

wall

Crater

wall

Crater

wall

Figure 20. East-west cross section A–A ́ (location shown in figures 1 and 21) 

illustrating extrusion and thrusting of spines 6 and 7. Top panel shows known profiles 

for dates given, and the other panels illustrate geologic interpretations at labeled 

times. Dashed lines indicate thrust faults.
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 197

12/15/05 features, glacier margin

10/24/05 features, glacier margin

0 100 200 300 400 500

METERS

Squares, immobile features

Circles, moving features

Numbers, vertical changes, in meters

8/10/05 features, glacier margin

9/20/05 features, glacier margin

10/4/04 vent

2/9/06 features, glacier margin

GPS stations

Vectors: black, GPS; yellow, this study

+6,+12,+6,_

+12,+17,_,_

+23,+13,_,_

 wg1 

-41,-40,-13,_

-43,-30,-31,-16

_,-30,-31,-14

_,-23,-33,_

no change

Opus

no change

no change

1980–86 dome

-11,-8, 0,_

eg1

0,-2, -2, -2

-49

-39

-20, 0, 0,_

-24, -1, 0, 0

-65,-16,+3,+19

-53,-16,+1,-2

-20,-7,-3,-5

West

Crater

Glacier

East

Crater

Glacier

Spine 3

Spine 5

Spine 4

ELEA4.4

HIE5.1 WES6

SEV7

ELE4.3

HIE5.0

Spine 6

A

A'

D

-12

12/15/05 features, glacier margin

0 100 200 300 400 500

METERS

Squares, immobile features

Circles, moving features

Numbers, vertical changes

2/9/06 features, glacier margin

GPS stations

Vectors: black, GPS; yellow, this study

+4

-16

-14

no change

Opus

no change

no change

1980–86 dome

 -2

no change +19

-2

-5

West

Crater

Glacier

East

Crater

Glacier

Spine 3

Spine 5 Spine 4

SEV7

A

A'

D

2.2 m/d

50 o

 up 

10/4/04 vent

Spine 7

Spine 6

562200 562400 562600 562800 563000 563200

5116400

5116000

5115600

5115800

5116200

5115400

562000 563400

562200 562400 562600 562800 563000 563200

5116400

5116000

5115600

5115800

5116200

5115400

562000 563400

A

B

GND0

1 m/d

2° up

GND0

1 m/d

D'

D'

Figure 21. Deformation 

vectors for the period from 

August 10, 2005, to February 

9, 2006, shown on DEMs. 

Squares and dots indicate 

features tracked for dates 

identified by color in key. 

These colored symbols 

delimit deformation vectors 

of each of the time intervals 

between relocation of 

features. Numbers are 

vertical components of 

vectors, in meters, for 

each interval; underscores 

indicate no data for interval. 

Red-lined pattern indicates 

spine 6; green-lined pattern 

indicates spine 7. Dotted 

lines indicate domeglacier margins for dates 

indicated by color in key. 

Lines A–A ́ and D–D ́ are 

traces of cross sections 

shown in figures 20 and 22. 

A, September 20, 2005. B, 

February 9, 2006.
198 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1,900

2,100

2,300

2,500

1,900

2,100

2,300

0 200 400 600 800 1,000 1,200 1,400 1,600 1,800

DISTANCE, IN METERS

ALTITUDE, IN METERS

West East

9/20/05

8/10/05

7/14/05

4/19/05

 1986

 1980

2/9/06

12/15/05

10/24/05

 1986

 1980

10/24/05

Ice

Spine 4

Ice

Ice

Ice

Spine 4

Spine 7

Spine 6

Spine 5

Spine 3

Spine 3

Spine 7

Spine 6

Spine 5

Spine 7

 Normal

 fault 

8/5/05–9/20/05

Debris

Debris

D D ́

2004–6 dome rocks

Figure 22. East-west cross section D–D ́ (location shown in figures 1 and 21) illustrating extrusion and thrusting of spines 6 

and 7. Panels illustrate geologic interpretations at labeled times. Dashed line indicates inferred normal fault, which is located 

at steep west-dipping trace of spine 4 on April 19, 2005, surface.

to northwest radially away from its source (fig. 21). Rock 

debris continually avalanching from the west face of spine 

7 formed hot talus slopes on this side. As the spine grew 

through November, its rubbly western slope began to bury 

adjacent sections of spine 6.

Spine 7 formed a distinct entity between, and overlapping, spines 5 and spine 6. By October 24, the spine had been 

extruding for 10 days, and we estimate its volume as about 

a third of the total volume erupted since September 20, or 

~1×106 m3. Between October 24, 2005, and February 9, 2006, 

time-averaged magma flux gradually diminished (table 2).

As spine 7 moved westward and thrust over spine 6, it 

grew higher and steeper, and its solid eastern buttress became a 

progressively more prominent, finlike structure (figs. 23A, 23B). 

Cracks penetrated the gouge coat in distinctive patterns and 

moved upward and westward along with the fin (fig. 23A). From 

examination of photographs, spine 7 was ascending westward 

at an angle of 50° at a rate of ~2 m/d in early April (fig. 23A). 

At the same time, GPS station GND0, ~100 m west of the fin, 

was only moving 1 m/d horizontally westward (fig. 23C). This 

discrepancy in rate of deformation within spine 7, with steeper, 

faster displacement near source and slower subhorizontal displacement to the west, implies internal shearing (figs. 20, 23D).

Deformation during Spine 7 Extrusion: October 9, 

2005–April 2006

All parts of spine 6 translated and subsided as spine 7 

pushed it westward (figs. 21, 22). Features on spine 6 were 

subject to substantial but gradually diminishing deformation 

from October 9, 2005, to February 9, 2006, with maximum 

total displacement over this period in excess of 200 m to 

the west and subsidence as much as 80 m (table 4). Toward 

the end of this period, spine 7 moved about 1 m/d westward 

with no subsidence, while adjacent spine 6 moved at half that 

rate and subsided (table 4). Such a pattern implies shearing 

between spines 6 and 7 (figs. 20, 23D).

West Crater Glacier continued to move westward to 

northwestward (fig. 21). Between October and December, a 

single traceable feature on the west glacier moved northwest 60 

m and rose 6 m in response to continuing spine impingement 

(fig. 21). Motion of GPS station ELE4.4, positioned farther 

south and closer to spine 7, slowed after October 24. It moved 

1.6 m/d N. 68° W. and rose 19 m in the 34 days before October 

24, whereas it moved 1.1 m/d N. 55° W. and rose 3 m in the 15 

days after that date (LaHusen and others, this volume, chap. 16). 

Farther north, the glacier accelerated to the north in response 
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 199

Ground contact

of 3/19/06

Ground contact

of 2/24/06

Ground contact

A of 2/15/06

B

Spine 6

Spine 7

Thrust faults

~2 m/d 50° up and West

~1 m/d to West

Spine 6

C

Spine 7

Talus,

 cold Talus,

 cold

Spine 6

Spine 7

Spine 5

West East

Hot

Hot

Hot

DTHGIEH

m004~

~600 m

DISTANCE

~200 m

Glacier

Ground contact

of 4/4/06

48 days

~ 2m/d

~ 100 m

48 days

~ 2m/d

~ 100 m

Figure 23. Photographs, thermal infrared image, and cross 

section of the Mount St. Helens crater on April 4, 2006. A, View 

looking west at 100-m fin showing previous ground-spine contacts. 

B, View looking northeast at spines 6 and 7. C, Thermal infrared 

image of view in B, looking northeast and showing thrust faults 

and relative motions. D, Schematic east-west cross section of 

growing spine 7, slowing spine 6, and stagnant spine 5. Orientation 

is roughly along line shown in B but extends east to spine 5. USGS 

photos by J.W. Vallance; thermal image by M. Logan.

to the 30–40 m of excess elevation it had gained through uplift 

between August 2005 and February 2006 (fig. 21A).

Discussion

We consider here factors that influenced dome growth 

during the 2004–6 eruption. Potential near-surface controls on 

spine growth during the 2004–6 eruption include thick glacial 

ice, initial vent position and geometry, the 1986 topographic 

surface, and backpressure caused by spines pushing through 

and thrusting over debris from previous spines. We also consider the effects of dome growth on Crater Glacier. Lastly, we 

compare the 2004–6 Mount St. Helens dome-building eruption 

with well-documented historical examples at other volcanoes.

Effect of Glacier on Spine Growth

Glacier ice as thick as 150 m has apparently had little 

effect on the extrusion of the dome or on the growth of various spines, except to conceal substantial parts of them and 

to prevent shedding of disintegrating dome talus beneath the 

level of the glacier surface. As discussed in the introductory 

section, we infer that dome-glacier contacts have remained 

steep. Near-vertical contacts are consistent with ice-hot rock 

marginal boundaries observed at other locations. Examples 

include tuyas in British Columbia, Canada (Mathews, 1947), 

and ice-lava contacts at Mount Rainier (Lescinsky and Sisson, 

1998). Glacial ice appeared not to impede spine growth significantly. Spines 2–4 grew recumbently to the south, pushing 

glacial ice aside as they progressed. Westward extension of 

spines 6 and 7 also pushed through thick glacial ice. The bed 

of the glacier was permeable enough that meltwater drained 

away without interacting with hot dome rock (Walder and 

others, this volume, chap. 13), except possibly during six brief 

phreatic explosions (Moran and others, this volume, chap. 6).

Vent Dimension and Location and Influence of 

1980–86 Dome on Spine Growth

The depression from which the initial phreatic eruptions originated and from which the initial spine extruded was 

located at the west end of Opus and had an approximate 
200 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Interval 9/20/2005–

10/24/2005

10/24/2005–

12/15/2005

12/15/2005–

2/9/2006

Days elapsed 34 52 56

Bearing N80°W N80°W N85°W

Horizontal translation

Displacement (m) 110 75 35

Rate (m/d) 3.2 1.4 0.6

Subsidence

Displacement (m) 30 31 14

Rate (m/d) 0.9 0.6 0.3

Table 4. Timing, bearing, magnitude, and rate of deformation of 

one feature located on spine 6 at Mount St. Helens, Washington, 

between September 20, 2005, and February 9, 2006.

[Timing, direction, and magnitude of deformation are inferred from DEMs on 

given dates (Schilling and others, this volume, chap. 8).]

diameter of 120 m (figs. 5, 12). Superimposing this vent location on the 1986 topographic surface (fig. 12) reveals that the 

initial vent was located on ice over the south-facing slope of 

the 1980–86 dome, in particular over the 1985 fault-formed 

ridge known as Opus. Just before the 2004 eruption, relief 

from the high point on Opus to the moat’s floor beneath the 

glacier was ~130 m, and south-facing slopes were as steep as 

40° under the trace of the vent (fig. 12).

The initial location of the 2004 vent, on steep south 

slopes of Opus, themselves buried beneath the glacier, clearly 

influenced the propensity of stiff spine extrusions to move 

southward. Several studies suggest that the magma ascending 

the conduit had largely solidified within a kilometer of the 

surface (Dzurisin and others, 2005; Iverson and others, 2006; 

and in this volume: Moran and others, chap. 2; Cashman and 

others, chap. 19; Pallister and others, chap. 30). We infer that, 

as the magma neared the surface in September 2004, the preexistent solid plug of 1980–86 dome rock deflected the rising 

mass southward so that, near the surface, it tilted southward.

We tracked successive extrusion points at the surface 

by centering circles on the center of each spine origin with 

appropriate diameters equal to spine width (fig. 24). As the 

extrusion transitioned from the initial spine formation of 

phase 2 to well-developed whaleback-style spines of phase 

3, the surface trace of the vent moved south and east. These 

trends continued until spine 3 began its breakup in December 

2004. During the spine 3–4 transition this sense of movement 

reversed. However, during growth of spine 4, the southward 

and eastward motion recommenced. With extrusion of spine 

5 during phase 4, the vent gradually moved westward back 

toward its original location. This motion increased with phase 

5 extrusion of spine 6. With the onset of spine 7 extrusion, the 

surface manifestation of the vent returned to and remained 

within 20 m of its original position. Between vent clearing 

of phase 1 and initial whaleback-style intrusion of phase 3, 

the surface trace of the vent rapidly rose 130 m because an 

increasingly thick pile of lava built beneath it. Elevation of the 

vent trace increased slowly through extrusion of spines 4 and 5 

during phases 3 and 4, then diminished with extension of spine 

6 as phase 5 commenced. Extrusion of spine 7 rebuilt the lava 

pile beneath the vent, so that by early 2006 its trace was 200 

-100

0

100

200

300

Easting

Northing

Vertical

0

100

200

300

400

10/2/04 11/21/04 1/10/05 3/1/05 4/20/05 6/9/05 7/29/05 9/17/05 11/6/05 12/26/05 2/14/06

DATE

Length

Height

Flux

DIMENSION, IN METERS POSITION OF VENT, IN METERS,

 RELATIVE TO 10/4/04 VENT

0

FLUX, IN CUBIC METERS PER SECOND

2

4

6

8

Spine 3 Spine 4 Spine 5 Spine 6 Spine 7

Spines 1 & 2 Phase 2

Spine 4 Spine 5 Spine 7 Spine 3 Spine 6

Spines 1 & 2

Onset of

recumbent

growth

Steady recumbent 

growth reestablished,

growth spurt Onset of westward

recumbent growth,

growth spurt

A Vent location

B Spine dimensions and extrusion rate

Spine 4

thrusts

over

spine 3

Spine 5

thrusts over

older spines

Spine 7 thrusts

over spine 6

Phase 3, recumbent growth Phase 4, overthrusting Phase 5, 

westward growth

Phase 1, vent clearing

Figure 24. Plots 

of positions of vent, 

dimensions of spines, and 

extrusion flux versus time. 

A, Center of vent located 

in plan. B, Length of active 

spine, height of active 

spine, and lava discharge.

m higher than its altitude in 

October 2004, though it was 

back to its original position in 

plan view.

Superimposing the 

position of the original vent 

and the seven spines on the 

1986 surface suggests that 

the vent or conduit need not 

have shifted substantially at 

the depth of the initial erup-
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 201

tion surface (figs. 12, 25). Geometry of the vent area permits 

origin of spines 1 through 4 from approximately the same 

initial vent location at the depth of the 1986 surface (fig. 12). 

Initially, extrusion of spines 1 and 2 covered the area directly 

above and to the south of the vent so that, as spine 3 began to 

grow upward and southward, previous spine and welt rubble 

diverted it eastward (fig. 12A). Additional dome rock and 

debris emplaced during spine 3 growth diverted spine 4 farther 

eastward. Eventually, so much dome rock had been pushed to 

the east that fragmented dome rock abutted the steep part of 

the east crater wall, and continuing spine extrusion could no 

longer push it aside. The subsequent spine (spine 5) therefore thrust instead over the older spines in a more southerly 

direction. When spine debris above the vent eventually built 

high enough, spine growth could extend westward across the 

rock debris of spines 1 and 2 (fig. 25). An inclined conduit did 

not simply increase in altitude; rather it shifted tens of meters 

southward, then south-southeastward, and finally southwestward as it evolved and grew higher (fig. 25). We suggest that 

vent geometry is such that spine extrusion throughout the 

eruption could have passed through approximately the same 

vent location at a depth near that of the 1986 surface and that 

migration of the vent’s surface expression resulted from diversion by remnants of previous spines.

Influence of 1980–86 Surface on Spine Growth

Topography inherited from the 1980–86 eruption controlled growth patterns of the laterally propagating spines 

(figs. 12, 25). We superimposed the outlines of actively 

growing spines on topographic features concealed by glacier 

ice to assess their influence on growth patterns. Generally, we 

found that the 1986 topographic surface controlled recumbent 

Figure 25. August 10, 2005, DEM showing outlines of spines 6 and 7 and their dome-glacier margins (solid lines) on 

dates indicated by colors in key. Dashed lines indicate contacts between spines 6 and 7 for appropriate dates indicated 

by colors in key. Select 20-m contours from 1986 DEM illustrate how composite spines 6 and 7 have migrated along a 

low topographic trough to west. Comparisons of altitudes on August 10, 2005, with those of 1986 are given for localities 

marked by red triangles.

(2160, 1967)

(2135, 1988) (2118, 2060)

(2118, 2100)

(2082, 1980)

(2220, 2020)

(2121, 2020)

(2281, 2035)

(2142, 2000)

2100 m

2000 m

2000 m

2000 m

2000 m

2100 m

2100 m

0 100 200 300 400 500 50

METERS

12/15/05 spine 6-7 margin, contact

10/24/05 spine 6-7 margin, contact

9/1/05 spine 6 margin

8/10/05 spine 6 margin

9/20/05 spine 6 margin

8/10/05 vent

Altitude, in meters (8/10/05, 1986)

2/9/06 spine 6-7 margin, contact 

1986 20-m contours

562200 562400 562600 562800 563000 563200

5116000

5115600

5115800

5116200

5115400

562000 563400

5116400

5116600

Crater

Glacier 1980–86 dome

Crater

Glacier

10/4/04 vent
202 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

growth directions, the rotation of spines, and how spines ultimately fragmented, but the ultimate barrier to continued lateral 

growth proved to be the walls of the 1980 crater modified by 

talus at their base and modified slightly by subsequent erosion.

After spine 1 grew, subsequent spines grew southward, 

in part because the crater floor sloped to the south. Spine 

2 advanced due south until it encountered the steep opposing slopes of the south crater wall. Because spine 2 was then 

positioned directly south of the vent, the prominent whaleback 

of spine 3 that followed was forced slightly eastward as it 

advanced. Spine 3 extended across a broad basin to the southsoutheast that was filled with glacial ice and welt debris. It then 

encountered the crater wall at an oblique angle (20°–30° from 

perpendicular) in mid-November 2004, and the influence of the 

wall deflected the snout of the spine eastward (fig. 12A). Once 

spine 4 had pushed remnants of its predecessor aside, it too 

progressed south-southeastward, encountered the 1980 crater 

wall at an oblique angle, and then rotated eastward (figs. 7, 15). 

A lack of such rotation before the spine arrived at buttressing 

slopes and fairly rapid rotation subsequently is strong evidence 

in support of the oblique-incidence hypothesis of rotation.

Both spines 3 and 4 began to fracture, crumble, and 

ultimately decouple from the source as a result of resistance to 

motion when they impinged on the crater wall. Spine 3 began 

to slow and break apart as it pushed against the crater walls 

in mid-November 2004. Two voluminous blocks and numerous smaller ones separated from the spine before it finally 

broke and fractured near its source on or about December 18, 

2004. Thereafter, lava near the source slowly decoupled from 

remnants to the south to form spine 4. A similar sequence 

occurred when spine 4 itself met the wall, fractured at its root, 

and decoupled in April 2005 to form spine 5. In the April case, 

however, debris had filled the area east as far as the east crater 

wall, and thus spine 5 grew by southward thrusting across 

previous spine remnants.

Rotational motion of spines 3 and 4 in turn caused rotation of previously emplaced dome debris, welt debris, and glacier ice to the east. Southward spine propagation and oblique 

impingement on the crater wall apparently caused the rotation, 

and the less resistant expanse of glacial ice to the east permitted the rotation to proceed in that direction. By mid-April, 

when spine 4 and associated debris outboard of it had encountered steep slopes to the southeast as well as to the south, the 

counterclockwise rotation (map view) of phase 3 ceased (fig. 

12B). Also at this time, the sense of rotation reversed from 

counterclockwise to clockwise, with a pivot 300 m south of 

the vent (fig. 19). The south end of spine 4 became fixed and 

the north end began to move east because the only available 

space east of the 2004–5 dome was located directly east of 

the vent (labeled “mid-April space east of vent” in fig. 12); no 

such space was available to the southeast.

Spine 5 thrust over remnants of spines 3 and 4 at steepening angles (fig. 18) until its perch atop those remnants became 

unstable. Failure along a north-south zone of weakness dipping 

steeply to the west allowed motion to resume along a 1986 

topographic low, as spine 5 slumped together with westward 

growth of spine 6 (fig. 22). From April to August 2005, preexisting topography had little effect on spine growth because the 

active spine, 5, was shearing over previous spines rather than 

following old topographic surfaces. With extrusion of spine 6, 

westward spine migration pushed previous spine remnants and 

affiliated rubble westward into the topographic trough defined 

by the 1980–86 dome and the crater wall (fig. 25). This moatlike topography channeled growth of spine 6 such that it barely 

impinged on steeper slopes to the south and rode up on topographically high areas of the 1980–86 dome only near the vent, 

where some northward spreading and rockfall was underway. 

Similarly, as spine 7 thrust west into parts of spine 6, it pushed 

the earlier spine westward along the same topographic trough. 

As the volume of material to the west built, the rate of westward recession of the dome-glacier boundary slowed, and slabs 

of spine 7 extruded at steepening angles.

Extrinsic Control of Spine Growth Rate

An intrinsic exponential decline in overall extrusion rate 

(fig. 24) that probably derives from declining magma supply 

and pressurization is apparent during the course of the present eruption. Overprinted on this decline are several apparent 

increases in magma flux that may have been controlled extrinsically (spurts in fig. 24B). Time-averaged effusion rates commonly rise rapidly to a peak before falling slowly, resulting in 

an exponential decrease in eruption rate and declining growth 

(Harris and others, 2000). Such trends can be explained by the 

tapping of enclosed and pressurized magma chambers (Wadge, 

1981; Harris and others, 2000). Extrinsic factors, such as 

changes in load, are thought to cause variation in effusion rates 

(Harris and others, 2003).

We hypothesized that increases or decreases in load near 

the surface might reduce or enhance extrusion rates, owing to 

increases and decreases in the mass displaced (fig. 24B). We 

plotted both relative height and length of active spines against 

time to test this idea. However, our results show no obvious 

correlation between spine length or height and extrusion rate 

(fig. 24B). A comparison between extrusion rate and style of 

extrusion suggests a possible correlation. Extrusion rate was 

greater when steady recumbent growth was established and 

smaller when active spines were thrusting upward at significant angles (fig. 24B). Steady lateral extrusion of spine 3 in 

late October and early November 2004 corresponded to an 

initial increase in discharge, and renewed steady extrusion of 

spine 4 in February 2005 corresponded to a slight increase in 

extrusion rate. A transition from thrusting to westward migration and slumping correlates to a third localized peak from 

late July to September 2005. Periods of resistance to movement caused by thrusting of spines at increasing angles over 

previous remnants also correlate with periods of diminished 

extrusion rate (fig. 24B). We suggest that such growth conditions may have acted to resist extrusion, suppressing the flux 

by backpressure. Slumping events and vent reorganization 

eased backpressure and thus enhanced flux.
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 203

Impact of Dome Growth on the Crater Glacier

Whereas the glacier had little affect on dome growth, 

spine growth did have a profound impact on the Crater 

Glacier—slicing it in two, pushing it hundreds of meters first 

one way then another, doubling it in thickness, but not melting 

it. As the response of Crater Glacier is the detailed subject of 

another contribution to this volume (Walder and others, chap. 

13), we merely summarize the impact of spine growth on the 

glacier from the perspective of surface deformation vectors, 

which allowed us to track certain glacier features throughout nearly the entire course of the eruption. During October 

and November 2004, subsurface deformation owing to spine 

extrusion caused the glacier surface to take on first the appearance of a migrating wave of fractured ice and rock, then, with 

surfacing of spine 3 through that material, the appearance of 

bow waves of fractured snow and ice both west and east of the 

whaleback form. Once spine 3 divided the glacier in December 2004, it and its successor, spine 4, began slewing to the 

east, rotating about their tails and plowing the ice of the east 

glacier into a 100-m-high berm by January 2005. The berm 

then sluiced northward through the gap between the 1980–86 

dome and the crater wall between April 2005 and February 

2006 (fig. 26). Growth of spines 6 and 7, plus subsidence of 

dome remnants into west Crater Glacier, created a similar 

response between August and December 2005 whereby the 

glacier was first pushed up and westward (fig. 26) and then 

began to flow through the gap between the 1980–86 dome and 

crater wall. Despite its mistreatment, the glacier has lost no 

more than about 10 percent of its volume to contact melting as 

of February 2006 (Walder and others, this volume, chap. 13). 

Apparently, gouge-coated spines and shrouds of cold debris 

have effectively insulated glacier ice from hot spine interiors.

Comparison with Other Dome-Building 

Eruptions

Several factors set the 2004–6 dome-building eruption of 

Mount St. Helens apart from that of other well-documented 

historical domes such as those of Mount St. Helens 1980–86, 

Montserrat 1995–98, Santiaguito 1922–2006, and Unzen 

1990–96. These include extrusion of solid spines, a propensity 

to form recumbent spines, interaction with glacier ice, and 

topographic setting. Mont Pelée in 1903 produced a spine that 

grew vertically but otherwise resembled whalebacks at Mount 

St. Helens. However, perhaps the historical dome-building 

eruption most similar to that of Mount St. Helens 2004–6 was 

the 1944–45 extrusion of the Showa-Shinzan dome at Usu 

volcano in Japan.

Mount St. Helens’ 1980–86 dome extrusion differed 

from the present extrusion in extrusion rate, morphology, and 

process of emplacement. Excepting a one-year endogenous 

phase, extrusion of the 1980–86 dome occurred in discrete episodes, 16 of which were preceded by periods of accelerating 

endogenous growth, followed by extrusion of a viscous lobe, 

and terminated with periods of subsidence and lateral spreading (Swanson and Holcomb, 1990). In contrast, the present 

eruption has proceeded continuously with a general decline in 

discharge (fig. 24). Output of fresh dacite in the early phases 

of the present eruption occurred at a rate that is about one-half 

to one-third of rates measured during the 1980–86 episodes 

(Chadwick and others, 1988). The relentless growth during the 

2004–6 eruption, however, has produced a total volume similar 

to that of the 1980–86 eruption in about one-fourth of the time.

Swanson and Holcomb (1990) document distinctive 

profiles for individual lobes emplaced during 1980–86, showing that lobes tended to adopt a characteristic slope (33°) and 

a characteristic height-to-diameter ratio (h/d) of about 0.32. 

The slope was approximately the angle of repose for coarse 

angular talus. In a more elaborate analysis, Iverson (1990) was 

able to model the characteristic slope in terms of a pressurized viscous magma enclosed by a brittle shell. Such a model 

is not applicable for spines of the present eruption because of 

the complicating influence of topographic barriers, glacial ice, 

and the subsolidus character of the magma, with the spines 

extruding in a near-solid state. Glacial ice within the crater 

has buttressed spines at slopes much greater than the angle of 

repose during the current eruption. Overall, the h/d ratio and 

slope for spines 3–7 ranged from 0.6 to 0.7 and from 50° to 

60°, respectively. Those of initial spines were greater. Slopes 

higher than the angle of repose also reflect the massive and 

solid character of extruded spines. Such high h/d ratios are 

typical of upheaved domes and peléean spines (Blake, 1990).

Well-documented dome extrusion at Unzen, Japan, and 

Santiaguito, Guatemala, was continuous but varied in extru-

-200

-100

0

100

200

300

10/2/04 12/1/04 1/30/05 3/31/05 5/30/05 7/29/05 9/27/05

Easting, eg1

Northing, eg1

Vertical, eg1

Easting, wg1

Northing, wg1

Vertical, wg1

DISPLACEMENT, IN METERS

DATE

East Crater Glacier

West Crater Glacier

Figure 26. Displacement plots of a feature on east Crater Glacier 

(eg1 located at site a in fig. 2) and one on west Crater Glacier 

(wg1 located at site i in figs. 2A, 2B) versus time (both features are 

located at intervals in figs. 5, 7, 15, 19, 21).
204 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

sion rate and included both endogenous intrusion of viscous 

magma and extrusion of lava flows (Rose, 1980, 1987; Nakada 

and others, 1999; Harris and others, 2003). At Unzen, endogenous growth typified slow discharge, and exogenous growth 

typified more rapid discharge (Nakada and others, 1999). 

Over the course of the eruption, discharge slowed and endogenous growth increased proportionately (Nakada and others, 

1999). Although Santiaguito has erupted continuously during 

1922–2006, its growth has been episodic—waxing and waning over time scales of several years (Rose, 1987). In contrast 

to recent activity at Unzen, during Santiaguito’s 84-year and 

ongoing eruption there has been a general tendency for the 

proportion of exogenous to endogenous growth to increase 

with time (Harris and others, 2003). Neither Unzen nor Santiaguito has shown the propensity to build solid-state spines at 

a low extrusion rate as observed during the current eruption at 

Mount St. Helens.

At Soufrière Hills volcano, Montserrat, Watts and others 

(2002) documented the morphology of lobes and spines, some 

of which superficially resemble those described here, and 

correlated them with extrusion rates. Watts and others (2002, 

their fig. 33) documented near-vertical spines, whaleback 

spines, and mega-spines that superficially resemble spines 

formed during the current eruption of Mount St. Helens, 

though similar forms here have had dimensions on a scale of 

hundreds of meters rather than tens of meters as observed at 

Soufrière Hills. With the exception of spine 1, the spines of 

the Mount St. Helens eruption have been larger and remained 

active longer than those at Soufrière Hills. Whaleback spines 

3 and 4 grew during periods of months rather than days as 

at Montserrat, and spines 5 and 7 thrust over previous spines 

during periods of four months or more. The unusual, simultaneous slump and westward extrusion of spine 6 has no analogy 

to spine growth during any documented episode at Soufrière 

Hills volcano. More fluid morphologies, such as shear lobes 

and pancake lobes, did not occur during the Mount St. Helens 

2004–6 eruption.

After its notorious eruption of 1902, Mont Pelée, Martinique, built a vertical spine that shares some characteristics 

with spines of the current Mount St. Helens eruption. That 

spine grew vertically to a height of more than 200 m in 1903 

(Lacroix, 1904). Photographs (Lacroix, 1904) suggest that 

the spine was solid, had a gouge-coated and striated surface, 

and exposed a broken and massive surface on its opposite 

side. Like the Mount St. Helens spines, the Mont Pelée spine 

crumbled as it grew and eventually stagnated (Lacroix, 1904).

During its 1943–45 eruption, Showa-Shinzan dome of 

Mount Usu uplifted an area of as much as 1.5×1.5 km as 

much as 140 m (Mimatsu, 1995), generating a deformed 

zone reminiscent of the welt. Extrusion followed the 

deformation, as a jagged solid spine punched through the 

older roof rocks. As described by Mimatsu (1995), within 

nine days of initial unrest in December 1943, an area west 

of Mount Usu began to experience uplifting, folding, and 

faulting. Uplift to the west was initially strongest, after 

which its locus migrated in stages, eastward back toward the 

volcano. By June 1944, uplift ranged from 10 to 40 m. From 

June to November 1944, an additional 100 m of deformation accompanied 17 phreatic or phreatomagmatic explosions (Mimatsu, 1995). Finally, at the end of November 

1944, the first lava spine pushed through the deformed and 

cratered uplifted area. Dacite spines continued to grow until 

August 1945. The spines commonly had a jagged appearance (Mimatsu, 1995), unlike those at Mount St. Helens. 

Like those at Mount St. Helens since 2004, the spines were 

completely solid on extrusion and showed no tendency to 

flow. The spines all rose more or less vertically, and none 

were described as having appreciable lateral components of 

motion (Mimatsu, 1995).

One spine at Mount Usu, Kobu-yama or Bump Mountain, did have a form more analogous to those of the present 

eruption of Mount St. Helens. Mimatsu (1995) describes it 

as having a shape like the bottom of a boat, with a coating of 

pulverized rock or dirt and grooves or scratches parallel to the 

direction of extrusion. The powdery surface is probably analogous to the gouge-coated surface of spines at Mount St. Helens. Mimatsu (1995) also described horizontal bands on Bump 

Mountain that were similar to the bands commonly observed 

on smooth surfaces of spines during 2004–6 at Mount St. Helens. The bands on Kobu-yama tended to form during periods 

of rain. Apparently, ash and debris that was constantly sloughing from the steep slopes of Kobu-yama stuck when the slopes 

were wetted. Furthermore, ash and debris accumulated at the 

base of the growing spine and stuck after heavy rains to form 

ledgelike bands that later rose as the spine extruded (Mimatsu, 

1995). According to Mimatsu (1995), parts of these ledges 

were later shorn by falling debris.

Horizontal bands at Mount St. Helens are probably 

analogous, though not quite identical in origin, to those at 

Showa-Shinzan. Horizontal bands on spines at Mount St. 

Helens seem to have three forms, but all are plausibly related 

to moisture. The first, those most closely matching Mimatsu’s 

(1995) description, are shelf-like accumulations of fine to 

coarse debris that stick to the smooth surfaces of spines at 

certain times, then rise with spine growth (fig. 27). At Mount 

St. Helens, these seem to correspond to stormy periods at 

the volcano. We suggest that, as at Showa-Shinzan, addition 

of water to mixtures of fine and coarse debris immediately 

adjacent to the hot base of the extruding spine creates a weak 

cement that subsequently dries against the hot spine. A variation on this process involves only the fine-grained ash at the 

base of the spine and requires relatively less moisture. These 

bands are much less prominent. Photographs show fine debris 

concentrated at the base of the spines (fig. 27). This material 

requires less water to form a cement and, hence less moisture 

is required to create horizontal bands of such fine ash. A third, 

most common but least prominent variety of band, appears to 

involve periodic darkening of the gouge-coated surface. No 

one has examined these closely enough to understand if there 

is an accumulation of material associated with them or if they 

are merely stains. We speculate that many of them are related 

to nightly dew. All of these bands are fragile and ephemeral.
9. Growth of the 2004–2006 Lava-Dome Complex at Mount St. Helens, Washington 205

Conclusions

Spine extrusion and associated near-vent deformation at 

Mount St. Helens during 2004–6 presented an opportunity to 

test and apply various methods to track, measure, and characterize the dynamics and morphologies of extrusion and nearby 

deformation during a dome-forming eruption. We summarize 

here our chief conclusions drawn from our data and the methodology that generated it.CUMULATIVE RAINFALL, IN CENTIMETERS

0

5

10

10/29/04 10/30/04 10/31/04 11/1/04 11/2/04 11/3/04 11/4/04 11/5/04

DATE

Loowit falls; dashed where no data

Spirit Lake outlet

Sense 

of motion,

~7 m/d

~1 day

End heavy rain of 

11/2/04; rills 

less prominent

Rain stops 11/2/04;

coarse band 1 forms 

as gouge dries

Ash band 3 formed

early morning 11/4/04

Faint color

band

Ash sloughs and

accumulates 

at source 1980–86 debris

Photo taken

11 am, 11/4/04

Ash band 2,

early morning 11/3/04

A

B

Band 2

Band 1 Band 3

Photo

taken

Figure 27. Photograph of emerging spine and plot of rainfall from October–November 

2004. A, Close-up view of spine 3 east flank at 11 a.m. on November 4, 2004. Coarse band 

1 consists of mixtures of ash and coarse lithic lapilli, more subtle bands 2 and 3 consist 

of ash, and very faint ones may be stains. Band formation is associated with increases in 

moisture. USGS photo by J.S. Pallister. B, Cumulative rainfall versus time from USGS gages 

at Loowit falls (crater mouth) and at Spirit Lake outlet and time of emergence of notable 

bands. Rainfall data courtesy of K. Spicer, USGS.

Thermal infrared (TIR) surveys proved useful in differentiating events and structures that were cold from those that 

were hot. The TIR surveys showed that the explosions of October 1–5 were phreatic rather than magmatic. Imagery from TIR 

surveys also proved useful in identifying areas where spines 

were about to emerge. These were apparent as broad areas that 

warmed substantially in the days immediately prior to extrusion. Once spines began extruding, regular TIR surveys helped 

document their growth, character, and structures within them. 

Finally, TIR images helped monitoring crews identify places on 
206 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

the dome cool enough to place GPS instruments and accelerometers (Schneider and others, this volume, chap. 17).

Data collected as part of this study along with others in 

this volume suggest solid-state extrusion throughout the current dome-building episode. TIR measurement of deep cracks 

and newly exposed surfaces give temperatures well below the 

solidus of the dacite magma being erupted. Morphology of the 

various spines exhibits no flowage features like lobes, coulees, 

or ramp structures. Indeed, yield strength has apparently been 

so high that the spines can stand at steep slopes until they 

crumble. New spines have typically formed as previous ones 

have undergone brittle failure and fracturing, then shearing off 

to form stagnant crumbling masses.

Tracking of features in successive sets of aerial photographs and DEMs has enabled the development of surfacedeformation vector fields during 17 time intervals, which have 

varied in duration from 1 to 55 days. Each vector field gives a 

comprehensive spatial sense of deformation during that interval. Each also indicates the nature of advance or motion during 

the interval. Time-lapse photography and GPS instruments 

provided extra detail on much finer time scales for specific 

localities and localized fields of view, but the DEM tracking 

provided a valuable synoptic perspective.

Chief near-surface controls on spine extrusion during 

the 2004–6 eruption have been vent location, relict surfaces 

such as the 1980 crater structure and the 1980–86 dome, and 

spine remnants emplaced during previous phases of the present eruption—but not glacial ice. Ice as thick as 150 m has 

obscured eruptive processes, prevented formation of marginal 

angle-of-repose talus fans, and encouraged steep boundary 

slopes to the new dome complex through buttressing, but 

it has not significantly impeded spines pushing through it. 

Spines initially emerged at a location over the steep southfacing slope of the 1980–86 dome, which dictated their initial 

southward propagation. The glacier-filled space of the moat 

between the 1980 crater walls and the 1980–86 dome permitted southward propagation of spines 2 to 4 and funneled spines 

6 and 7 westward. Spine 2 impinged on the opposing slope 

of the crater and stopped. In contrast, recumbent whaleback 

spines 3 and 4 impinged at oblique angles and rotated eastward before cracking up. Although the vent location at the 

2004–6 surface shifted east and south more than 100 m before 

moving back to the west, its altitude increased ~200 m due to 

piling up of lava over the initial vent. The vent position relative to its initial trace at the 1986 surface need not have moved 

substantially. Once spine remnants occupied all available open 

space to the south, new spines thrust over previous remnants. 

Resistance to extrusion during intense periods of thrusting 

may have slowed extrusion rates because of backpressure 

effects during certain time intervals.

Although Crater Glacier had minimal influence on the 

growing spines, spine growth affected the glacier dramatically, initially dividing it into two arms and then bulldozing 

it hundreds of meters first east (east arm), then west (west 

arm), while heaping it more than 100 m higher than its original altitude.

The 2004–6 eruption has thus far differed from other 

well-documented historical eruptions in its solid-state character, its recumbent growth style, and its interaction with the 

glacier. On the basis of historical records, most domes grow 

endogenously; exogenously to produce thick units with a high 

aspect ratio and, sometimes, longer lava flows; or a combination of both (Blake, 1990). Peléean spines like those at Mont 

Pelée or Soufrière Hills are similar to those of the current 

eruption, but the 1943–45 eruption of Mount Usu provides 

the closest historical analogue. Perhaps the most similar of 

historically documented domes are those sometimes referred 

to as upheaved plugs (Blake, 1990). Such plugs appear to push 

up bodily like pistons and, when they reach the surface, have 

sufficient strength not to deform or spread outward but instead 

ascend vertically. Mimatsu (1995) beautifully documents the 

evolution of one such upheaval dome from 1943 to 1945 at 

Mount Usu in Japan. The intriguing variations in pluglike 

dome construction and evolution at Mount St. Helens since 

2004, not previously well documented, have been recumbent 

growth and interaction with an unusual combination of topographic constraints and glacial ice.

Acknowledgments

We thank Andy Harris and Michelle Coombs for careful 

reviews of this manuscript. Andy’s very detailed comments, 

in particular, greatly improved the paper. Cascades Volcano 

Observatory staff were generous with their time, their ideas, 

and their data.

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P.H., eds., A volcano rekindled; the renewed eruption of 

Mount St. Helens, 2004–2006: U.S. Geological Survey 

Professional Paper 1750 (this volume).

Swanson, D.A., and Holcomb, R.T., 1990, Regularities in 

growth of the Mount St. Helens dacite dome, 1980–1986, in

Fink, J.H., ed., Lava flows and domes, emplacement mechanisms and hazard implications: Berlin, Springer-Verlag, 

International Association of Volcanology and Chemistry of 

the Earth’s Interior, Proceedings in Volcanology 2, p. 3–24.

Thelen, W.A., Crosson, R.S., and Creager, K.C., 2008, 

Absolute and relative locations of earthquakes at Mount St. 

Helens, Washington, using continuous data; implications for 

magmatic processes, chap. 4 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Turcotte, D.L., and Schubert, G., 1982, Geodynamics: New 

York, John Wiley, 450 p.

Wadge, G., 1981, The variation of magma discharge during 

basaltic eruptions: Journal of Volcanology and Geothermal 

Research, v. 11, p. 139–168.

Walder, J.S., Schilling, S.P., Vallance, J.W., and LaHusen, 

R.G., 2008, Effects of lava-dome growth on the Crater 

Glacier of Mount St. Helens, Washington, chap. 13 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano 

rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).

Watts, R.B., Herd, R.A., Sparks, R.S.J., and Young, S.R., 

2002, Growth patterns and emplacement of the andesitic 

lava dome at Soufrière Hills Volcano, Montserrat, in Druitt, 

T.H., and Kokelaar, B.P., eds., The eruption of Soufrière 

Hills Volcano, Montserrat from 1995 to 1999: Geological 

Society of London Memoir 21, p. 115–152.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 10

1

 Department of Earth Science, Webb Hall, University of California, Santa 

Barbara, CA 93106; now at P.O. Box 750255, Fairbanks, AK 99775

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Photogeologic Maps of the 2004–2005 Mount St. Helens 

Eruption

By Trystan M. Herriott1, David R. Sherrod2, John S. Pallister2, and James W. Vallance2

Abstract

The 2004–5 eruption of Mount St. Helens, still ongoing 

as of this writing (September 2006), has comprised chiefly 

lava dome extrusion that produced a series of solid, faultgouge-mantled dacite spines. Vertical aerial photographs 

taken every 2 to 4 weeks, visual observations, and oblique 

photographs taken from aircraft and nearby observation 

points provide the basis for two types of photogeologic 

maps of the dome—photo-based maps and rectified maps. 

Eight map pairs, covering the period from October 1, 2004, 

through December 15, 2005, document the development 

of seven spines: an initial small, fin-shaped vertical spine; 

a north-south elongate wall of dacite; two large and elongate recumbent spines (“whalebacks”); a tall and elongate 

inclined spine; a smaller bulbous spine; and an initially 

endogenous spine extruded between remnants of preceding spines. All spines rose from the same general vent area 

near the southern margin of the 1980s lava dome. Maps 

also depict translation and rotation of active and abandoned 

spines, progressive deformation affecting Crater Glacier, 

and distribution of ash on the crater floor from phreatic and 

phreatomagmatic explosions. The maps help track key geologic and geographic features in the rapidly changing crater 

and help date dome, gouge, and ash samples that are no 

longer readily correlated to their original context because of 

deformation in a dynamic environment where spines extrude, 

deform, slough, and are overrun by newly erupted material.

Introduction

During its 2004–5 eruptive activity, Mount St. Helens 

extruded solid dacite lava onto the May 18, 1980, glaciercovered crater floor south of the 1980s lava dome. Beginning 

in October 2004, a series of seven spines extruded during 

the first 15 months of the eruption, following seismic unrest 

that began September 23, 2004. Two prominent whalebackshaped spines (3 and 4) that erupted from late October 2004 

to April 2005 followed two small spines (1 and 2) extruded 

in October 2004. From April through July 2005, tall, inclined 

spine 5 overrode remnants of previous spines to reach the 

dome’s maximum altitude as of September 2006 (Schilling 

and others, this volume, chap. 8); this spine subsequently 

subsided and partially disintegrated. Finally, beginning in 

early August 2005, spines 6 and 7 extruded westward, a 

marked change from previous spines, which were shoved 

south along the crater floor. Remnant spines (those no 

longer actively extruding) were rapidly degraded and overrun by subsequent spines. Crater Glacier, which formed 

after the cataclysmic 1980 eruption (Schilling and others, 

2004; Walder and others, this volume, chap. 13), was nearly 

bisected during the initial two months of the eruption and 

formed distinct east and west limbs that were substantially 

deformed in response to emplacement of the new lava dome.

Intensive monitoring of the eruption by the U.S. Geological Survey’s Cascades Volcano Observatory (CVO) has 

been supplemented by photographic documentation of the 

dome’s growth (this volume: Poland and others, chap. 11; 

Major and others, chap. 12; Dzurisin and others, chap. 14), 

including a series of 9×9-in. vertical aerial photographs. We 

present eight photogeologic maps traced from vertical aerial 

photographs of the Mount St. Helens crater that encompass 

15 months of the eruption from October 1, 2004, through 

December 15, 2005. Each map is presented as both a photobased and an accompanying rectified map without a photo 

base. The maps depict (1) the growth, stagnation, and subsequent burial or degradation of seven dacite spines (fig. 1); (2) 

the translation and rotation of geologic and geographic features throughout the evolution of the ongoing dome eruption; 

and (3) the progressive deformation of Crater Glacier as the 

growing dome displaced ice south of the 1980–86 dome. The 

photo-based maps also serve as a base for plotting locali-
210 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ties of, and giving geologic and geographic context to, rock, 

gouge, and ash samples that were collected on the growing 

lava dome. Photographic documentation of sample context is 

crucial in a constantly changing setting in which a sample’s 

original location loses significance, as when the base of a 

coherent spine sampled in January 2005 becomes a pile of 

transported and undifferentiated rubble by May 2005.

Methods

We compiled maps of the 2004–5 dacite dome eruption 

at Mount St. Helens on a sequence of ~1:12,000-scale vertical 

aerial photographs taken by Bergman Photographic Services 

on contract to CVO (Schilling and others, this volume, chap. 

8). Photo pairs encompassing the crater rim, new lava dome, 

and the 1980–86 lava dome were selected and examined with 

a stereoscope. The crater rim and 1980s lava dome provided 

a reasonably stable frame of reference in the dynamic crater environment. On the photographs we mapped units that 

highlighted changes in dome growth and morphology, as well 

as changes in the 1980s dome, rock debris, and glacial ice surrounding the dome. Uncertainties of the map units result from 

working without the benefit of on-the-ground field mapping, 

which early in the eruption was deemed too hazardous owing 

to possible explosions or rockfalls and rock avalanches. Our 

intent is to provide a heightened visual record of the locations 

and characteristics of major features and deposits in the crater 

during the ongoing eruption.

Oct 1 Nov 4

2004 2005

Dec 11 Mar 10 Apr 19 Aug 10 Sep 20 Dec 15 

5

10

15

20

25

30

35

40

7

6

5

4

3

1

2Dec 15

Oct 11 Oct 15

Oct 27

Apr 15

Jul 25

AREA COVERED, IN HECTARES

DATE OF AIR PHOTO (LINEAR SCALE)

Oct 9

D a t e

o f n e w s p i n e a p p e a r a n c e

Figure 1. Area of Mount St. Helens’ new dome and proportional coverage of individual spines 

through time. Areas measured from rectified photogeologic maps produced for dates of air photos 

shown. Date of new spine appearance from Vallance and others (this volume, chap. 9).

Two types of photogeologic maps were prepared—photobased and rectified. The latter lacks a base image in its final 

presentation. The visual content of the aerial photographs is 

immense, and unit boundaries are commonly vivid; however, 

the aerial photographs are not rectified, and therefore the 

photo-based maps are subject to scale variability and distortion within an individual image. Consequently, areas of, and 

distances between, map units on the photo-based maps are 

only comparable qualitatively. The rectified geologic maps 

allow quantitative comparison and complement the visually 

rich, nonrectified photo-based maps.

Although the two map series are based on the same photo 

pairs for any given date, variations exist within each map pair 

(photo-based and rectified). Dissimilarities result primarily 

from compiling the two types of maps at separate times but are 

compounded by employing different methods for assessing the 

contact locations (see below). Without doubt, the greater magnification and clarity possible with the plotter used to make the 

rectified maps resulted in a finer degree of detail, especially 

in differentiating gouge-covered surface, unroofed spine, and 

talus. Ultimately, small differences within map couplets persist. The two map series are not intended to be rigorously comparable but rather to be viewed and used for their individual 

strengths as discussed above.

Photo-Based Geologic Maps

We selected photographs on the basis of observable 

changes since previous photo coverage, visual clarity of crater 
10. Photogeologic Maps of the 2004–2005 Mount St. Helens Eruption 211

features, and pertinence to appropriate geographic setting of 

sample localities. These selection criteria winnowed the aerial 

photograph series from 18 to 8 sets (table 1). Our procedure 

began by scanning (300 dpi) each of the chosen set of photographs. The scanned images were enlarged and framed to 

encompass a similar area at roughly similar scale. We then 

traced rock, ash, and glacier units using Adobe ® Illustrator 10 

software. Although the photo-based maps are neither rectified 

nor georeferenced, we attempted to maintain the same field 

of view and azimuthal orientation (top to the north) of maps 

throughout the photo-based series.

We studied photo pairs stereoscopically and analyzed 

apparent three-dimensional surface morphology to identify 

photogeologic features and draw detailed map-unit contacts. 

Dashed lines on the maps indicate ambiguous contact locations. Thermal imagery and digital elevation models (DEMs) 

constructed from each set of photos (Schilling and others, 

this volume, chap. 8; Vallance and others, this volume, 

chap. 9) augmented the interpretation of geologic features 

and deposits viewed on the aerial photographs. We resolved 

additional crater details by referring to the extensive collection of oblique aerial photographs taken in the crater during 

field work by CVO scientists and by comparing aerial photographs with repeat images from time-lapse cameras at three 

fixed locations near the crater rim (this volume: Poland and 

others, chap. 11; Major and others, chap. 12; Dzurisin and 

others, chap. 14). These supplementary data elucidate areas 

within the crater that are obscured in the aerial photographs 

by condensed steam, shadows, or snow, or are difficult to 

interpret in plan view.

Date of photographyFigure 

No.

Photo 

used

Samples shown on 

photo map

Pair used 

for rectified 

map

Georegistration 

points

10/1/2004 2, 3A 6-5 No new dome 6-4, 6-5 19

11/4/2004 4, 3B 6-4 300, 301-1, 302-2, 

304-1, 304-2

6-3, 6-4 15

12/11/2004 5, 3C 6-3 305 6-2, 6-3 19

3/10/2005 6, 3D 7-2 306, 307, 308, 309, 

310, 311-1B, 312

7-2, 7-1 10

4/19/2005 7, 8A 6-4 313, 314, 315, 316 6-3, 6-4 19

8/10/2005 9, 8B 1-3 317, 318, 319, 320, 

321, 322, 323

1-2, 1-3 22

9/20/2005 10, 8C 1-2 1-2, 1-3 15

12/15/2005 11, 8D 1-3 324 1-2, 1-3 25

Table 1. Imagery used for photogeologic maps.

[Photo number (ex: 6-5) reported as flightline and frame number. Samples are whole-rock and gouge samples 

used for geochemical analyses described elsewhere in this volume. For brevity, sample numbers lack “SH” prefix 

(for example, complete sample No. is SH300). For stereographic pair, first listed is left-hand photo, southerly of 

pair; all flightlines oriented north-south. Column for georegistration points indicates number of points used in 

transformation to better rectify the photogrammetric geologic map.]

Rectified Geologic Maps

Conventional photogrammetric methods were used to 

rectify photogeologic maps. We mapped contacts while viewing paired vertical aerial photos stereoscopically on a Kern 

PG-2 optical-mechanical stereographic plotter. Our instrument at CVO lacks an electronic digitizer, so the resulting 

pencil-on-mylar maps were scanned and hand digitized using 

MapInfo ® Geographic Information System (GIS) software. 

Points of known geographic position, such as prominences on 

the crater rim and on the 1980s dome within the crater, served 

as pseudobenchmarks by which the maps were registered 

geographically. The crater-rim control points are prominent 

apices and craggy summits whose locations were derived from 

a lidar image with 2-m cells produced in November 2004. The 

1980s-dome control points are rock spires recognizable on a 

1988 high-resolution topographic map of the crater floor.

The PG-2 plotter produced rectified maps under most 

conditions; however, the photos encompass substantial altitude 

variation, more than 600 m, across short horizontal distances. 

Parallax-free models are therefore difficult to achieve, and a 

small amount of distortion may occur in the resulting maps. 

The problem is exacerbated by using the crater rim for geographic registration, whereas the area of geologic interest lies 

entirely on the crater floor, substantially lower in altitude. 

Moreover, snow cornices modify the crater rim in winter, with 

some forming prominent topographic apices recognizable in 

DEMs and aerial photos; these windblown snow and ice features can rebuild in slightly different geographic positions from 

photo to photo, a source of error for which we cannot system-
212 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

atically account. The crater rim itself retreats episodically by 

erosion, but no noticeable changes caused by this effect were 

observed in the 15 months that span our geologic maps.

To minimize distortion, the scanned pencil-on-mylar 

images were transformed by rubbersheeting methods. Specifically, the image was brought into its approximately correct 

position by a transformation matrix (general perspective 

projection transformation) and then by triangular irregular net 

adjustment (for fuller discussion of methodology, see Schilling and others, this volume, chap. 8). The transformations 

involved 15 to 25 registration points, except for the imagery of 

March 10, 2005, which used 10 points (table 1).

Ultimately, the resulting precision and probably the 

accuracy of the rectified maps is within plus or minus 12 m. 

This value is acceptable owing to the 1:12,000 scale of the 

aerial photographs and the standard convention that a geologic 

map should portray precision to at least 1 mm on the map. 

The precision was tested empirically by draping the resulting 

linework on existing DEMs. The crater rim was traced as part 

of each rectified map to provide a visual-empirical guide for 

comparison with the crater rim on the DEM. Rock spires on 

the 1980–86 dome provided intracrater tie points to further 

assess precision. The crater-rim test shows that all tracings lie 

within 5 m of each other. As a test of accuracy, the rock-spire 

test gives geographic coincidence generally within 1–5 m, 

with a few strays as far as 10 m. The resulting rectified maps 

are more than adequate, in both precision and accuracy, for the 

cataloging and archiving of geologic information.

Sample Localities

CVO staff collected (chiefly by helicopter dredge) and 

analyzed 26 samples of 2004–5 lava-dome dacite after the first 

dacite spine erupted in mid-October 2004 (Pallister and others, 

this volume, chap 30; Thornber and others, 2008). Sample 

localities (table 1) are plotted on the photo-based maps that 

depict relevant geologic and geographic features, providing 

approximate spatial and temporal context for the samples.

Description of Photogeologic Map 

Units

Map units described here include rock and debris from 

the 2004–5 dacite dome, phreatic or phreatomagmatic and 

rockfall-generated ash deposits, crater floor debris, 1980s 

lava dome, and deformed and undeformed glacial ice. Labels 

coupled with a plus sign (+) indicate where thin surficial 

deposits blanket other map units. The underlying map unit is 

listed first, followed by the symbol of the blanketing deposit. 

For example, unit label gd+a indicates deformed glacier overlain by a veneer of ash (see Dome Map Units, below).

We use lava dome, or simply dome, to describe the 

composite feature that comprises extrusive spines or lobes, 

endogenous growth, and talus. We mapped each spine as an 

individual unit of extrusive lava, typically massive, and analogous to a lobe or flow. Endogenous growth of some spines 

results in a rubbly surface. A whaleback is a smooth, striated 

spine, be it recumbent or vertical. The terms “new” and “old” 

dome are used informally to differentiate between the 2004–5 

and 1980–86 lava domes, respectively.

The term Opus identifies an informally named part of 

the 1980s dome that was displaced southward during a 1985 

eruption. The name originated from a benchmark used to track 

this movement and later was extended through casual usage 

to denote the entire elongate geographic ridge that resulted. 

Opus became important during the renewed eruption in 2004 

because the conduit breached the surface near it, deforming 

both Opus and the adjacent Crater Glacier.

Crater Glacier deformed because of upwarping early in 

the eruption and then by lateral compression as extruded lava 

shoved the glacier aside. We demarcate deformed and undeformed glacial ice by tracing the abrupt topographic break in 

slope between them. This boundary was a deformation front 

within a compositionally coherent unit, unlike lithologic contacts, which are planar features separating discrete bodies of 

rocks or deposits.

2004–2005 Dacite Dome Map Units

Spines—Lava lobes, numbered 1–7 according to eruptive 

sequence of dacite lava extruded onto floor of Mount 

St. Helens crater during eruption ongoing since 

October 2004 (Vallance and others, this volume, 

chap. 9). Moderately porphyritic, with phenocrysts 

of plagioclase, amphibole, hypersthene, and Fe-Ti 

oxides. Includes three gross textural features––

surface gouge; unroofed, ragged spine; and remnant 

spine—described below. Thus, map symbol may be 

a composite; for example, s4u indicates unroofed, 

ragged part of spine 4 after erosion has destroyed its 

smooth carapace. Spines defined as follows:

s7 Spine 7—Initially endogenous dacite spine; extruded 

between remnants of spines 5 and 6 beginning in 

October 2005. Longest-lived and latest in sequence as 

of September 2006

s6 Spine 6—Rubble-covered dacite spine; extruded during 

August and September 2005. Grew coincident with 

and adjacent to a graben, or sag, that developed west 

of spine 5

s5 Spine 5—Large south-trending dacite spine; extruded 

from late April through July 2005. Highest spine 

measured as of September 2006 (Schilling and others, 

this volume, chap. 8)
10. Photogeologic Maps of the 2004–2005 Mount St. Helens Eruption 213

b Ballistic craters—Depressions in glacier caused by 

bomb impacts during October 1, 2004, and March 8, 

2005, explosions

Other Crater Map Units

rp Crater-floor roof pendant—Crater-floor debris that 

rested atop spine 3 upon its initial extrusion

gl Crater Glacier—Glacial ice and enclosed rock debris 

that accumulated subsequent to May 18, 1980, craterforming eruption. Compression and thickening by 

crater-floor uplift and extrusion of new lava dome 

increased rate of flow northward around 1980s dome. 

Shown separately are the following:

gd Deformed Crater Glacier—Deformed, typically uplifted 

and crevassed glacial ice shoved aside as lava dome 

grew

gs Stranded glacial ice—Remnant of ice from east glacier 

left perched on flank of Opus by crater-floor uplift

la Lahar deposits—Remnants of small muddy debris flows 

and tiny pyroclastic flows generated by interaction of 

hot dacite dome rock and snow or glacial ice

cd Crater-floor debris and talus—Rubble uplifted by 

actively growing 2004–5 dacite dome

cw Crater wall—Strata forming crater walls and outer 

flanks of Mount St. Helens. Unit only appears on 

rectified maps

op Opus—Ridge created on south side of 1980s dome 

by a small graben that formed during October 1985 

eruption

od 1980–86 lava dome—Older dome emplaced during a 

6-year period following the May 18, 1980, craterforming eruption

h Melt pit—Steeply walled pits at glacier/lava dome 

contact melted out by fumaroles. Unit only appears 

on rectified maps

* 2004–5 sample location—Sample location, showing 

number (for example, 300; the SH prefix is not 

included in the map label); most samples collected 

by helicopter dredging tools. Shown only on photobased maps

s4 Spine 4—Large, elongate, south-trending dacite 

whaleback; extruded from late December 2004 

through late April 2005

s3 Spine 3—Large, elongate, south-trending dacite 

whaleback; extruded from late October through late 

December 2004. First occurrence of gouge-mantled 

spine with striated carapace, characteristic of all 

subsequent spines

s2 Spine 2—Elongate south-trending dacite spine; 

extruded in middle to late October 2004

s1 Spine 1—Small northeast-trending finlike dacite spine; 

initial effusive product; extruded in mid-October 

2004

u Unroofed spine—Part of actively growing spine that 

no longer retains gouge-mantled carapace, which 

typically was shed through repeated rockfalls

r Spine remnants—Partially intact rubble of inactive 

spines. Moved constantly as growth of new spine 

wedged them away from vent

g Gouge—Cataclastic carapace 1–2 m thick that 

characterized surfaces of spines 3–7 where they 

first emerged from the vent. Commonly removed by 

fracturing and degradation

t Talus—2004–5 dacite dome talus resulting from 

rockfall from active and abandoned spines. May be 

darker in color than adjacent hot talus where wet or 

covered by damp ash

a Ash—Tephra, produced chiefly by phreatic or 

phreatomagmatic explosions and rockfalls. Shown 

separately on some maps are the following units:

a2 Ash from March 8, 2005, explosion—Primary 

magmatic and accidental lithic tephra generated 

during largest phreatomagmatic explosion of ongoing 

eruption. Resulted from near-vent fallout from plume 

of steam and ash that rose to about 11 km height

a1 Ash from October 1, 2004, explosion—Fine ash that 

blanketed northwest sector of Mount St. Helens crater

v Vent for October 1, 2004, explosion—Pit about 50 m in 

diameter that produced a short-lived steam-and-ash 

plume of phreatic or phreatomagmatic origin. Source 

of first explosive activity of ongoing eruption
214 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Discussion of Photogeologic Maps

Our discussion of the maps focuses on salient geologic, 

geographic, topographic, and glacial features generated 

during the first 15 months of the eruption. These include 

spine evolution, phreatic and phreatomagmatic explosions, 

uplifted crater-floor rocks, deformation of Opus, deformation of the Crater Glacier, and extensive rockfall-generated 

ash deposits. We do not discuss all features depicted in the 

images, as our goal is to provide a guide for visualizing the 

most notable changes on the crater floor between the dates 

of successive maps.

October 1, 2004

The October 1 phreatic explosion left its mark on the crater scene photographed that afternoon (fig. 2). Dark-gray ash 

(a1) mantles the western crater floor, having emanated from a 

prominent vent (v) in the Crater Glacier. Numerous craters on 

west Crater Glacier define a more restricted field of ballistic 

craters (a1b).

Adjacent to the vent on the east is highly fractured and 

uplifted glacial ice (gd), a welt approximately 52,000 m2

in extent that grew rapidly as magma ascended toward the 

surface (fig. 3A). The base of the topographic rise to the welt, 

distinctive in stereo viewing but not readily apparent in the 

single base image of figure 2, defines the contact separating deformed from undeformed glacier. With time, some of 

the deformed ice seen in figure 2 became isolated from the 

main mass of Crater Glacier. The Opus ridge (op) is also 

deformed.

November 4, 2004

The November 4 image shows spine 3 (s3) and remnants 

of spines 1 (s1r) and spine 2 (s2r) of the new dome (fig. 4). 

The east-flank carapace of spine 3 is a well-preserved striated 

gouge surface (s3g). A distinctly whiter band marks the freshest gouge along the eastern base. Remnants of spines 1 and 2 

are north and west, respectively, of spine 3.

At least one roof pendant (rp) of old dome and craterfloor debris is atop spine 3. In the photo, the pendant is seen 

as the darker area near the crest of the spine (fig. 4). Dark-gray 

Figure 2. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken October 1, 2004. For explanation 

of unit symbols, see text. Shoestring notch is an informal geographic name for the topographic cleft that is the truncated head 

of the Shoestring Glacier, which was largely destroyed in the 1980 eruption of Mount St. Helens.
10. Photogeologic Maps of the 2004–2005 Mount St. Helens Eruption 215

debris (c d) along the eastern periphery of spine 3 represents 

deformed fragments of the 1980–86 dome, first dragged southward 100–200 m on the roof of spine 3 and then eroded off 

(fig. 3B). Uplifted crater floor material lies east of the 1980s 

dome debris (fig. 3B).

East of the new dome is the highly deformed east limb of 

Crater Glacier (gd), crisscrossed by crevasses to form a serac 

field. The growing dome has nearly split the glacier, such that 

its east and west limbs have become geographically distinct. 

Nevertheless, the glacier remains contiguous around the south 

side of the new dome.

Deformation of Opus (op) is largely complete but is 

not easily discerned at the scale of this aerial image (fig. 

4); stagnant and broken glacial ice mantles its south side. 

Increased heat flow on its western side has melted snow from 

an irregular area, but the exposed, darkened mass of jumbled 

rock is likely little different in form from the snow-covered 

nonthermal area of Opus.

Two or three small debris fans and more lobate, slurried deposits (la) originate from spine 1 remnants. Thinner 

lahar deposits extend along a narrow trace down the center of 

the cleft that separates the west flanks of the new and 1980s 

domes from the west limb of Crater Glacier (gl) (fig. 4).

December 11, 2004

Rain, melted snow, or condensed steam has left the snowfree areas very dark gray on this overcast day (fig. 5). Spine 

3 has greatly lengthened, earning the name “whaleback.” Its 

gouge-mantled surface (s3g), where wet, is dark grayish brown. 

The western slope of spine 3 has spalled to form talus (t), including some large house-size blocks. All the uplifted crater-floor 

debris seen along the west flank in previous images is now indistinguishable, likely buried by talus. Along the southeast side, 

steam and coarse blocks mark an elongate slab of spine 3 (s3u), 

detached, deformed, and wedged between spine 3 and the glacier 

(fig. 3C). Spine 1 and 2 remnants remain visible, although they 

are somewhat obscured by steam in this image (fig. 5).

Notable deformation continues on the east limb of Crater 

Glacier (gd), but only small changes occur along its west limb. 

The glacier remains intact along the south side of the new 

dome, but it has narrowed greatly to only about 60 m (fig. 

3C). Glacial ice on parts of Opus has disappeared, causing an 

expansion of the area of Opus (compare figs. 3B, 3C).

March 10, 2005

Geomorphically prominent spine 4 has by this time 

supplanted spine 3, which underwent rapid fragmentation 

during early January 2005 (fig. 6). Spine 4 evolved as the 

northernmost part of spine 3 split along a vertical fracture 

and decoupled at its root from the rest of spine 3 to the south. 

As it grew southward, spine 4 isolated a small relic of spine 

3 (s3ru) to the west as it bulldozed the remainder to the east 

(Vallance and others, this volume, chap. 9); these remnants 

emit steam where they abut glacial ice along the east margin. 

The stunning striated and gouge-mantled carapace of spine 

4 (s4g), as well as the crevassed and thickened tongue of the 

deformed glacier (gd, east of spine 4; see below), take center 

stage in this photo. Remnants of spines 1 and 2 are increasingly ragged in appearance. Virtually no uplifted crater-floor 

debris remains adjacent to the new spines, as it has been 

buried by talus (t).

A large phreatomagmatic explosion on March 8 

distributed ash (a2) across the north and east sectors of the 

crater, but large ballistic fragments fell chiefly north and 

northwest of the new dome. Opus and its shroud of snow 

and ice lie pinched between the new and 1980s domes. The 

stagnant ice on Opus (unit gs on fig. 3D) is now fully separated from the actively deforming glacier.

The east limb of Crater Glacier deformed dramatically 

through March and into April, producing the prominent, 

lobate deformation front visible near the southeast margin of 

the 1980–86 dome. The amplified glacial deformation slowed 

significantly in subsequent weeks.

April 19, 2005

Spine 5 (s5) at this time lies at the north end of the 

spine 4 whaleback remnants, with much of its gouge-covered 

surface (s4rg still intact (fig. 7). Extensive sloughing of spine 

4 was first observed following a 2-week stormy period in late 

March 2005. Spine 4’s southward growth and collision with 

the southern crater wall caused its breakup. Spine 3 (s3ru) is 

reduced to remnants flanking spine 4. Spine 1 and 2 remnants 

persist, although spine 2 (s2r) steams vigorously where shoved 

into the western limb of the glacier, so that little of it is visible 

in this photo (fig. 7).

Snowfall subsequent to the March 8 phreatomagmatic 

explosion buried the thin veneer of ash and ballistic fragments 

from that event. Distinctive ash (a) blanketing the east limb 

of the crater glacier on this image is a localized deposit that 

resulted from a large hot rockfall off the northeast face of 

spine 4 remnants (fig. 8A). Blocky rockfall material stopped 

abruptly at a topographic rise on the deformed east glacier; 

the rise is well pronounced in stereo viewing, but not readily 

apparent in figure 7. Ash elutriated and surged ahead for a 

short distance, forming the darkest of the deposit seen in the 

photo (gd+a, northern extent) (fig. 7). Even finer ash traveled 

south, presumably downwind, before being deposited (gd+a, 

southern extent).

August 10, 2005

Spine 6 (s6) separates from spine 5 at this time and edges 

westward (fig. 9), contrasting with previous spines that pushed 

along a south-southeastward trajectory. This new growth pattern focuses deformation of Crater Glacier in its west arm. 

The heretofore relentlessly moving terrain of spine 3, 4, and 5 

remnants stagnates over the next few months (compare spatial 

relations in figs. 8B, 8C, 8D). Spine 3 remnants (s3ru) are vis-
563000 563500 5116500 562000 562500 5116000 5115500561500 0 200 400 m 5117000 CRATER RIM Shoestring notchBrutus op od gl gd 562000 562500 

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op

November 4, 2004

December 11, 2004 March 10, 2005

A

B

C D

October 1, 2004

0 300 m

0 300 m 0 300 m

Figure 3. Rectified geologic maps of Mount St. Helens’ new dome. A, October 1, 2004. B, November 4, 2004. C, December 11, 2004. D, March 
Lahar deposits from area of new domeTalus from flank of spineCrater-floor debrisWedge of downdropped crater-floor debris

Melt pit at glacier marginTalus from crater walls and 1980−86 domeStrata of crater walls and outer flanks

TEPHRA DEPOSITS ON MAP OF MARCH 10, 2005

Northern and southern limits of lapilli and ash

Ash distribution

Ballistic fallout  Inner, dense fall  Outer limit of sparse fall

Spine of 2004−2005 eruption (ongoing in 

December 2005)—Darker area mantled by gouge

Notable crevasse on glacier—Resulting fro dome emplacement. Only crevasses that 

developed after the previous map date ashown

Distribution of coarse ballistics—Dashed 

where lack of snow cover makes limit of cratering indistinct

1980–86 dome—Darker area is Opus ridge 

(op)

Glacial ice—Darker area deformed by dom

growth

Vent for phreatomagmatic eruption of October 

1, 2004

op

od

gd gl

h

cw

la

t

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s3g

cd

FEATURES ASSOCIATED WITH NEW DOME OTHER FEATURES IN CRATER

EXPLANATION

TEPHRA DEPOSITS ON MAP OF OCTOBER 1, 2004

Figure 3.—Continued.
218 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 4. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken November 4, 2004. 

Stars mark sample localities; see table 1.

Figure 5. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken December 11, 

2004. Star marks sample locality; see table 1.
10. Photogeologic Maps of the 2004–2005 Mount St. Helens Eruption 219

Figure 6. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken March 10, 2005. Stars 

mark sample localities; see table 1.

Figure 7. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken April 19, 2005. Stars 

mark sample localities; see table 1.
562000 562500 563000 5116500 5116000 gl h h la gl gd gd gs s1r s2r s3ru s4ru t t s4rg s4rg s3ru s4ru s4ru s5g od op cw cw h 562000 562500 

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April 19, 2005September 20, 2005 December 15, 2005

August 10, 2005

D

CRATER RIM CRATER RIM

CRATER RIM CRATER RIM

SAG

0 300 m 0 300 m0 300 m 0 300 m

Figure 8. Rectified geologic maps of Mount St. Helens’ new dome. A, April 19, 2005. B, August 10, 2005. C, September 20, 2005. D, December 15, 2
Talus from flank of spine

Melt pit at glacier margin

FEATURES ASSOCIATED WITH NEW DOME OTHER FEATURES IN CRATER

Talus from crater walls and 1980−86 domStrata of crater walls and outer flanks

Limit, density current of ashy sediment from rockfall

Outer limit of dilute ash from rockfall

Spine of 2004−2005 eruption (ongoing in 

December 2005)—Darker area mantled by gouge

Notable crevasse on glacier—Resulting fr dome emplacement. Only crevasses tha

developed after the previous map date shown

Steaming lineament—On figure 8C. Likely 

marks incipient renewed exposure of spine 2 from beneath mantle of talus as seen on figure 8D

1980–86 dome—Darker area is Opus ridg

(op)

Glacial ice—Darker area deformed by do

growth

op

od

gd gl

h

cw

t

s3us3g

EXPLANATION

TEPHRA DEPOSITS ON MAP OF APRIL 19, 2005

Figure 8.—Continued.
222 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 9. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken August 10, 2005. 

Stars mark sample localities; see table 1.

Figure 10. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken September 20, 2005.
10. Photogeologic Maps of the 2004–2005 Mount St. Helens Eruption 223

ible only on the eastern periphery of the dome, whereas talus 

from spines 5 and 6 (t5, t6) now buries all remnants of spines 

1, 2, and 3 to the west.

The washed-out appearance of the August 10 image 

(fig. 9) results from dust generated by summer rockfalls from 

both the new dome and the crater walls. Much of the lightdark color contrast in this photo arises from the water content 

of the dusted surface, which is darker where underlain by 

snow or ice.

September 20, 2005

Spine 6 (s6) continues to propagate westward (fig. 10). 

On its east side, rockfalls continue to incise the relatively 

small remaining parts of the gouge-mantled carapace (s6g). 

Abundant rockfalls from spine 5 have unroofed the core of that 

spine (s5ru), which now has a prominent axial crest. Spine 6’s 

net westward motion is accompanied by subsidence between 

the actively growing part of the dome and the relatively stationary remnants of spines 3, 4, and 5 (fig. 8C).

Deformation in the west arm of Crater Glacier increases 

as spine 6 grows westward. Large new crevasses radiate 

outward from the contact with the expanding dome (fig. 8C). 

To the south, the band of glacial ice that bridges the east and 

west limbs of Crater Glacier has narrowed to about 6 m in 

width (fig. 8C).

December 15, 2005

In early October, spine 7 (s7) emerged from the sag that 

formed between spine 6 and remnants of spine 5 (fig. 11; Vallance and others, this volume, chap. 9). Spine 7 grew westward, pushing spine 6 remnants westward, overthrusting them, 

and burying them in talus. A small sliver of spine 6 (s6ru) still 

crops out along the margin of west Crater Glacier. Snow covers the largely cooled terrain of spines 3, 4, and 5, distinguishing it from the snow-free, hot, actively growing region of the 

dome. Spine 4 and 5 remnants are becoming more prominent 

as their cores are exhumed by rockfalls.

The deformation of Crater Glacier along its west arm is 

remarkable in this image (fig. 11). Deformation since August 

2005 has created a jumbled serac field west of spine 6 remnants, and numerous crevasses underlie the snow-covered 

areas of the deformed west arm of the glacier.

Acknowledgments

We thank Christina Heliker and Steve Schilling for their 

insightful manuscript reviews, which led to greater clarity. 

Any remaining errors are the sole responsibility of the authors. 

The first author also thanks A.R. Wyss for his comments on an 

early version of the text.

Figure 11. Photo-based geologic map of Mount St. Helens’ new dome from aerial photo taken December 15, 2005.
224 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

References Cited

Dzurisin, D., Lisowski, M., Poland, M.P., Sherrod, D.R., and 

LaHusen, R.G., 2008, Constraints and conundrums resulting from ground-deformation measurements made during the 2004–2005 dome-building eruption of Mount St. 

Helens, Washington, chap. 14 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Major, J.J., Kingsbury, C.G., Poland, M.P., and LaHusen, 

R.G., 2008, Extrusion rate of the Mount St. Helens lava 

dome estimated from terrestrial imagery, November 

2004–December 2005, chap. 12 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Pallister, J.S., Thornber, C.R., Cashman, K.V., Clynne, M.A., 

Lowers, H.A., Mandeville, C.W., Brownfield, I.K., and 

Meeker, G.P., 2008, Petrology of the 2004–2006 Mount St. 

Helens lava dome—implications for magmatic plumbing 

and eruption triggering, chap. 30 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Poland, M.P., Dzurisin, D., LaHusen, R.G., Major, J.J., Lapcewich, D., Endo, E.T., Gooding, D.J., Schilling, S.P., and 

Janda, C.G., 2008, Remote camera observations of lava 

dome growth at Mount St. Helens, Washington, October 

2004 to February 2006, chap. 11 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Schilling, S.P., Carrara, P.E., Thompson, R.A., and Iwatsubo, 

E.Y., 2004, Posteruption glacier development within the 

crater of Mount St. Helens, Washington, USA: Quaternary 

Research, v. 61, no. 3, p. 325–329.

Schilling, S.P., Thompson, R.A., Messerich, J.A., and Iwatsubo, E.Y., 2008, Use of digital aerophotogrammetry to 

determine rates of lava dome growth, Mount St. Helens, 

Washington, 2004–2005, chap. 8 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Thornber, C.R., Pallister, J.S., Rowe, M.C., McConnell, S., 

Herriott, T.M., Eckberg, A., Stokes, W.C., Johnson Cornelius, D., Conrey, R.M., Hannah, T., Taggart, J.E., Jr., 

Adams, M., Lamothe, P.J., Budahn, J.R., and Knaack, 

C.M., 2008, Catalog of Mount St. Helens 2004–2007 dome 

samples with major- and trace-element chemistry: U.S. 

Geological Survey Open-File Report 2008–1130, 9 p., with 

digital database.

Vallance, J.W., Schneider, D.J., and Schilling, S.P., 2008, 

Growth of the 2004–2006 lava-dome complex at Mount St. 

Helens, Washington, chap. 9 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Walder, J.S., Schilling, S.P., Vallance, J.W., and LaHusen, 

R.G., 2008, Effects of lava-dome growth on the Crater 

Glacier of Mount St. Helens, Washington, chap. 13 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano 

rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).
1

 U.S. Geological Survey, PO Box 51, Hawaii National Park, HI 96718

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

3

 U.S.D.A. Forest Service, Gifford Pinchot National Forest, 10600 NE 51st 

Circle, Vancouver, WA 98682

Chapter 11 

Remote Camera Observations of Lava Dome 

Growth at Mount St. Helens, Washington, 

October 2004 to February 2006

By Michael P. Poland1, Daniel Dzurisin2, Richard G. LaHusen2, Jon J. Major2, Dennis Lapcewich3, 

Elliot T. Endo2

, Daniel J. Gooding2, Steve P. Schilling2, and Christine G. Janda2

Abstract

Images from a Web-based camera (Webcam) located 8 

km north of Mount St. Helens and a network of remote, telemetered digital cameras were used to observe eruptive activity 

at the volcano between October 2004 and February 2006. The 

cameras offered the advantages of low cost, low power, flexibility in deployment, and high spatial and temporal resolution. Images obtained from the cameras provided important 

insights into several aspects of dome extrusion, including 

rockfalls, lava extrusion rates, and explosive activity. Images 

from the remote, telemetered digital cameras were assembled 

into time-lapse animations of dome extrusion that supported 

monitoring, research, and outreach efforts. The wide-ranging 

utility of remote camera imagery should motivate additional 

work, especially to develop the three-dimensional quantitative 

capabilities of terrestrial camera networks.

Introduction

During the 20th century, advances in technology have 

added an array of geophysical and geochemical instrumentation to the modern volcanologist’s toolkit. The study 

of active volcanoes has relied increasingly upon datasets 

derived from such technology to infer the mechanics of volcanic processes, which often occur at depth. As detailed in 

this volume, many geophysical and geochemical techniques 

have been applied to improve understanding of eruptive 

activity at Mount St. Helens in 2004–6. Visual surveillance 

in volcanology, however, remains critical for providing 

“ground truth” necessary to confirm inferences drawn from 

geophysical and geochemical data.

Visual observations can be recorded by imaging systems 

on the ground or in an aircraft or spacecraft. For example, photogrammetric applications of aerial photography to volcanoes 

include quantification of large-scale deformation before the 

1980 eruption of Mount St. Helens (Moore and Albee, 1981) 

and calculations of erupted volumes at Stromboli, Italy, in 

2002–3 (Baldi and others, 2005) and at Mount St. Helens in 

2004–6 (Schilling and others, this volume, chap. 8). Groundbased visual imagery is equally important for observing 

volcanic activity, having the advantages of low cost, frequent 

image acquisition, and flexibility in deployment. Starting in 

September 2004, we made extensive use of terrestrial cameras 

to investigate activity at Mount St. Helens using a continuously operating Webcam located 8 km north of the volcano 

and repeat photographs from a network of remote, telemetered 

digital cameras. The imagery was used to evaluate broad-scale 

eruptive activity in near real time, correlate geophysical signals with changes in eruptive activity, investigate dome extrusion processes, track the evolution of the eruption (including 

deformation of glacial ice) over time, and assess weather 

conditions for planning fieldwork.

We describe here the remote camera deployments and the 

activity recorded at Mount St. Helens during the period October 2004 to February 2006. Other types of camera deployments and applications at Mount St. Helens are described 

elsewhere in this volume. Results from high-rate, small fieldof-view photography experiments designed to measure smallscale changes in dome extrusion are described by Dzurisin and 

others (this volume, chap. 14). Major and others (this volume, 

chap. 12) discuss quantitative dome-growth measurements 

A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008
226 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

using images from a single remote camera in combination 

with a digital elevation model.

Previous Uses of Visual Observation 

Systems to Monitor Active Volcanoes

Volcanology is fundamentally an observational science, 

and repeat observations from fixed locations have proven critical to the documentation and interpretation of many volcanic 

processes. During dome extrusion in 1902–5 at La Montagne 

Pelée, Martinique, Lacroix (1908) collected photographs 

from fixed vantage points to record the development of the 

dome over the course of the extrusive phase of the eruption. In 

1944–45, Mimatsu Masao, the postmaster of the Sobetsu Post 

Office in Japan, lacked camera equipment but documented 

the growth of the Showa-Shinzan dome at the base of Mount 

Usu in a detailed diary and with careful sketches. His unique 

surveying methods included a fixed observation point behind 

the post office, from where he viewed the growing dome by 

resting his chin on a level and by using a series of horizontally 

stretched cords as reference lines. His drawings of the uplift 

and dome growth from this vantage point were presented at 

the 1948 International Association of Volcanology conference 

in Oslo, Norway, and what came to be known as “Mimatsu 

diagrams” were praised as “the only existing records of the 

entire birth of a volcano” (Mimatsu, 1995).

One of the best known volcano photographic sequences 

was taken by Gary Rosenquist at Mount St. Helens during the 

landslide and lateral blast of May 18, 1980. The Rosenquist 

photos, and similar sequences taken from other locations 

around the volcano at the start of the eruption, were critical 

to understanding the development of the landslide and lateral 

blast (Voight, 1981; Voight and others, 1981; Moore and Rice, 

1984), lahar initiation (Pierson, 1985), pyroclastic stratigraphy 

(Criswell, 1987), and the question of whether or not the blast 

was a product of one or two explosions (Hoblitt, 2000). Fixedvantage-point cameras from more than 100 repeat terrestrial 

photography and time-lapse film stations were also a key 

tool for studying dome building at Mount St. Helens during 

1980–86 (Topinka, 1992).

The development of digital cameras has facilitated the 

use of visual observation systems at volcanoes. At Kïlauea 

Volcano, Hawai‘i, time-lapse digital cameras powered by 

solar panels and encased in weatherproof boxes now record 

details of volcanic events, including ground deformation, 

vent collapses, and surface breakouts of lava (Orr and Hoblitt, 

2006). Repeat views from remote digital cameras have also 

been employed at Soufrière Hills volcano4

, Montserrat, where 

they provided important visual documentation of dome growth 

(Watts and others, 2002) and of the 2003 catastrophic dome 

collapse (Herd and others, 2005).

Remote Camera Systems Used at 

Mount St. Helens

During October 2004 to February 2006, two types of 

remote camera systems were used for visually monitoring 

eruptive activity at Mount St. Helens—a Webcam and a network of remote, telemetered digital cameras. These systems 

are described below.

Webcam

A Webcam, herein referred to as the “VolcanoCam,” 

was installed in 1996 at the U.S. Department of Agriculture–

Forest Service’s Johnston Ridge Observatory (JRO; fig. 1), 

8 km north of Mount St. Helens (fig. 2A). At that time, the 

installation of the camera was more of an Internet novelty for 

the Gifford Pinchot National Forest (GPNF) and the Mount 

St. Helens National Volcanic Monument. The GPNF had just 

established one of the first Web sites within the Forest Service, 

and the addition of the VolcanoCam, they hoped, would 

provide a boost to forest recreation use by stimulating general 

interest in the area.

The VolcanoCam operated with minimal problems for 

7 years until it suffered a mechanical failure in June 2003. 

Funding problems delayed replacement of the camera for more 

than a year. New equipment was finally procured and installed 

on September 23, 2004—coincidentally the day that seismic 

unrest began at Mount St. Helens. The new VolcanoCam was a 

color charge-coupled camera that provided a signal of 525 TV 

lines at 30 frames per second (terminology from the National 

Television System Committee standards). Still images were 

uploaded every five minutes to the Forest Service’s national 

Web server. The clock on the camera was not synchronized to 

Internet time and was probably only accurate to within about 

1 minute.

Access to the camera was initially limited to Forest 

Service and USGS staff, but the VolcanoCam was opened 

for public access on September 27, 2004, and immediately 

became a major attraction (http://www.fs.fed.us/gpnf/volcanocams/msh/, last accessed January 28, 2008). The number 

of hits on the VolcanoCam Web site became so large that the 

main Forest Service Web server crashed several times, and 

excessive bandwidth use threatened the main U.S. Department 

of Agriculture Web servers. A Web caching system alleviated 

the most serious bandwidth concerns. Fourth-quarter 2004 

statistics for all Federal government Web sites later revealed 

that the Forest Service enjoyed the largest quarterly increase 

in customer satisfaction ever recorded for a Federal government Web site, due mainly to the worldwide popularity of the 

VolcanoCam.

4

 Capitalization of “Volcano” indicates adoption of the word as part of 

the formal geographic name by the host country, as listed in the Geographic 

Names Information System, a database maintained by the U.S. Board on Geographic Names. Noncapitalized “volcano” is applied informally—eds.
11. Remote Camera Observations of Lava Dome Growth at Mount St. Helens, Washington, October 2004 to February 2006 227

A

B

SOUTH RIM

BRUTUS

GUACAMOLE

SUGAR

BOWL

CRATER

Seismic

onset

10/18/1

2004 2005 2006

12/1 2/1 4/1 6/1 8/1 10/1 12/1 2/1

A

B

Brutus

Crater

South Rim

Sugar Bowl Guacamole

VolcanoCam

KILOMETERS

SHW

YEL

46.2°

0 3

46.25°

46.15°

122.25° 122.2° 122.15°

Figure 1. Field setting for U.S. Department of Agriculture Forest 

Service VolcanoCam at Mount St. Helens, Washington. A, Camera 

position beneath eave of Johnston Ridge Observatory. USGS 

photo by J.P. Griswold, August 25, 2006. Inset shows camera. 

USGS photo by S.P. Schilling. B, Example of VolcanoCam view of 

the volcano acquired on September 24, 2004.

Figure 2. Remote, telemetered digital cameras used during 

monitoring of Mount St. Helens, Washington, 2004–6. A, Map 

showing locations of U.S. Geological Survey remote cameras 

(red dots) and U.S. Department of Agriculture Forest Service 

VolcanoCam (yellow dot). Black dots show locations of seismic 

stations that are referred to in figures 10 and 11. Hillshade-relief 

base map is from digital elevation model (DEM) of October 2005. B, 

Timeline with dates of operation (red bars) for remote, telemetered 

digital cameras through February 2006. Usable images were lacking 

on about half of all operating days, owing to inclement weather.

Despite its relative simplicity, the VolcanoCam was a 

remarkably useful educational resource and volcano-monitoring tool. Many of the thousands of emails received by the Forest Service in late 2004 regarding the VolcanoCam were from 

teachers across the United States, offering their thanks for the 

opportunity to view volcanic activity in their classrooms. In 

addition, the VolcanoCam proved to be a valuable tool for volcanologists, enabling rapid assessment of volcanic activity and 

weather conditions from any location having Internet access. 

The camera also demonstrated limited infrared capabilities. 

Nighttime observations were important for detecting magma 

extrusion and rockfall events, and they garnered substantial 

interest from the general public. An independent Web site 

managed by Mr. Darryl Luscombe even made available daily 

movies from sequential “glow” images collected during the 

previous night (http://www.luscombe-carter.com/index.html, 

last accessed January 28, 2008).

Remote, Telemetered Digital Cameras

In early October 2004, the value of having a visual monitoring station close to the volcano became obvious because 

the VolcanoCam’s view of the locus of renewed activity was 

blocked by the 1980–86 lava dome. To meet this need, staff 

at the USGS Hawaiian Volcano Observatory constructed a 

remote, telemetered digital camera, based on models used at 

Soufrière Hills volcano, Montserrat (Herd and others, 2005), 

and sent it to the Cascades Volcano Observatory (CVO) for 

deployment at Mount St. Helens. The system included an 

Olympus C–3030 3.3-megapixel camera with a × 3 optical 

zoom lens. The camera was connected through a serial port to 

a 900-MHz radio mounted in a weatherproof box (fig. 3). The 
228 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

box was fastened to a tripod and pointed towards the deforming area in the southeast part of the crater (fig. 4). Power was 

supplied by a solar panel and batteries with enough capacity 

to ensure that the camera and radio would operate even during 

long periods of cloudy weather. Image resolution, zoom, and 

timing of acquisition were controlled from a computer located 

at the Forest Service’s Coldwater Ridge Visitor Center, about 

13 km northwest of the crater, using PhotoPC public domain 

software (http://www.lightner.net/lightner/bruce/photopc/

ppc_use.html, last accessed January 28, 2008). Images were 

time stamped according to the camera time. The controlling 

computer could be reached via ftp from CVO through a satellite link, thereby providing access to imagery in near real time.

The remote camera was installed on October 10, 2004, 

on Sugar Bowl dome, immediately northeast of the breach in 

the 1980 crater wall and 2.3 km from the intensely deforming area, or welt (Dzurisin and others, this volume, chap. 

14), in the southeast part of Mount St. Helens crater (fig. 2). 

The goals for the Sugar Bowl camera deployment were to (1) 

establish a visual record of volcanic activity, which could be 

used to test inferences drawn from geophysical, geological, 

and geochemical measurements, (2) monitor volcanic activity 

in near real time, and (3) provide a means of assessing general 

conditions in the volcano’s crater to support field operations. 

Sugar Bowl offered a good view of the welt and subsequent 

dome growth (fig. 4B) from a point relatively safe from the 

mild explosive activity that characterized the early stages of 

the eruption.

A few problems resulted in a loss of imagery from the 

digital camera. Although high winds minimized snow accumulation, rime ice built up when temperatures were below freezing, obscuring the camera’s view (fig. 5). The ice was removed 

manually during site visits, but it often persisted for weeks at a 

time when no field work was conducted. The system functioned well during the period October 2004 to February 2006, 

with only a few lapses in image acquisition (fig. 2B) caused by 

mechanical breakdowns and abrasion of the viewing window by blowing volcanic ash. When the camera was operating, cloudy or icy weather resulted in no usable imagery for 

approximately half of the total deployment time.

The Sugar Bowl camera became an important tool in 

monitoring, research, and public outreach efforts, and it motivated the deployment of four additional instruments (Crater, 

Brutus, South Rim, and Guacamole; fig. 2) by the end of 2005. 

These new systems used similar equipment and software as the 

Sugar Bowl camera (fig. 6A). The Crater camera was installed 

by a helicopter sling operation within a few hundred meters 

of the growing dome on January 14, 2005. A location close to 

the dome was selected to provide close-up images that might 

be used to test dome-growth models (fig 6B). This camera 

suffered a mechanical failure several hours after it was put into 

Camera

Viewing window

Radio

Figure 3. Interior of weatherproof box that contains camera and 

radio telemetry for Sugar Bowl remote, telemetered digital camera 

system. USGS photo by M.P. Poland, October 10, 2004.

Figure 4. Setting and view for Sugar Bowl remote telemetered 

camera at Mount St. Helens, Washington. A, Field site atop Sugar 

Bowl dome, 2 km from active vent. Solar panel is out of view to the 

right. USGS photo by M.P. Poland, October 10, 2004. B, Example of 

camera view, acquired on February 10, 2005, showing spine 4.

SPINE 4

1980–86

LAVA DOME

EAST CRATER

WALL

SOUTH CRATER WALL

BATTERY BOX

CAMERA HOUSING

TELEMETRY ANTENNA

A

B
11. Remote Camera Observations of Lava Dome Growth at Mount St. Helens, Washington, October 2004 to February 2006 229

place and was subsequently destroyed by a small explosion 

during the early morning of January 16, 2005. During the summer and fall of 2005, additional cameras were installed on the 

crater rim: Brutus (fig. 6C), 1.1 km east-northeast of the vent, 

and South Rim (fig. 6A, D), 0.7 km southwest of the vent. As 

dome building focused in the southwestern part of the crater 

during late 2005, a camera was established on the floor of the 

breach in the 1980 crater: Guacamole (fig. 6E), 2.6 km north 

of the vent. Taken together, these camera systems provided a 

variety of different views of the growing lava dome.

Insights from Remote Camera Imagery

The remote cameras provided important, and sometimes 

unexpected, insights into volcanic activity at Mount St. Helens 

during 2004–2006. For example, the VolcanoCam confirmed 

that the extrusion of lava had begun in October 2004. Visual 

and infrared observations from a helicopter on October 11, 

2004, noted a craggy, hot (maximum temperature of 580°C), 

rocky “fin,” indicating that lava had reached the surface (Scott 

and others, this volume, chap. 1; Vallance and others, this 

volume, chap. 9). During the night of October 11 and the 

morning of October 12, the VolcanoCam showed signs of glow 

reflected off steam in the vicinity of the new spine, providing a 

valuable supplement to the earlier visual and infrared data and 

accessible to anyone with Internet access.

Both the VolcanoCam and the Sugar Bowl remote camera also had excellent views of explosive activity at Mount 

St. Helens. VolcanoCam photos posted to the Internet every 

five minutes provided useful, though approximate, constraints 

on the duration, magnitude, and timing of the early October 

2004 explosions. Following that period, only two additional 

Figure 5. Rime-ice accumulation on Sugar Bowl telemetered 

digital camera at Mount St. Helens, Washington. Ice buildup was 

a common problem on camera systems during fall and winter 

months. USGS photo by S.P. Schilling, October 24, 2004.

significant explosions occurred, on January 16 and March 8, 

2005 (Scott and others, this volume, chap. 1; Moran and others, this volume, chap. 6). The January 16 explosion occurred 

shortly after 0300 Pacific standard time (PST, Greenwich 

mean time minus 8 hours) during a period of poor weather 

in the middle of the night and was not visible to either the 

VolcanoCam or Sugar Bowl systems. In contrast, the March 8 

event took place at approximately 1725 PST during a time of 

clear weather (Scott and others, this volume, chap. 1; Moran 

and others, this volume, chap. 6). Analysis of Sugar Bowl 

imagery proved useful for the interpretation of seismic and 

acoustic data recorded during the event (Moran and others, 

this volume, chap. 6).

Visual imagery from remote cameras was useful in the 

recognition and analysis of rockfall from the lava spines. 

Although background glow from the growing dome had 

been observed in VolcanoCam imagery starting on the night 

of October 11, 2004, brief, brighter flashes were noticed by 

Internet observers beginning on January 13, 2005 (such flashes 

probably occurred earlier than this date but were not observed 

because of either their low intensity or poor weather). These 

images prompted seismologists to review the overnight seismic records and led to the recognition that the flashes were 

associated with rockfall signals. A major VolcanoCam flash 

occurred at about 0303 PST on February 22, 2005, and was 

accompanied by a large seismic signal (fig. 7). Visual inspection by field crews on the following day recognized a new scar 

on the growing lava dome, confirming the occurrence of a 

large rockfall during the previous night.

Significant rockfall events during daylight hours were 

accompanied by bursts of ash that often drifted above the crater 

rim (Moran and others, this volume, chaps. 2 and 6). Combining imagery from the remote, telemetered digital cameras, 

which was available within minutes of acquisition, with realtime seismic data allowed for rapid recognition of the rockfall 

source. An example occurred on April 26, 2005, at approximately 1126 Pacific daylight time (PDT, Greenwich mean time 

minus 7 hours), when a part of spine 4 disintegrated, sending a 

small ash plume above the crater rim (fig. 8).

The volume of extruded lava at Mount St. Helens during 

2004–2006 was calculated every 1–2 months by differencing 

digital elevation models (DEMs) derived from aerial photography or lidar data (Schilling and others, this volume, chap. 

8). More frequent, but necessarily qualitative, estimates of 

the relative rate of lava extrusion could be made by examining time-lapse sequences acquired by the remote, telemetered 

digital cameras. For example, in December 2004, a marked 

decline in the release of seismic energy (Moran and others, 

this volume, chap. 2) suggested that the eruption was slowing. When a sequence of daily images from the Sugar Bowl 

remote camera was reviewed, however, it became clear that 

the overall rate of lava extrusion had not changed significantly across the lull in seismicity. The measurement of extrusion rates can be quantified by combining a DEM with the 

remote camera imagery, as demonstrated using data from the 

Sugar Bowl camera by Major and others (this volume, chap. 
230 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 6. Other remote-camera setups and views at Mount 

St. Helens, Washington. Numbers on photos refer to spines as 

defined by Vallance and others (this volume, chap. 9). A, South 

Rim site, which also exemplifies equipment and installation style 

of Crater, Brutus, and Guacamole cameras. USGS photo by S.P. 

Schilling, August 19, 2005. B, Example of view from Crater camera, 

acquired January 15, 2005. C, Example of view from Brutus 

camera, acquired September 20, 2005. D, Example of view from 

South Rim camera, acquired August 19, 2005. E, Example of view 

from Guacamole camera, acquired February 24, 2006. Pink streak 

in middle of image is caused by sun damage to camera.

12). In an attempt to assess whether dome extrusion occurred 

smoothly or by a series of irregular surges correlative with 

seismicity, high-rate photography of a small field of view of 

patches on the growing lava dome was performed, but several 

factors limited the success of this experiment (Dzurisin and 

others, this volume, chap. 14).

Evolution of the Dome Complex Shown 

by Animations of Camera Imagery

Perhaps the most useful aspect of remote camera 

observations during 2004–6 at Mount St. Helens was the 

4

SOUTHWEST

CRATER WALL

B

SOLAR PANEL

CAMERA HOUSING

TELEMETRY ANTENNA

BATTERY BOX

A

6

5

1980–86

LAVA DOME

CRATER WALL

WEST CRATER

GLACIER

C

7

WEST CRATER

GLACIER TERMINUS

1980–86

LAVA DOME

E

WEST CRATER

GLACIER

WEST CRATER

GLACIER TERMINUS

1980–86

LAVA DOME

WEST CRATER

GLACIER

CRATER

WALL

5

CRATER WALL

4

6

D

6
11. Remote Camera Observations of Lava Dome Growth at Mount St. Helens, Washington, October 2004 to February 2006 231

02:56:38 02:59:31 03:02:24

PACIFIC STANDARD TIME (hh:mm:ss), FEBRUARY 22, 2005

03:05:17 03:08:10 03:11:02

02:58:36 03:03:36 03:08:36

Figure 7. Use of nighttime images from the U.S. Department of Agriculture Forest Service VolcanoCam to track rockfalls at Mount St. 

Helens, Washington. Top, consecutive images from February 22, 2005 (PST), showing one of the largest rockfalls of that year. Bright 

patch is a reflection in steam clouds of incandescence created by the sudden exposure of hot material. Outline of Mount St. Helens and 

1980–86 lava dome provided for context. VolcanoCam clock was not synchronized to Internet time, so it is probably only accurate to plus 

or minus 1 minute. Bottom, seismic record from station SHW (see fig. 2A for station location).

documentation of the lava dome complex’s morphological 

evolution over time. During clear weather and rime ice-free 

conditions, and regardless of the presence of field personnel, 

the remote, telemetered cameras provided high-quality views 

of the volcano from common vantage points. Imagery was 

therefore directly comparable over time, and visualizing the 

changing morphology of the lava dome complex by animating images into time-lapse movies proved to be an important 

tool for monitoring and interpreting volcanic activity.

Time-lapse animations of images from the Sugar Bowl, 

Brutus, South Rim, and Guacamole remote camera are provided as supplementary digital data to this report (Major and 

others, this volume, chap. 12, appendix 1, found on the DVD 

accompanying the volume and on the Web version of the 

work). Below, we describe and interpret the time-lapse animations of dome growth at Mount St. Helens obtained from 

the remote, telemetered digital cameras during the period 

October 2004 to February 2006. This account relies heavily 

on the Sugar Bowl camera for observations during the first 

year of the eruption, when that was the only remote camera 

that had been deployed. The observations that follow are 

drawn solely from remote camera imagery and do not rely on 

other data. The account is not meant to supplant but rather to 

complement descriptions of dome growth derived from other 

types of observations and data that are contained elsewhere 

in this volume. The chronology of 2004–6 activity is reported 

in this volume by Scott and others (chap. 1), Schilling and 

others (chap. 8), Vallance and others (chap. 9), and Herriott 

and others (chap. 10). In addition, geophysical and geochemical time series from the eruption are summarized by Moran 

and others (chaps. 2 and 6), Lisowski and others, (chap. 15), 

LaHusen and others (chap. 16), Gerlach and others (chap. 

26), and Pallister and others (chap. 30).

Between October 2004 and February 2006, dome 

growth at Mount St. Helens occurred through the extrusion 

of seven distinct spines (Scott and others, this volume, chap. 

1; Vallance and others, this volume, chap. 9). Spines 1 and 

2, formed in mid-October 2004, were the smallest of the 

extrusions, and they were active for the shortest periods of 

time. Owing to their location along the south margin of the 

1980–86 lava dome, they were not visible to the Sugar Bowl 

camera or the VolcanoCam and were documented only by 

observations (including thermal imagery) from helicopter 

overflights. The growth of spines 3–7, however, was visible 
232 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 8. Seismic record from station YEL (bottom) and consecutive images from Sugar Bowl remote camera (top) for April 26, 2005 

(PDT), showing major rockfall (source identified by red circle). See figure 2A for location of YEL.

from the Sugar Bowl camera except for a brief interval during 

the growth of spine 6.

The Sugar Bowl camera was deployed after the welt had 

largely formed in the southeastern part of the crater. During the first week of image acquisition, the camera recorded 

growth of a small knob that protruded from the welt along the 

southeast margin of the 1980–86 lava dome. The knob disappeared during October 20–27, 2004, a period of inclement 

weather when no visual observations (either by remote camera 

or field personnel) were possible. Judging from oblique aerial 

photos, the bulge appeared to be crater-floor debris and ice 

that was pushed up and later collapsed during the initial stages 

of dome extrusion (J. Major, written commun., 2006).

Spine 3 first became apparent in Sugar Bowl camera 

imagery on October 29, 2004, when uplift of the welt accelerated rapidly. The spine continued to grow steadily towards the 

southeast until mid-December, when imagery from December 

17 showed that cracks had formed along the north side of 

the extrusion. These cracks continued to develop throughout 

the remainder of the month, eventually leading to the formation of an independent spine of lava (spine 4). The breakup 

of spine 3 may have been caused when it impinged upon the 

Ch. 11, Poland, Figure 8

11:23:02 11:24:29 11:25:55 11:27:22

PACIFIC DAYLIGHT TIME (hh:mm:ss),

APRIL 26, 2005

11:28:48 11:30:14 11:31:41

1 2 3 4

1 — 11:23:45 2 — 11:26:21

3 — 11:28:56 4 — 11:31:26
11. Remote Camera Observations of Lava Dome Growth at Mount St. Helens, Washington, October 2004 to February 2006 233

southeastern crater wall (Vallance and others, this volume, 

chap. 9). The apparent extrusion velocity at the vent did not 

seem to change over this time period, judging from the photographic sequence from the Sugar Bowl camera and calculations of lineal extrusion rates (Major and others, this volume, 

chap. 12). If extrusion was constant, spine 3 was undergoing 

horizontal compression as lava continued to extrude from the 

vent, which probably caused it to fracture (Moran and others, 

this volume, chap. 2).

Spine 4, which was also characterized by dominantly 

southeastward motion, first rose to a higher altitude than 

spine 3, then began to override the latter in early January 

2005. Sugar Bowl imagery suggests that the smooth surface of spine 4 began to fracture and disintegrate sometime 

between March 15 and April 13, 2005. A more definitive 

date is difficult to assign, because the view from Sugar Bowl 

was obscured by ice and clouds between those two dates, but 

field observations show that disintegration began about April 

10 (Vallance and others, this volume, chap. 9). The increasingly fractured, but still coherent, spine continued to move to 

the southeast until April 24, 2005, by which time significant 

motion (that is, motion on the order of meters per day) had 

ceased and extrusive activity shifted from spine 4 to spine 

5. Major and others (this volume, chap. 12) document this 

transition using quantitative lineal extrusion rates based on 

Sugar Bowl imagery.

The Sugar Bowl camera observed the initial formation of 

spine 5 between April 14 and 18, 2005, when upward motion 

and lineal extrusion rate (Major and others, this volume, 

chap. 12) of that spine became independent of spine 4. Spine 

5 extruded at a steep angle (about 60°–70° from horizontal) 

from the vent and was subject to two cycles of construction 

and destruction during its life. Construction dominated until 

May 13, 2005, when a large part of the spine collapsed. Spine 

disintegration competed with extrusion to keep the spine at a 

relatively constant height from that time through June 4, 2005, 

when a second period dominated by construction began. By 

July 1, 2005, spine 5 had reached its highest altitude, although 

a period of more frequent collapses began around June 30, 

2005. A few tens of meters of dome elevation were lost 

between July 1 and 3, 2005. Upward motion of spine 5 continued, but the highest altitude of the extrusion remained nearly 

constant between July 3 and 14, 2005, as collapses from the 

upper part of the spine compensated for the addition of new 

lava at its base. By July 15, 2005, destructive processes began 

to outpace spine construction, and the height of the spine 

decreased daily. Sugar Bowl imagery suggests that the second 

cycle of growth and destructive phases of spine 5 had mostly 

ceased by August 2, 2005.

The distinction between spines 5 and 6 is difficult to 

constrain, but Sugar Bowl images indicate that spine 6 was 

moving independently of spine 5 by August 1, 2005. Growth 

of spine 6 was mostly vertical until August 10, 2005, when 

it began to move to the west without increasing in height. 

Sugar Bowl imagery and lineal extrusion rates (Major and 

others, this volume, chap. 12) suggest an apparent acceleration in westward motion starting on about August 16, 2005, 

and the spine height began to decrease as large collapses 

destroyed its upper reaches. A consequence of the motion of 

spine 6 towards the west was the development of a depression between spine 6 and the mostly inactive spine 5. Continued extrusion to the west occurred throughout the remainder 

of August and September, with the spine’s motion becoming 

almost completely horizontal. During this time period, the 

westernmost part of spine 5 gradually slumped into the growing depression, probably because it was left unsupported as 

spine 6 moved to the west (Vallance and others, this volume, 

chap. 9). Interestingly, unlike other spines, spine 6 apparently did not experience an extended period of collapse and 

destruction towards the end of its activity. This may have 

been a result of its relatively low height, compared to spines 

3, 4, and 5.

Growth of spines 6 and 7 was well documented by the 

Brutus, South Rim, and Guacamole remote cameras, which 

were installed during late summer 2005. Sometime between 

September 28 and October 17, 2005 (a period of poor 

weather when few observations were possible), spine 6 gave 

way to spine 7, which grew out of the western side of the 

depression between spines 5 and 6. The direction of spine 7’s 

motion was also toward the west but included a significant 

component of upward motion. As a result, spine 7 pushed 

and overrode spine 6, obscuring the distinction between the 

two extrusions. Although poor weather characterized much 

of late 2005 and early 2006, limited imagery indicates that 

spine 7 continued to grow into February 2006 with two 

cycles of alternating height increase (when the spine was 

gravitationally stable) and decrease (when the spine disintegrated gradually).

Time-lapse animation sequences from the remote, 

telemetered cameras reveal that spines 3 to 7 each experienced cycles of growth and destruction that lasted several 

months. The growth stages generally involved the extrusion 

of smooth-sided spines (with the exception of spine 6, which 

was mostly covered by rubble) with little accompanying 

large disintegration events. During destructive phases, spine 

extrusion continued, but abundant rockfall destroyed the 

smooth carapaces and resulted in highly fractured and blocky 

formations surrounded by talus. The onset of the destructive 

phase preceded the transition to a new spine in the cases of 

spines 3, 4, and 5. Spines 5 and 7 both experienced multiple 

constructive and destructive phases, perhaps related to their 

steeper extrusion angles. The “great spine” at La Montagne 

Pelée, Martinique, which was similar in appearance to the 

Mount St. Helens spines, also experienced multiple construction and destruction cycles during 1902–3, although the 

cycles appear to have been related to an unsteady, pulsing 

eruption rate (Jaupart and Allègre, 1991; Tanguy, 2004). At 

Mount St. Helens, lineal extrusion rates derived from Sugar 

Bowl imagery suggest that the eruption rate during the extrusion of spine 5 was nearly constant (Major and others, this 

volume, chap. 12); thus, alternating cycles of spine-height 

increase and erosion must have been controlled by other fac-
234 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

tors, for example, the strength of the dome carapace, thermal 

cooling, propagation of fractures, or gravitational stresses. 

Vallance and others (this volume, chap. 9) discuss the history 

and driving mechanisms of spine construction and destruction and the transitions between spines.

In addition to lava-dome processes and morphology, 

deformation of the Crater Glacier—which surrounded the 

1980–86 lava dome on the east, west, and south before the 

onset of eruptive activity in 2004 (Schilling and others, 

2004)—was recorded by several of the remote cameras. The 

Brutus camera’s field of view included the contact between 

the western part of the dome complex and the west arm of 

the glacier. By the time the camera was installed in midSeptember 2005, the west arm of Crater Glacier had already 

been extensively compressed, thickened, and fractured. The 

Brutus sequence of images showed continued thickening 

and cracking of glacial ice as spines 6 and 7 grew toward 

the west. A complementary perspective was provided by the 

Guacamole camera, which had a view of much of the glacier’s west arm (including its terminus) and recorded glacier 

deformation from the time of its installation in mid-November 2005. Motion of the glacier’s terminus occurred at an 

accelerated rate between January 23 and February 16, 2006, 

perhaps because of a downstream-moving bulge caused by 

compression of the glacier by spines 6 and 7 (Walder and 

others, this volume, chap. 13).

Strategies for Future Deployments of 

Remote Camera Systems

The bulk of the contributions from visual observation systems to monitoring efforts at Mount St. Helens in 

2004–2006 are necessarily qualitative because of limitations in camera views and weather conditions. Images were 

generally used to support inferences drawn from geophysical and geological observations or to characterize transient 

events and long-term processes. As demonstrated by Major 

and others (this volume, chap. 12), however, quantitative 

measurements of surface change from single camera deployments are possible.

Future camera deployments at active volcanoes should 

take advantage of photogrammetric principles, which will 

allow for more detailed analyses of surface change. For 

example, oblique aerial photographs that include ground 

control points with known positions can be used to construct 

DEMs of the ground surface. The technique has been demonstrated in laboratory conditions (Cecchi and others, 2003), at 

small scales on active lava flows at Mount Etna (James and 

others, 2006), and at larger scales on an entire lava dome at 

Soufrière Hills volcano, Montserrat (Herd and others, 2005). 

Expanding the use of photogrammetry to terrestrial cameras 

can be accomplished by deploying a pair (or more) of remote, 

telemetered digital cameras with views that are separated by 

30°–60° in azimuth from the target area and include several 

ground control points. A DEM of the areas viewed in common 

by a pair of cameras can then be constructed, and displacements, perhaps on the order of centimeters, may be calculated 

by differencing DEMs from different time periods. Although 

the principles involved in deriving such DEMs are not new, 

they have yet to be applied extensively using ground-based 

cameras. Terrestrial systems, although limited by weather conditions, offer the benefits of low cost, low power, flexibility, 

and high temporal and spatial measurement density. Domebuilding eruptions characterized by steady topographic change 

over time, like the 2004–6 activity at Mount St. Helens, offer 

an excellent opportunity for developing terrestrial photogrammetric systems.

Conclusions

Remote camera systems have provided important information regarding volcanic activity at Mount St. Helens during 

2004–6. A Webcam and a network of remote, telemetered 

digital cameras observed rockfalls, explosive activity, and the 

steady extrusion of lava on a nearly continuous basis, interrupted only by periods of inclement weather and infrequent 

mechanical failures. Time-lapse animations from the remote, 

telemetered digital cameras are outstanding records of lava 

dome emplacement that can be used to aid interpretations of 

volcanic activity and support education and outreach efforts.

Acknowledgments

The authors would like to thank the staff of the Hawaiian Volcano Observatory, particularly Rick Hoblitt (now at 

the USGS Cascades Volcano Observatory), Tim Orr, Richard 

Herd (now at the University of East Anglia in the United 

Kingdom), and Kevan Kamibayashi, for constructing and 

testing the first of the Mount St. Helens remote, telemetered 

cameras and sending the instrument to CVO for deployment at 

Sugar Bowl dome. Remote camera deployment was conducted 

with the assistance of Gene Iwatsubo, Stephanie Konfal, and 

Kirstie Simpson; Jon Major maintained the camera after it 

was installed and analyzed the imagery. Seismic data used 

in figures 7 and 8 were provided through the Incorporated 

Research Institutions for Seismology by the Pacific Northwest 

Seismic Network. Noise-filtered images from the VolcanoCam 

used to produce figure 7 were provided by Darryl Luscombe; 

David Ramsey generated the DEM that formed the backdrop 

for figure 2A. Jane Takahashi, Richard Herd, and Seth Moran 

provided valuable reviews.
11. Remote Camera Observations of Lava Dome Growth at Mount St. Helens, Washington, October 2004 to February 2006 235

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Mount St. Helens, 2004–2005, chap. 6 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Orr, T., and Hoblitt, R.P., 2006, Monitoring Kilauea Volcano 

using non-telemetered time-lapse camera systems [abs.]: 

Eos (American Geophysical Union Transactions), v. 87, no. 

52, Fall Meeting Supplement, Abstract G53A-0875.

Pallister, J.S., Thornber, C.R., Cashman, K.V., Clynne, M.A., 

Lowers, H.A., Mandeville, C.W., Brownfield, I.K., and 

Meeker, G.P., 2008, Petrology of the 2004–2006 Mount St. 

Helens lava dome—implications for magmatic plumbing 

and eruption triggering, chap. 30 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Pierson, T.C., 1985, Initiation and flow behavior of the 1980 

Pine Creek and Muddy River lahars, Mount St. Helens, 

Washington: Geological Society of America Bulletin, v. 96, 

no. 8, p. 1056–1069.

Schilling, S.P., Carrara, P.E., Thompson, R.A., and Iwatsubo, 

E.Y., 2004, Posteruption glacier development within the 

crater of Mount St. Helens, Washington, USA: Quaternary 

Research, v. 61, no. 3, p. 325–329.

Schilling, S.P., Thompson, R.A., Messerich, J.A., and Iwatsubo, E.Y., 2008, Use of digital aerophotogrammetry to 

determine rates of lava dome growth, Mount St. Helens, 

Washington, 2004–2005, chap. 8 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Scott, W.E., Sherrod, D.R., and Gardner, C.A., 2008, Overview of the 2004 to 2006, and continuing, eruption of 

Mount St. Helens, Washington, chap. 1 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Tanguy, J.-C., 2004, Rapid dome growth at Montagne Pelée 

during the early stages of the 1902–1905 eruption; a reconstruction from Lacroix’s data: Bulletin of Volcanology, v. 

66, no. 7, p. 615–621, doi:10.1007/s00445-004-0344-z.

Topinka, L., 1992, Basic photography at Mount St. Helens 

and other Cascade volcanoes, in Ewert, J. and Swanson, D., 

eds., Monitoring volcanoes; techniques and strategies used 

by the staff of the Cascades Volcano Observatory, 1980–90: 

U.S. Geological Survey Bulletin 1966, p. 195–217.

Vallance, J.W., Schneider, D.J., and Schilling, S.P., 2008, 

Growth of the 2004–2006 lava-dome complex at Mount St. 

Helens, Washington, chap. 9 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Voight, B., 1981, Time scale for the first moments of the May 

18 eruption, in Lipman, P.W., and Mullineaux, D.R., eds., 

The 1980 eruptions of Mount St. Helens, Washington: U.S. 

Geological Survey Professional Paper 1250, p. 69–86.

Voight, B., Glicken, H., Janda, R.J., and Douglass, P.M., 1981, 

Catastrophic rockslide avalanche of May 18, in Lipman, 

P.W., and Mullineaux, D.R., eds., The 1980 eruptions of 

Mount St. Helens, Washington: U.S. Geological Survey 

Professional Paper 1250, p. 347–377.

Walder, J.S., Schilling, S.P., Vallance, J.W., and LaHusen, 

R.G., 2008, Effects of lava-dome growth on the Crater 

Glacier of Mount St. Helens, Washington, chap. 13 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano 

rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).

Watts, R.B., Herd, R.A., Sparks, R.S.J., and Young, S.R., 

2002, Growth patterns and emplacement of the andesitic 

lava dome at Soufrière Hills Volcano, Montserrat, in Druitt, 

T.H., and Kokelaar, B.P., eds., The eruption of Soufrière 

Hills Volcano, Montserrat, from 1995–1999: Geological 

Society of London Memoir 21, p. 115–152.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 12

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, PO Box 51, Hawaii National Park, HI 96718

Extrusion Rate of the Mount St. Helens Lava Dome 

Estimated from Terrestrial Imagery, 

November 2004–December 2005

By Jon J. Major1, Cole G. Kingsbury1, Michael P. Poland2, and Richard G. LaHusen1

Abstract

Oblique, terrestrial imagery from a single, fixed-position 

camera was used to estimate linear extrusion rates during 

sustained exogenous growth of the Mount St. Helens lava 

dome from November 2004 through December 2005. During 

that 14-month period, extrusion rates declined logarithmically 

from about 8–10 m/d to about 2 m/d. The overall ebbing of 

effusive output was punctuated, however, by episodes of fluctuating extrusion rates that varied on scales of days to weeks. 

The overall decline of effusive output and finer scale rate 

fluctuations correlated approximately with trends in seismicity 

and deformation. Those correlations portray an extrusion that 

underwent episodic, broad-scale stick-slip behavior superposed on the finer scale, smaller magnitude stick-slip behavior 

that has been hypothesized by other researchers to correlate 

with repetitive, nearly periodic shallow earthquakes.

Introduction

Aerial and terrestrial photography are effective ways of 

monitoring morphological changes that occur during 

volcanic eruptions (see, for example, Zlotnicki and others, 

1990; Yamashina and others, 1999; Baldi and others, 2000, 

2005; Honda and Nagai, 2002; Herd and others, 2005; Thompson and Schilling, 2007; Poland and others, this volume, chap. 

11; Schilling and others, this volume, chap. 8). Some camera 

deployments and photography campaigns are aimed chiefly 

at monitoring volcanic activity qualitatively, whereas others 

endeavor to gather photographs sufficient to make quantitative 

(photogrammetric) measurements of static features or dynamic 

processes. For example, aerial photography has been used to 

estimate volumes of volcanic deposits and lava domes, to measure magnitudes of edifice deformation, to estimate volumetric loss during dome collapse, and to model development of 

volcano glaciers (for example, Moore and Albee, 1981; Jordan 

and Kieffer, 1981; Zlotnicki and others, 1990; Sparks and others, 1998; Schilling and others, 2004; Herd and others, 2005; 

Schilling and others, this volume, chap. 8). Similarly, groundbased photographs have been used to provide dynamic data 

of active processes, such as velocity estimates of large debris 

avalanches and lateral blasts (Voight, 1981), short-term (days) 

growth rates of lava domes (Yamashina and others, 1999), 

and motion of lava flows (James and others, 2006), and for 

geometric reconstructions of volcanic stratigraphy (Dungan 

and others, 2001). Commonly, stereoscopic imagery is used in 

such analyses, but apparent parallax caused by movement of 

an object in repeat photographs from a fixed position has also 

been exploited.

During the 2004–5 eruption of Mount St. Helens, oblique 

terrestrial imagery from remotely stationed cameras was one 

of the chief methods for monitoring the nature and pace of the 

eruption (Poland and others, this volume, chap. 11). Cameras 

were deployed principally to monitor the eruption visually 

without exposing scientists to unnecessary risk and to provide ancillary information on conditions in the crater (such as 

weather, the amount of steaming, or blowing ash) for purposes 

of planning field work. They were not deployed specifically 

for photogrammetric purposes.

Quantitative analysis of oblique terrestrial imagery 

commonly requires stereoscopic imagery or well-controlled 

nonstereoscopic imagery from multiple camera positions 

(Wolf and Dewitt, 2000). From October 2004 through December 2005 and beyond (Poland and others, this volume, chap. 

11), however, we obtained repeat, ground-based imagery of 

the Mount St. Helens eruption from a single, fixed-position 
238 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

camera located near the mouth of the volcano’s crater. After 

August 2005, other ground-based cameras were located 

elsewhere around the volcano (Poland and others, this volume, 

chap. 11), but none of those provided stereoscopic imagery. 

In this paper, we employ a methodology for quantifying the 

linear extrusion rate of a growing silicic lava dome from the 

imagery obtained by the camera deployed near the crater 

mouth, summarize the results of spatial and temporal variations of dome growth and extrusion rates from November 

2004 through December 2005, and compare our results with 

corresponding time series of seismic-energy release and 

local deformation measured by continuous Global Position 

System (GPS) receivers (Moran and others, this volume, 

chap. 2; LaHusen and others, this volume, chap. 16). Previously, average extrusion rates of silicic lava domes have been 

determined for discrete, short-lived eruptions or eruptive 

episodes (for example, Huppert and others, 1982; Swanson 

and others, 1987; Nakada and others, 1995; Sparks and others, 

1998), although extrusion rates of a continuous, decadal-scale 

eruption have been measured by Rose (1987) and by Harris 

and others (2003). Demonstration of fine-scale temporal and 

spatial rate variation during a long-term, continuous extrusion is rare, however (for example, Sparks and others, 1998). 

Our data provide further insights on fine-scale behavior of 

sustained silicic dome growth, and they provide constraints 

for mathematical models that elucidate the physics of dome 

growth (for example, Barmin and others, 2002; Melnik and 

Sparks, 2005; Iverson and others, 2006; Iverson, this volume, 

chap. 21; Mastin and others, this volume, chap. 22).

The Sugar Bowl Camera

To photographically monitor dome growth at Mount 

St. Helens, we deployed a camera at the northeast end of the 

volcano’s crater, on the Sugar Bowl lava dome, about 2.3 km 

from the locus of eruptive activity (fig. 1A). We used an Olympus 3030Z camera having a serial port (fig. 1B).

The camera utilizes a 1/1.8 inch solid-state sensor and 

has a nominal 6.5–19.5 mm focal length lens, equivalent to a 

32–96 mm lens on a 35-mm camera. The camera was programmed to expose images at its maximum focal length and 

to record them at 1,280×960 pixels (~1.2-megapixel images). 

The relatively low-resolution setting was intended to permit 

rapid capture and transmission of images of explosions or 

other dynamic phenomena. Communications with the camera 

and transmission of images were enabled via radio through 

the serial port. Images were transmitted to a base-station 

computer located at the Mount St. Helens National Volcanic 

Monument’s Coldwater Ridge Visitor Center, ~13 km downvalley (northwest) from the volcano. Once an image was 

transmitted, the camera acquired another image. The rate at 

which an image could be transmitted through the radio link 

governed the frequency of image acquisition, but in general 

images were acquired about every 3 minutes during daylight 

hours. Approximately once per hour the base-station computer 

transmitted an image to the Cascades Volcano Observatory. As 

needed, the computer could be queried to retrieve images more 

frequently. Periodically, all images were retrieved from the 

Coldwater Ridge Visitor Center.

The camera, deployed in a weatherproof housing (fig. 

1B), began operating on October 10, 2004 (Poland and others, 

this volume, chap. 11). It was replaced by another camera 

and remounted on a fixed pipe in February 2006 and replaced 

again in May 2006. Owing to episodic cloud cover, ice rime, 

pitting of the window glass, and unit failures, usable imagery 

is discontinuous over the period of deployment. In this paper 

we focus on imagery obtained from November 2004 through 

December 2005, roughly the first year of the eruption that 

began in October 2004.

Viewing window

Dessicant

Radio

 1980–86

lava dome

Camera housing

Battery box

Antenna

A

B

Figure 1. Photographs showing the camera system at Sugar 

Bowl lava dome. A, Deployed system at crater mouth. View is to 

southwest. Inset shows site location (white dot) on Sugar Bowl 

dome. Locus of eruptive activity is marked by emission of steam 

immediately to left rear of the partly snowclad 1980s lava dome. B, 

Camera and radio inside weatherproof housing.
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 239

Methodology

Theory of Terrestrial Image Analysis

The style of terrestrial imagery is defined on the basis of 

the orientation of the camera axis at the time of exposure. A 

horizontal terrestrial image is obtained when the camera axis 

is oriented horizontally at the time of exposure. If the camera 

is properly leveled before exposure, the x and z photographic 

axes are, respectively, oriented horizontally and vertically. 

Armed with precise information about camera and lens characteristics and an ability to pinpoint the optical center, or principal point, of a photograph, one can employ simple trigonometry to determine the horizontal and vertical angles between 

the camera axis and the rays to points in an object space (Wolf 

and Dewitt, 2000; fig. 2). If the camera is inclined from the 

horizontal, the resulting photograph is considered an oblique 

terrestrial image. In that case, computations of the angular 

differences between the camera axis, the rays to objects of 

interest, and a horizontal plane must account for the angle of 

inclination of the camera (Wolf and Dewitt, 2000).

The geometric relations among a camera’s position, L, 

its focal length, f, the principal point of the photograph, o, the 

inclination angle of the camera axis, h, and the horizontal and 

vertical angles to an image point are shown in figure 3. The 

horizontal angle, aa

, between the vertical plane containing 

image point a and the vertical plane containing the camera 

axis, Lo, is given by (Wolf and Dewitt, 2000):

sec/(tan )sin 1 a = a − zfx a  − . (1)

To conform to sign conventions, negative inclination angles 

refer to depression below, and positive inclination angles to 

elevation above, the horizontal (Wolf and Dewitt, 2000). In 

equation 1, correct algebraic signs must be applied to xa, za, 

and h. The vertical angle,ba, to image point a is given by

a = tan–1( za cos / ( f sec − za sin ) sec αa ) . (2)

In equation 2, the algebraic sign of b is automatically obtained 

from the sign of the za coordinate (Wolf and Dewitt, 2000).

Although equations 1 and 2 allow determination of 

the horizontal and vertical angles to any image point relative to the camera axis, they represent an underdetermined 

system of equations with respect to quantifying horizontal 

or vertical distances unless some geodetic control can be 

established between the camera and an object of interest. Typically, control is provided by interior and exterior 

referenced orientations. Interior orientation includes camera 

X

Z

x z

A

a

f

L

o

O

Y

 

1980–86 lava dome

Crater rim

2004–5 lava dome

Crater wall

Figure 2. Diagram showing 

relations among optical rays 

from photographed objects and 

images projected in the focal 

plane for a horizontal terrestrial 

image. For clarity, positions of 

the focal plane and lens nodal 

point have been reversed. 

Focal length of the lens, f, and 

principal point of the image, o, 

are shown relative to L, the lens 

nodal point or position of the 

camera. Projections of object 

point, A, and the horizontal (X) 

and vertical (Z) distances of 

the object point relative to the 

camera axis are shown on the 

projected image as a, x, and z, 

respectively. Object distance, 

Y, horizontal angle, a, between 

the vertical plane containing 

the object point and the vertical 

plane containing the camera 

axis, and vertical angle, b, from 

a horizon line to the object point 

are also shown.
240 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

+

+

Horizon line

a

a





L 

h

f

xa

z

za

x

a’ a’’

o 

a 

k

calibration parameters such as lens focal length, location of 

the image principal point, and corrections for lens distortion. Exterior orientation refers to the position and orientation of the camera with respect to a ground-based reference 

frame or with respect to the photographed object (Wolf and 

Dewitt, 2000; Molander, 2001). Orientation with respect to 

a ground-based reference frame is commonly accomplished 

by combining camera control (position and orientation) 

with object-space control (through established control 

points within the field of view).

In the absence of sound calibration parameters, one can 

approximate interior orientation of a digital camera from 

metadata contained in an image’s exchangeable image format 

(EXIF) file and by assuming that the principal point is at the 

center of the image. At maximum focal length, zoom lenses 

commonly have minimal distortion, even on consumer-grade 

cameras. Such assumptions, of course, can introduce large 

errors to a photogrammetric analysis.

In the absence of independently established control points, 

one way to establish exterior orientation of a fixed camera is by 

Figure 3. Diagram showing relations among optical rays from 

photographed objects and images projected in the focal plane 

of the camera for an oblique terrestrial image (slightly modified 

from Wolf and Dewitt, 2000; diagram copyright The McGraw-Hill 

Companies, Inc.). For clarity, the positions of the focal plane and 

lens nodal point have been reversed (see fig. 2). Focal length of 

the lens, f, principal point of the image, o, and angle of inclination 

of the camera axis, h, are shown relative to L, the lens nodal point 

or position of the camera, and a horizontal line (+x ). The projection 

of object point A (not shown) and its position relative to the 

photograph coordinate system (+x, +z) are shown on the projected 

image as a, xa, and za, respectively. Line Lk is a horizontal line that 

intersects the photograph at point k. Line aa ́ is a vertical line, with 

a ́ located in a horizontal plane. Points h and a’’ represent geometric 

projections used in the derivation of equations 1 and 2. Angles aa

and ba

 are the horizontal and vertical angles of the image point, a, 

relative to horizontal and vertical axes. See Wolf and Dewitt (2000) 

for detailed discussions of geometric relations and derivations of 

equations 1 and 2.

measuring the orientation and position of the camera with respect 

to a photographed object. For the Sugar Bowl camera, we lack 

rigid calibration parameters. We do, however, have empirical calibration parameters from four similar cameras. In the absence of a 

solid camera calibration, we used averaged parameters obtained 

from the calibrations of those other cameras (table 1). We imposed 

exterior control on the imagery by measuring the camera orientation (direction of the camera axis and angle of inclination) and its 

location (using GPS), and we established its position relative to the 

proximal part of the actively growing dome by measuring coordinates and distances between the camera and the dome on sequential digital elevation models (DEMs) (Schilling and others, this 

volume, chap. 8; fig. 4). By fixing the distance between the camera 

and the near-vent area of the dome between sequential DEMs, 

knowing the camera orientation, and employing averaged lens 

characteristics, we solved all necessary trigonometric equations 

and roughly quantified magnitudes and rates of dome growth.

Resolving horizontal and vertical displacement rates 

within the focal plane of the image provided only apparent 

rates of extrusion, however. Quantifying more accurate linear 

Figure 4. Digital elevation model of Mount St. Helens for 

February 21, 2005, showing positions of Sugar Bowl camera 

and 2004–5 lava dome, and relation between camera axis and 

obliquity of extrusion, c. Measured distance is from camera to an 

identifiable feature on the dome. Trigonometric manipulation of 

this distance and the angular difference between that axis and 

the camera axis allowed computation of the distance along the 

camera axis.

plane

Camera axis

Measured distance

Camera

Photographic

Direction of γ extrusion

N

0 500 METERS
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 241

Table 1. Camera characteristics and orientation parameters for the Sugar Bowl imagery, Mount St. Helens, 

Washington.

extrusion rates of the dome required resolving trigonometric 

differences between the apparent direction of motion in the 

focal plane and the actual direction of motion. Again, by using 

sequential DEMs of the dome, we measured the angle of 

obliquity, c, between the focal plane and the principal direction 

of dome growth (fig. 4). Analysis of a three-dimensional pyramid formed by the principal motion vectors (fig. 5) revealed 

that the angle, d, between the vector of apparent motion in the 

focal plane (vector c in fig. 5) and the vector of actual motion 

(vector V in fig. 5) is related to the obliquity angle, c, measured on the DEMs by

)tan(tan 1  R  − = , (3)

where R is the ratio of the horizontal motion vector (vector b

in fig. 5) to the resolved motion vector, c, in the focal plane 

(fig. 5). From that relation, simple trigonometric relations 

can be solved to estimate the extrusion rate in the direction of 

dome growth. Our methods and analysis are, of course, subject 

to many possible errors (discussed below) and are limited to 

extrusion rates that exceed about 1 to 1.5 m/d (displacements 

of about 2 to 3 pixels per day; table 1). Nevertheless, they 

represent useful procedures for extracting quantitative information from relatively low-resolution images from a fixed position 

obtained from an uncalibrated (and now damaged) camera.

Tracking Distinctive Features

To estimate extrusion rates, we followed movements 

of distinctive features on the lava dome between selected 

images. These features consisted of sharp edges, intersecting fractures, fracture tips, spots, or other stable, distinctive 

markers that could be readily identified and that persisted 

through multiple images, typically from a few days to about 

two weeks. To begin, an initial image (November 10, 2004) 

was imported into graphical design software and its apparent 

principal point (the center of the image) identified. Coordinates of features of interest on the dome were then computed 

relative to that apparent principal point (fig. 6). Subsequent 

images (separated by roughly 24 hours when possible) were 

then imported and manually coregistered with the preceding 

image, and the positions of displaced features of interest were 

updated. Coordinates of distinctive features on the dome in 

proximity to the vent area were entered into a spreadsheet, and 

equations 1 and 2 solved for the employed camera parameters 

(table 1) and measured distances between the camera and the 

dome (the distance was updated with each new DEM). The 

solutions, along with elapsed times between photographs, 

provided a time series of apparent horizontal and vertical 

displacement rates for selected proximal points on the actively 

growing dome. Typical rates of apparent horizontal and vertical displacement, and standard deviations of those rates, were 

estimated by averaging apparent displacement rates of 3 to 10 

points per image. Apparent linear extrusion rates in the focal 

plane were determined by resolving the averaged horizontal 

and vertical displacement rates, and the standard deviation 

errors on those averaged rates were propagated to the resolved 

solution using standard methods (for example, Bevington and 

Robinson, 1992). An average extrusion rate in the direction of 

dome growth was determined by resolving the average apparParameter Value

Camera characteristics

Maximum sensor width 1 7.76 mm

Maximum pixels at 1.2 megapixel setting 1,280 (1,280 × 960)

Field of view (at maximum focal length) 21°±1°

Resolvable pixel footprint at nominal distance to dome ~ 0.65 m

Lens characteristics

Lens focal length (maximum) 1 20.8 mm±1.2 mm

Lens distortion (radial) 1 ≤ 10−4 mm

Camera orientation

Inclination angle 7°±0.5°

Axis azimuth 207°±0.5°

Spatial measurements

Camera location (UTM coordinates, zone 10, datum is WGS84) 5118137N, 563633E, ±10 m

Distance from camera to plane tangent to proximal dome 2,315 m ± 30 m

Obliquity of extrusion relative to focal plane 27°–53°±2°

1

 Values of these parameters represent averages obtained from empirical calibrations (using Photomodeler Pro 5) of four other Olympus C3030Z cameras.
242 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Crater wall

a

e d

c

V

b

Focal plane

1980–86

lava dome

Crater rim

2004–5 lava dome





Figure 5. Diagram of a three-dimensional pyramid showing relations among the apparent vector of the 

average rate of motion, c, within a vertical plane perpendicular to the camera axis, the actual vector of 

extrusion, V, and the angle of obliquity of extrusion, c. Triangle abc is oriented vertically and perpendicular 

to the camera axis, and triangle ebd is oriented orthogonally to triangle abc (that is, in the plane of the 

topographic digital elevation model). As a result, triangle ecV is a right triangle. Note that leg e is common 

to triangles ecV and ebd. Therefore, e = b tan c = c tan d. Solving this equality leads to tan d = (b/c) tan c, 

or d =tan-1(R tan c), where R = b/c. Thus, by measuring the angle of obliquity of extrusion, c, on a DEM, one 

can employ simple trigonometry to resolve the apparent average rate of motion in the focal plane to the 

estimated rate of extrusion in the direction of dome growth.

ent extrusion rate in the focal plane with respect to the angle of 

obliquity of extrusion.

Sources of Error

Several sources of error are inherent in a quantitative 

analysis of oblique imagery from the Sugar Bowl camera. A 

fixed source of error involves the quantification of lens characteristics, camera orientation, and spatial measurements (table 

1) that affect the interior and exterior control imposed on the 

imagery. More random is the operator error incurred during 

image analysis.

Lens characteristics used in this analysis represent average values obtained from calibrations of four similar cameras. 

Variations of 1.2 mm about the assumed focal length (table 1) 

cause ±7 percent variation of the averaged apparent extrusion rates we report. Calibrated radial distortion in imagery 

was essentially negligible (table 1) and was thus ignored in 

our analyses. Repeated measurements of the inclination and 

azimuth of the camera housing during field visits minimized 

orientation error (table 1), although frequent misalignments 

among sequential photographs show that wind and transient 

snow and ice loads caused minor variations in orientation.

Spatial measurement errors revolve around accuracy of 

the camera location, inherent errors in the DEMs, measurement of the distance from the camera to the active extrusion, 

and measurement of the obliquity of the extrusion with respect 

to the camera axis. Replicable GPS measurements of the 

camera location and coordinate determinations in a GIS of 

selected locations on the active extrusion limit the error of the 

distance between the camera and the active extrusion to about 

±30 m, or about 1-percent error on the measured distance 

(table 1). Inherent errors in the DEMs (Schilling and others, 

this volume, chap. 8) are small compared to other measurement errors and are thus ignored. For much of the period 

from November 2004 to August 2005, the dome grew along 

azimuths ranging from 150° to 170° (Schilling and others, this 
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 243

1

2

3 4 5

6

Figure 6. Example of photograph taken by the Sugar Bowl 

camera of the 2004–5 dome in the crater of Mount St. Helens, 

illustrating a distribution of points used to track motion. Because 

the image has been cropped, the principal point of the original 

photograph, which serves as the origin to which point coordinates 

(for example, 1 to 6) are referenced, is offset. In this image, the 

2004–5 dome is about 130 m tall. Photograph taken on March 6, 

2005, at 17:08:38 UTC.

volume, chap. 8; Vallance and others, this volume, chap. 9), 

about 30°–50° oblique to the focal plane of the camera (which 

had an azimuth of 117°; table 1). From about mid-August 

through December 2005, motion was chiefly westward (~270° 

azimuth), about 25°–30° oblique to the focal plane of the camera. Repeated measurements of these obliquity angles within 

a GIS constrained these values to within a few degrees (table 

1). Errors of ±2° about the measured obliquity angles cause ±2 

percent variation in the reported averaged extrusion rates.

Aside from assumptions about lens characteristics and 

estimates of the distance from the camera to the dome, our 

most critical sources of error revolve around selection of the 

photographic principal point, the accuracy to which sequential 

photographs were coregistered, and the accuracy with which 

moving points were identified and tracked. For purposes of the 

analysis reported here, the principal point to which all measurements are referenced was selected simply as the center of 

the first base image, and all subsequent images were coregistered to that principal point. Owing to minor variations in 

fields of view resulting principally from winds and snow and 

ice loads on the housing, the principal point of the base image 

was not always the principal point of subsequent images. The 

distinctive dome features we followed were identified at the 

pixel level where possible, but the relatively low resolution 

of the images and the sometimes challenging lighting made 

pixel-level identification frequently difficult. Manual identification of displaced point positions could thus be in error by 

a few pixels. Overall, the greatest errors in the analysis were 

introduced through assumptions about lens focal length and 

through manual image registration and feature tracking.

Given the several sources of error that potentially affect 

our analysis, quantifying the cumulative error on our estimates 

of extrusion is a daunting task. Instead, we computed standard 

deviations of mean vertical and horizontal displacement rates 

in the focal plane from a collection of points on the proximal 

part of the active extrusion. Using standard practices (for 

example, Bevington and Robinson, 1992), we propagated the 

standard deviations of those mean values to the resolved extrusion rates in the focal plane and carried the propagated errors 

over to the resolved extrusion rates in the direction of dome 

growth to approximate a magnitude of error associated with 

our estimates of extrusion rate.

Results

Episodic growth and disintegration of several solidified 

lava spines characterized emplacement of the 2004–5 dome 

(Schilling and others, this volume, chap. 8; Vallance and others, this volume, chap. 9). Two minor spines, not visible from 

Sugar Bowl, breached the surface in October 2004. From late 

October 2004 until mid-April 2005, dome growth proceeded 

chiefly through emplacement and disintegration of two recumbent spines (spines 3 and 4) dubbed ‘whalebacks’ (Schilling 

and others, this volume, chap. 8; Vallance and others, this volume, chap. 9; fig. 7; supplemental movie in appendix 1). The 

character of growth changed following disintegration of spine 

4 in mid-April 2005. More vertical, rather than recumbent, 

growth characterized spine 5 from mid-April until late July 

2005 (fig. 7; supplemental movie). Following disintegration 

of spine 5 in late July 2005, vertical growth diminished and 

horizontal motion intensified as a new spine (spine 6) emerged 

and migrated rapidly westward. This phase of development 

persisted into October 2005, as a graben opened along the 

central part of the dome complex (see fig. 7 and supplemental 

movie) and growth of spine 6 became more endogenous. From 

mid-October through December 2005, the rapid westward 

migration of spine 6 slowed and vertical growth of another 

spine (spine 7) became notable (fig. 7; supplemental movie).

Rates of extrusion and associated motion of assorted 

segments of the 2004–5 lava dome varied in time and space, 

but for extended periods extrusion occurred at nearly steady 

rates. The average rate of vertical displacement varied between 

about 1 and 4 m/d, with a central tendency toward 2–3 m/d 

from November 2004 through June 2005 during growth of 

spines 3, 4, and 5 (fig. 8A). From about late June to mid-July 

2005, the rate of vertical displacement of spine 5 slowed 

substantially, then increased from mid- to late July during a 

period of reinvigorated growth. Following a series of rockfalls 

in late July, vertical motion of spine 5 diminished, but another, 

though less vigorous, growth spurt occurred in early to midAugust as spine 6 emerged. In December 2005 the average 

vertical motion of spine 7 hovered around 1 m/d (fig. 8A). In 

contrast, the average rate of horizontal displacement (in the 

focal plane) during emplacement of recumbent spines 3 and 
October 11, 2004 December 28, 2004 

Spine 3

Spine 3 Spine 4

Spine 3

Figure 7. Time series of images from the Sugar Bowl camera illustrating growth, disintegration, and morphologic change of the new lava dome at October 2004 and January 2006.
February 10, 2005 April 19, 2005 

Spine 4

Spine 4

Spine 5

Spine 4

Spine 4

Spine 5

Figure 7.—Continued.
Figure 7.—Continued.

June 24, 2005 August 24, 2005 

Spine 4 Spine 4

Spine 4 Spine 4

Spine 5 Spine 5

Spine 5

Spine 5 Spine 6 
October 22, 2005 December 15, 2005 

Spine 4

Spine 4

Spine 5

Spine 7

Spine 4

Spine 5

Spine 5

Spine 7

Spine 4

Spine 5 

Figure 7.—Continued.
248 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

4 decreased progressively from about 5–7 m/d in November 

2004 to about 3 m/d by mid-March 2005, then to fractions of a 

meter per day by late July 2005 after vertical growth of spine 

5 became predominant (fig. 8B). Horizontal motion increased 

substantially in August (to >2 m/d) when growth of spine 6 

migrated sharply westward; in December 2005, apparent horizontal motion of spine 7 hovered around 2 m/d.

Resolving the average horizontal and vertical displacement rates of the various spines into average rates of motion in 

the focal plane and then correcting those measurements for the 

angle of obliquity of extrusion shows that estimated extrusion 

rates declined logarithmically from November 2004 through 

December 2005 (fig. 8C). This decline is comparable to 

changes in magma discharge documented through analyses of 

DEMs (Schilling and others, this volume, chap. 8). However, 

several spurts of accelerated extrusion are superposed on the 

overall diminishing extrusion rate, and for several months lava 

extruded at a nearly steady rate (fig. 8C).

The estimated average linear extrusion rate was as great 

as 8–10 m/d in early November 2004 during emplacement 

of spine 3, hovered between 4 and 5 m/d from late December 2004 through mid-March 2005 during growth of spine 4, 

dropped to less than 2 m/d by mid-July 2005 during the waning stages of growth of spine 5, and then increased to about 4 

m/d before spine 5 disintegrated in late July 2005. Following 

disintegration of spine 5, the extrusion rate again dropped to 

as low as 2 m/d before accelerating to about 4 m/d in August, 

when spine 6 emerged and migrated sharply westward. In 

December 2005, during growth of spine 7, the extrusion rate 

had again declined to about 2 m/d. GPS receivers deployed on 

spines 3 and 4 tracked extrusion rates in real time for limited 

periods (LaHusen and others, this volume, chap. 16). During the periods when usable imagery and GPS deployment 

coincided, estimates of extrusion rate compiled from imagery 

analyses are about 20–30 percent lower than the measured 

rates (fig. 8C). Given the assumptions in our analyses, our 

estimates of linear extrusion rate compare relatively favorably 

with measured rates.

Motion of the extruded lava varied spatially as well as 

temporally. Movement of the distal ends of recumbent spines 3 

and 4 slowed as those spines enlarged, plowed over fragments 

of earlier spines, and impinged upon the crater wall. Differential rates of motion between distal and near-vent segments of 

those spines prompted development of thrust faults, large-scale 

fractures, and partial to complete disintegration in December 

2004, January 2005, and April 2005 (fig. 7; supplemental 

movie; Schilling and others, this volume, chap. 8; Vallance 

and others, this volume, chap. 9). Such disintegration generated hot rockfalls and minor pyroclastic surges but did not trigger substantive pyroclastic flows similar to those produced by 

dome collapses elsewhere (for example, Ui and others, 1999; 

Herd and others, 2005). Following the disintegration of spine 

4 (the largest of the recumbent spines) in mid-April 2005, the 

growth of spine 5 became focused along the western margin 

of the dome complex. During this phase of growth, active 

extrusion was effectively decoupled from the remnants of the 

earlier spines. From mid-April on, several remnants of spines 

3 and 4 remained stable as spine 5 emerged (for example, note 

the stability of the intact block of spine 4 on the east margin of 

the dome complex between April and December 2005; fig. 7; 

supplemental movie). Local displacement of ground near the 

vent, however, showed that, on occasion, extruding solidified 

lava was well coupled to the immediately surrounding terrain 

(see supplemental movie).

Following the disintegration of spine 5 in July 2005, the 

dome became further segmented. As spine 6 emerged and 

migrated westward, the eastern segment of the dome complex 

remained stable, but the central segment slumped as its westward buttress was removed (fig. 7; supplemental movie). Such 

spatially differential motion and migration of the extruding 

solidified lava led to the segmented morphology of the 2004–5 

dome (fig. 9), in contrast to the composite, but uniform morphology of the 1980s lava dome (see, for example, Swanson 

and others, 1987).

Extrusion rates estimated from the Sugar Bowl imagery 

are broadly correlated with overall trends in seismicity and 

deformation (fig. 10). From November 2004 until January 

2005, the rapidly decreasing extrusion rate was synchronous 

with a general decline in seismic-energy release as indicated 

by real-time seismic amplitude measurements (RSAM) (Endo 

and Murray, 1991; Murray and Endo, 1992; Moran and others, 

this volume, chap. 2) (figs. 10A, C). The RSAM values are 

commonly invoked as a proxy highlighting overall seismicenergy release at volcanoes (for example, Power and others, 

1994, 1995; Harlow and others, 1996; Mori and others, 1996; 

Voight and others, 1998). During phases of more or less steady 

extrusion from January to June 2005, minor increases in extrusion rate generally correlated with slight increases in RSAM. 

The extrusion-rate nadir and subsequent significant growth 

spurt in mid- to late July 2005 correspond with an equivalent 

nadir and subsequent minor increase in RSAM.

Correlations between trends in extrusion rate and geodetic measurements of deformation of a part of the 1980s 

lava dome are evident, but they are more subtle than are those 

with trends in seismicity (figs. 10B, C). We compare trends 

in extrusion rate with trends in the motion of station DOM1 

located on the west side of the 1980s lava dome north-northwest of the vent of the 2004–5 eruption (Dzurisin and others, 

this volume, chap. 14; LaHusen and others, this volume, chap. 

16). The overall logarithmic decay of the rate of lava extrusion 

is approximately mimicked by a corresponding logarithmic 

change in the northing component of motion of DOM1. Finer 

scale fluctuations in lava extrusion rate and motion of DOM1 

are, however, typically out of phase, particularly after about 

the first 5 months of the eruption. For example, an increase in 

extrusion rate between late April and mid-June 2005 corresponds with southward movement of DOM1 (that is, a relaxation of the 1980s dome toward the 2004–5 vent). From late 

June through mid-July 2005, an ensuing period of declining 

extrusion rate corresponds with a period of increased northward movement of DOM1 (that is, movement of the 1980s 

dome away from the vent). An ensuing increase of extrusion 
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 249

0

2

4

6

8

10

DATE

A

B

C

0

2

4

6

8

0

1

2

3

4

5

6

MEAN EXTRUSION RATE, IN METERS PER DAY MEAN APPARENT DISPLACEMENT RATE,

IN METERS PER DAY

MEAN APPARENT DISPLACEMENT RATE,

IN METERS PER DAY

2004 2005

NOV DEC JAN FEB MAR APR MAY JUN JUL AUG SEP OCT NOV DEC

Resolved extrusion rate

Horizontal displacement rate

Vertical displacement rate

Figure 8. Time-series plot of displacement rates and extrusion rates of the 2004–5 lava dome at Mount 

St. Helens. Error bars represent ±1 standard deviation about the mean. A, Average rate of vertical 

displacement within the camera focal plane. B, Average rate of horizontal displacement within the 

camera focal plane. C, Estimated average extrusion rate of the 2004–5 lava dome in the direction of 

growth. The gray bands illustrate periods when GPS receivers were deployed on the lava dome, and 

the triangles represent the average rates of motion measured by GPS (LaHusen and others, this volume, 

chap. 16). A logarithmic decay curve is superposed on the data.
250 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1980s lava dome

Stable dome

Deformed glacier

Deformed glacier

Active extrusion

Spine 4

Spine 5

Spine 6

Spine 7

rate in late July again corresponds with southward movement 

of DOM1 (toward the vent), whereas the decrease of extrusion 

rate in early August corresponds with increased northward 

motion of DOM1 (again, away from the vent). The increase of 

extrusion rate documented in mid-August 2005 is not as well 

out of phase with the motion of DOM1 as are other periods of 

changing rates of extrusion, but it does correspond to a gradual 

change from northward motion (away from the vent) to southward motion (toward the vent). Only the northing component 

of motion of DOM1 appears to broadly correspond with 

variations in extrusion rate; there is little if any substantive 

correlation between extrusion rate and the easting or vertical 

components of motion of DOM1 (figs. 10B, C).

Discussion

Within the constraints of the interior and exterior control 

we imposed, the available oblique, terrestrial imagery from 

the Sugar Bowl camera provided a valuable means of estimating long-term linear extrusion rates over periods of weeks to 

months during the 2004–5 eruption of Mount St. Helens. The 

greatest sources of error revolved around our assumptions 

regarding the camera focal length and our ability to accurately 

coregister sequential images and follow features of interest 

through time. Pitting of the glass on the camera box, shadows, 

steam and clouds, weather-related loss of usable images, and 

physical changes to the dome during its growth and partial 

disintegration all contributed to the challenge of using the 

Figure 9. Oblique aerial photograph looking southwest into the 

crater of Mount St. Helens. Photograph illustrates the segmented 

morphology of the 2004–5 lava dome complex. Remnants of spines 

4 and 5 are visible as discrete snowclad humps on the left and 

left-center of the image, and the locus of active growth, spine 7, 

is visible on the right side of the image. Fractured and deformed 

glacier ice is visible on the far left and far right sides of the image, 

and part of the snow-covered 1980s lava dome is visible in the 

bottom center of the image. USGS photograph by J.S. Pallister, 

December 6, 2005.

imagery. Although the locus of extrusion varied slightly during 

the period of analysis, errors in the imposed external control 

and measurements of the distance between the camera and 

the dome were relatively small compared to other sources of 

error. Indeed, the errors associated with the assumed focal 

length alone make the greatest difference in our comparisons 

of extracted versus measured extrusion rates.

Our methodology for extracting quantitative information from the fixed-position imagery hinged upon an ability 

to impose external control. In the absence of a time series of 

DEMs, or even a single DEM, we would not have been able 

to solve the underdetermined system of equations posed by 

having imagery from only a single viewpoint. Traditional 

surveying could have provided the necessary measures of 

distance from the camera site to the dome, but the dangerous environment in the crater during the early phases of the 

eruption, the distance between the camera and the lava dome, 

the difficulty of precisely placing prisms even remotely, and 

the rapidly changing physical character of the dome inhibited 

such a strategy.

The ability to quantify the rate of extrusion, even 

crudely, greatly enhanced the value of the imagery. Rather 

than simply serving as a method to monitor the status of the 

eruption or the conditions in the crater, the imagery supplemented other geophysical monitoring equipment. For the 

most part, it provided the sole means of extracting long-term, 

semicontinuous quantitative information. Differencing of 

sequential DEMs provided long-term estimates of magma 

discharge (Schilling and others, this volume, chap. 8), but 

those DEMs were based on stereoscopic aerial photographs 

only acquired about once per month. Hence, DEM differencing could not provide information on the fine-scale fluctuations of extrusion rate apparent in the camera imagery. Thus, 

quantification of linear extrusion rate of the lava from the 

Sugar Bowl imagery provided information about the eruption 

that was unattainable by other means.

Deployment of remote cameras at Mount St. Helens 

clearly enhanced documentation and analysis of the eruption (Poland and others, this volume, chap. 11). The semicontinuous imagery obtained from the Sugar Bowl camera 

was particularly useful for analyzing long-term, relatively 

fine-scale (days to months) variations of the linear extrusion 

rate. It proved less useful for examining extrusion rate over 

periods of hours, given the low average rates of movement, 

the challenges of accurately following features of interest, and 

the resolution of the pixel footprint at the distances involved. 

Hence, the scale of the imagery had a direct bearing on the 

quality and utility of the information that could be extracted. 

Finer resolution at shorter time scales requires more narrowly 

focused, larger scale imagery, but obtaining such imagery can 

be extremely challenging (Dzurisin and others, this volume, 

chap. 14). Of course, measurements of linear extrusion rate 

do not necessarily correlate directly with measurements of 

volumetric extrusion rate (that is, magma discharge). Nevertheless, long-term estimates and documentation of fluctuations 

of this one-dimensional parameter are useful and shed insights 
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 251

0

2

4

6

8

10

DATE

A

B

C

–0.2

0.2

0.4

0.6

0.8

0

0

100

200

300

MEAN EXTRUSION RATE, IN METERS PER DAY DISPLACEMENT, IN METERS REAL-TIME SEISMIC AMPLITUDE, IN COUNTS

2004 2005

NOV DEC JAN FEB MAR APR MAY JUN JUL AUG SEP OCT NOV DEC

Resolved extrusion rate

Station DOM1, L1 GPS

Station SHW

Northing

Easting

Vertical

Figure 10. Time-series plots of RSAM, geodetic measurements, and extrusion rate at Mount St. Helens. 

A, Real-time seismic-amplitude measurements (RSAM) from station SHW located on the west flank of 

the volcano. Plot has been smoothed using a five-point moving average. B, GPS geodetic measurements 

at station DOM1 located on the western side of the 1980s lava dome. The data, obtained using the 

L1-band frequency, are referenced relative to a station located in the upper South Fork Toutle River 

valley (LaHusen and others, this volume, chap. 16). Plot has been smoothed using an 11-point moving 

median. C, Estimated average extrusion rate of the 2004–5 lava dome. See figure caption 8C for details.
252 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

into overall eruptive behavior, especially when trends in linear 

extrusion rate can be correlated with trends in other geophysical measurements.

The pulsating behavior of lava extrusion documented 

on time scales of days to weeks in our analysis has been 

observed on time scales of hours to years during emplacement of other silicic lava domes. At Soufrière Hills volcano, 

Montserrat, West Indies, Sparks and others (1998) noted that 

magma discharge pulsated on time scales of hours to months 

during a time when the overall discharge rate was increasing. 

In contrast, the pulsating behavior documented here occurred 

during an overall trend of declining magma discharge (Schilling and others, this volume, chap. 8). Fluctuating extrusion 

rates during an overall decline of magma discharge were also 

observed at Unzen volcano, Japan, from 1991 to 1995 (Nakada 

and others, 1999). Lava at Santiaguito dome, Guatemala, 

has extruded continuously since 1922, with 3–5-yr spurts of 

rapid discharge interspersed with 10–15-yr periods of slower 

discharge (Rose, 1987; Anderson and others, 1995; Harris and 

others, 2003). The Mount St. Helens 1980–86 lava dome grew 

through a series of 20 eruptive episodes between October 1980 

and October 1986. However, each discrete extrusive episode 

was separated by pauses in eruptive activity that ranged from 

about one month to one year (Swanson and others, 1987), 

in contrast to the sustained, but fluctuating, extrusion that 

characterized the 2004–5 eruption. Sparks and others (1998) 

attribute pulsating extrusive behavior to a combination of deep 

and shallow volcanic processes. They speculate that discharge 

fluctuations on time scales of months to years are influenced 

chiefly by deep processes that control magmatic influx from 

the mantle and magma-chamber processes that affect magma 

ascent. In contrast, fluctuations on time scales of hours to 

weeks are attributed chiefly to pressure fluctuations caused by 

shallow-level processes, such as gas loss, crystallization kinetics, and mineral precipitation, which affect changes in magma 

properties and pressurization.

Correlations among seismicity, deformation, and extrusion rate during the 2004–5 Mount St. Helens eruption portray 

an extrusion that likely consisted of episodes of broad-scale 

stick-slip behavior. Relatively fine-scale correlations among 

periods of declining extrusion rate, relatively low seismicity, 

and northward movement of the 1980s lava dome (away from 

the eruptive vent) suggest episodes when the extruding solidified plug of lava was relatively well coupled to, and having 

difficulty evacuating, the conduit. Such a period of “stick” 

would be expected to result in lower seismicity, a low rate of 

extrusion, and increased deformation (that is, swelling) of the 

local terrain. In contrast, broad correlation among accelerated extrusion rate, slightly elevated seismicity, and southward movement of the 1980s lava dome (toward the eruptive 

vent) suggest episodes when the extruding solidified lava was 

relatively poorly coupled to, and slipping through, the conduit. 

Such periods of “slip” generated more or larger earthquakes as 

the lava dome lurched along (Moran and others, this volume, 

chap. 2) and allowed the local terrain deformed during the preceding period of stick to relax. Such episodes of broad-scale 

stick-slip movement complement a dynamic model of repetitive, fine-scale, stick-slip movement during sustained extrusion as proposed by Iverson and others (2006) and Iverson 

(this volume, chap. 21).

Long-term (months) and short-term (days) correlations 

between seismic intensity, deformation, and dome growth, 

such as documented here, have also been noted at other 

volcanoes. Broad correlations between seismic intensity 

and dome growth have been documented at Usu (Wano and 

Okada, 1980), Augustine (Power, 1988), Redoubt (Power 

and others, 1994), Unzen (Nakada and others, 1999), and 

Montserrat (Rowe and others, 2004), and short-term synchronicity and cyclicity between seismicity and deformation 

also have been observed (Voight and others, 1998). However, 

long-term, fine-scale correlations among fluctuations in 

extrusion rate, seismicity, and deformation generally have 

not been reported. Such correlations, as noted above, are 

clearly related to eruption mechanics. Indeed, Denlinger and 

Hoblitt (1999) have modeled short-term synchronicity and 

cyclicity in RSAM and deformation at Montserrat as a function of the interaction of volatile overpressure in magma and 

the overburden of an extruding lava dome, and Iverson (this 

volume, chap. 21) has hypothesized that the small, repetitive 

earthquakes that have occurred during the 2004–5 Mount St. 

Helens eruption, dubbed drumbeat earthquakes (Moran and 

others, this volume, chap. 2), reflect repetitive, small-magnitude (a few millimeters) stick-slip behavior of the extruding lava. Extraction of semicontinuous extrusion rates from 

long-term camera imagery, in conjunction with time-series 

of other geophysical data, clearly plays an important role in 

constraining dynamic eruption models and enabling forecasts 

of hazardous activity.

Conclusions

Imposition of interior and exterior controls on a semicontinuous series of oblique, terrestrial imagery from a fixed 

vantage point on the Sugar Bowl lava dome allowed quantification of fine-scale temporal behavior of the linear rate of lava 

extrusion during the 2004–5 eruption of Mount St. Helens. 

Analysis of the imagery showed that over a period of 14 

months (November 2004–December 2005), the linear extrusion rate varied in both space and time. Overall, the extrusion 

rate declined approximately logarithmically from about 8–10 

m/d in November 2004 to about 2 m/d by December 2005. 

However, the overall decline in the rate of extrusion was 

punctuated by fine-scale (days to weeks) fluctuations. The 

overall logarithmic decline of extrusion rate and the finer scale 

fluctuations correlated, approximately, with trends in seismicity and geodetic deformation. Those correlations portray an 

extrusion that underwent episodes of broad-scale stick-slip 

movement in addition to finer scale, smaller magnitude stickslip episodes that others hypothesize to correlate with small, 

so-called drumbeat earthquakes. The ability to extract linear 
12. Extrusion Rate of the Mount St. Helens Lava Dome Estimated from Terrestrial Imagery 253

extrusion rates from the imagery from this fixed-position camera provided a significant, and sometimes the sole, means of 

semicontinuously quantifying eruption dynamics during much 

of the first year of eruption, and those data provide an important constraint for dynamic eruption models.

Acknowledgments

As noted in this chapter, the Sugar Bowl camera was not 

specifically deployed with the intention of extracting quantitative information. However, when viewing the images, especially 

after they have been compiled into a time-lapse movie, one 

cannot help but realize that they contain a plethora of information about the growth, and the rate of growth, of the new lava 

dome. The trick is to extract that information. We thank our 

colleagues at the Hawaiian Volcano Observatory for providing 

us with the original camera and communications software; Gene 

Iwatsubo, Kirstie Simpson, and Stephanie Konfal for helping 

install the camera; Seth Moran for providing us with RSAM 

data; Steve Schilling for stimulating discussions as we muddled 

along through our analyses; Matt Logan and Dan Gooding for 

compiling the imagery into the movie included in the supplemental movie of appendix 1; our many colleagues at CVO who 

asked pointed questions at staff meetings when we unveiled our 

analyses; and Ren Thompson and Ricky Herd for insightful, 

critical reviews that kept us from wandering beyond the limits of 

what we could reasonably do in our efforts to extract information from these photographs.

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domes and lava flows: Geological Society of America Special Paper 212, p. 17–27.
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0273(90)90121-U.

Appendix 1. Time-Lapse Photography of Mount St. Helens, 2004–2006—Movie

[This appendix appears only in the digital versions of this work in the DVD-ROM that accompanies the printed volume and 

as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

The appendix is a time-lapse movie showing dome growth at Mount St. Helens from November 10, 2004, to May 10, 2006. 

The movie, in mpeg-1 file format and titled “Sugarbowl to May10 2006_5000.mpg,” is composed of 188 photographs taken by 

the Sugar Bowl remote camera. It was assembled by choosing the best image per day for times when weather was suitable for 

viewing the dome. File size is 23 Mb. Compiled by Matt Logan and Dan Gooding (USGS).
256 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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257 A 

Volcano A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 13

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Effects of Lava-Dome Growth on the Crater Glacier of 

Mount St. Helens, Washington

By Joseph S. Walder1, Steve P. Schilling1, James W. Vallance1, and Richard G. LaHusen1

Abstract

The process of lava-dome emplacement through a glacier 

was observed for the first time as the 2004–6 eruption of 

Mount St. Helens proceeded. The glacier that had grown in the 

crater since the cataclysmic 1980 eruption was split in two by 

the new lava dome. The two parts of the glacier were successively squeezed against the crater wall. Photography, photogrammetry, and geodetic measurements document glacier 

deformation of an extreme variety, with strain rates of extraordinary magnitude as compared to normal temperate alpine 

glaciers. Unlike such glaciers, the Mount St. Helens crater 

glacier shows no evidence of either speed-up at the beginning 

of the ablation season or diurnal speed fluctuations during 

the ablation season. Thus there is evidently no slip of the 

glacier over its bed. The most reasonable explanation for this 

anomaly is that meltwater penetrating the glacier is captured 

by a thick layer of coarse rubble at the bed and then enters the 

volcano’s groundwater system rather than flowing through a 

drainage network along the bed. Mechanical consideration of 

the glacier-squeeze process also leads to an estimate for the 

driving pressure applied by the growing lava dome.

Introduction

Since October 2004, a silicic lava dome has been 

emplaced first through, and then alongside, glacier ice in the 

crater of Mount St. Helens. The dome has been emplaced in a 

near-solid state, not as liquid magma solidifying at the Earth’s 

surface (Vallance and others, this volume, chap. 9). Heretofore, dome emplacement through a glacier was known only 

from a single published photograph (Simons and Mathewson, 

1955, plate 6) showing a lava dome that had been emplaced 

through the caldera glacier of Great Sitkin Volcano, Alaska, 

sometime in 1945. Evidence bearing on lava-dome emplacement into ice has been presented by, for example, Gilbert and 

others (1996), who used geophysical methods to identify lava 

domes emplaced beneath the caldera glacier of Volcán Sollipulli, Chile, and by Tuffen and others (2001), who described 

a domelike rhyolite body that was evidently emplaced subglacially in Iceland and since exhumed. The 2004–6 eruption 

of Mount St. Helens has afforded the first-ever opportunity 

to actually document the process of lava-dome emplacement 

through a glacier.

The common picture of volcano-glacier interactions is 

one of rapid meltwater generation either as magma contacts 

the glacier bed—examples from Iceland have been especially 

well characterized, for example, the 1996 Gjálp eruption 

(Guðmundsson and others, 1997)—or as lava or pyroclasts are 

erupted onto the glacier surface (many examples are mentioned by Major and Newhall, 1989). At Mount St. Helens, 

however, glacier melt associated with dome emplacement 

has been minor, even as the glaciological consequences have 

been dramatic—Crater Glacier has been cut in half, and the 

resulting ice bodies have in succession been squeezed between 

the growing lava dome and the crater wall. In this paper we 

focus our attention on the glaciological consequences of the 

eruption. Condensed discussions of this material have been 

presented elsewhere by Walder and others (2005, 2007).

Field Setting: Crater Glacier Before 

October 2004

After the cataclysmic eruption of May 18, 1980, which 

beheaded, and in some cases completely destroyed, the glaciers that existed on the flanks of Mount St. Helens (Brugman 

and Meier, 1981), material from rock and snow avalanches 

began accumulating in the north-facing, amphitheaterlike 

crater (fig. 1). Mills (1992) used digitized topographic maps to 
258 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

calculate the volumes of material eroded from the crater walls 

and accumulated on the crater floor. His results show that as 

of mid-1988, the thickness of accumulated material was 60 to 

80 m across much of the crater floor south of the 1980–86 lava 

dome. The accumulated material as of mid-1988 was about 60 

percent rock debris by volume and contained interstitial snow, 

but it was not flowing. The first reasonably clear evidence that 

a crater glacier had come into existence—the appearance of 

crevasses, which reflect flow—comes from photographs taken 

in September 1996 (Schilling and others, 2004). The glacier 

(now called Crater Glacier) at that time had a surface area of 

about 0.1 km2; by September 2000, this area had increased to 

about 1 km2. Proceeding similarly to Mills (1992) but using 

digital elevation models (DEMs), Schilling and others (2004) 

calculated that the material that had accumulated in the crater 

between May 18, 1980, and September 2000 had a thickness locally as great as 200 m and a volume of 1.2×108

 m3, of 

which about one-third comprised rock debris. If we interpret 

these figures in the context of Mills’ discussion of what had 

accumulated on the crater floor as of 1988, it seems clear that 

the deepest part of the crater-floor fill consists primarily of 

rock-avalanche debris—a point to which we shall return—and 

would not be considered glacier ice by usual glaciological 

standards. The uppermost part of Crater Glacier, however, 

probably contains no more than 5 percent rock debris by volume, with such debris forming discrete, discontinuous layers 

A

B

Figure 1. Oblique view of Mount St. Helens crater on October 5, 

2000, looking south. Crater Glacier wraps around 1980–86 lava dome. 

East (left) arm of glacier is obscured by rock-avalanche debris; west 

(right) arm merges to the north of the lava dome with a rock-covered 

icy mass shed off the west crater wall. Crater width, as indicated by 

double-headed arrow, is about 2 km. USGS photograph by 

Bergman Photographic Services, Portland, Oreg.

Figure 2. Glacier features in crater of Mount St. Helens. A, 

Rock layers within uppermost part of Crater Glacier, as exposed 

on west side of new lava dome. Distance from glacier surface 

to prominent debris layer (arrow) is about 3 to 5 m. View to east. 

USGS photograph taken August 4, 2005, by W.P. Johnson. B, 

Surface of Crater Glacier on August 20, 2003, looking north along 

east side of 1980–86 lava dome. Rock-avalanche lobe in center of 

view extended from the crater wall nearly to the south side of the 

lava dome and had maximum thickness of about 1 m. 

USGS photograph by J.S. Walder.

that originate as rock-avalanche lobes (fig. 2), and glaciologists would call this material “dirty” firn and glacier ice.

To what extent is the material accumulated on the crater 

floor since 1980 a glacier? In framing an answer, we have to 

make explicit our reason for asking the question in the first 

place. Our focus here is not on morphology, but rather on 

the ice flow and deformation processes familiar to glaciologists, and how such processes affect the mechanical response 

of the crater-fill material to lava-dome emplacement. From 

this perspective, what one is tempted to call a glacier in a 

morphological sense is not the same as what is rheologically 

and mechanically glacier ice. Deformation of a material 

containing 60 percent rock debris by volume—Mills’ (1992) 

estimate for the composition of the pre-1988 crater-floor 

fill—is surely dominated by rock-to-rock friction, not creep 

of any interstitial ice.

We choose to exclude from our mechanically defined glacier, as best we can, the deepest, rock-rich crater-floor fill. We 

do this by picking the glacier bed as the crater-floor surface 
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 259

defined by DEMs for October 12, 1986, and November 12, 

1986. This is an approximate but defensible choice for several 

reasons: (1) The rate of accumulation of rock debris in the 

crater decreased markedly after 1986 (fig. 3). (2) 1986 marks 

the end of the previous dome-growth episode, so accumulation after 1986 occurred within a basin with reasonably stable 

boundaries. (3) As we argue in appendix 1, interstitial ice 

within the lowest, rock-rich crater-fill material has probably 

melted and not been replaced by ice intruding from above. 

With the 1986 surface thus defined as the glacier bed, we then 

differenced 2003 and 1986 DEMs to calculate the glacier 

thickness shortly before the start of the 2004 eruption (fig. 

4). Using the Mills (1992) and Schilling and others (2004) 

figures for rock-debris accumulation, we estimate that Crater 

Glacier, so defined, has an average rock content of 15 percent 

by volume.

A note about names. The U.S. Board on Geographic 

Names on June 6, 2006, approved the name “Crater Glacier” 

for the feature that existed before the recent eruption. However, as is documented below, Crater Glacier has been split in 

two by dome growth, and it is both sensible and convenient to 

use the informal names “west Crater Glacier” and “east Crater 

Glacier” for the ice masses that exist in the crater as of the 

time of writing.

Changes in Crater Glacier Since 

October 2004

Methods

Hazards posed by Mount St. Helens’ eruptions severely 

restricted field work in the crater, so we documented eruptive effects on Crater Glacier primarily by photography and 

YEAR

1980 1985 1990 1995 2000 2005

 CUMULATIVE VOLUME ERODED,

 IN MILLION CUBIC METERS

10

0

20

30

40

50

Figure 3. Cumulative volume eroded from walls of Mount St. 

Helens crater since the 1980 eruption, as determined by differencing 

digital elevation models for 1990, 2000, and 2003 with a DEM for 1980. 

Error bars (±1σ) are shown. Despite uncertainties in the data, it is 

clear that erosion rate has fallen sharply since the mid-1980s.

photogrammetry. We managed to collect some glacier-motion 

data using single-frequency global positioning system (GPS) 

stations slung by helicopter onto the glacier surface in 2005 

and 2006. The GPS stations (LaHusen and others, this volume, chap. 16) were available for glacier monitoring only 

intermittently, and on several occasions had to be moved, or 

else they would have toppled into crevasses. Station positions 

were determined from short-baseline differential fixed static 

solutions sampled at 10-second intervals over a 25-minute 

period every hour. Accuracy of individual solutions was 

approximately 20 mm in the horizontal and 50 mm in the 

vertical. A running-median filter was applied to solutions to 

remove spikes.

Morphological Changes

One of the first indirect signs of dome growth was the 

formation of a bulge in the south part of Crater Glacier during 

the last few days of September 2004 (fig. 5). An explosion on 

October 1, 2004, excavated a hole in the glacier (fig. 6). As the 

eruption proceeded, the southern part of Crater Glacier was 

eventually punctured by a rock spine surrounded by rubble 

(fig. 7), the latter perhaps comprising unconsolidated mateFigure 4. Map showing thickness of material accumulated on 

the crater floor of Mount St. Helens between October–November 

1986 and September 2003. Background is a hillshade-relief map 

constructed from September 2003 digital elevation model. The 

1980–86 lava dome is in center. As explained in text, the October–

November 1986 surface is approximately the glacier bed, and the 

isopachs represent approximate glacier thickness.
260 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

B

C

Figure 5. Bulge in Crater Glacier next to south side of 1980–86 

lava dome on September 30, 2004. Dark material on surface of 

fractured area is talus. Width of bulge is about 50 m. View to west. 

USGS photograph by D. Dzurisin.

Figure 6. Beginning of Mount St. Helens eruption through Crater 

Glacier on October 1, 2004. View to west. USGS photograph by 

J.S. Pallister.

Figure 7. Upwarped, tephra-covered firn and ice around 

margins of new lava dome on October 11, 2004. View to northeast. 

A, Deformed rock at ambient temperature. B, Deformed firn and 

ice. C, Spine 1 (hot rock). USGS photograph by C.A. Neal.

rial that underlay the glacier. The lava dome as it exists as of 

October 4, 2006, is a complex of seven such spines extruded 

sequentially in the solid state from the same general vent area 

(fig. 8; Vallance and others, this volume, chap. 9). Spine 3, 

which began to be extruded in late October 2004, grew preferentially southward, developing a whaleback form and pushing 

aside firn and ice in a way reminiscent of the bow wave that 

precedes a ship through water (fig. 9). After spine 3 ran into 

the south crater wall in mid-November 2004, Crater Glacier 

was for all practical purposes split into two parts.

East Crater Glacier

Spine 3 spread to the east until late December 2004, then 

spalled greatly and was shouldered aside by spine 4—another 

“whaleback”—which grew until mid-April 2005. The east 

Crater Glacier (ECG) was effectively caught in a vise formed 

by the whaleback spines and the east crater wall. Owing to 

drought conditions that prevailed throughout most of the 

winter of 2004–5, there was practically no snow accumulation, and thus glacier-surface features showed very clearly. As 

eastward dome growth proceeded, the upwarped glacier apron 

on the east side of the dome (compare fig. 9) impinged against 

the east crater wall. However, the northernmost part of this ice 

apron was rotated until it formed a steplike feature trending 

nearly east to west (fig. 10). The ECG surface buckled, with 

east-west-trending crevasses forming parallel to the direction 

of dome spreading (fig. 11). Comparison of DEMs reveals that 

between mid-November 2004 and mid-April 2005, the dome/

ECG contact migrated laterally by as much as 200 to 250 m 

and the glacier locally doubled in thickness (figs. 12, 13, 14). 

Expressed in terms of rates, the dome-ECG contact moved on 

average about 1 m/d and the glacier thickened at an astounding 

0.6 m/d. By way of comparison, the average thickening rate 

for the “reservoir area” of a surging glacier, during the interval 

between surges, is perhaps 0.02 to 0.04 m/d (Raymond, 1987, 

p. 9123, fig. 1).

Since spine 4 quit growing in mid-April 2005, east 

Crater Glacier has thinned in its upper reach and thickened 

in its lower reach as normal flow processes redistribute ice 

mass downslope. Longitudinal crevasses became obvious by 

late April 2005; these crevasses probably reflect transverse 

spreading as the bowed-up surface—so evident during the 

squeezing episode—relaxed. As a result, the glacier surface 

became a field of seracs (fig. 15). The ECG terminus became 

steep (fig. 16) and advanced by about 150 m between April 

19, 2005, and August 18, 2006.
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 261

West

Crater Glacier West Crater Glacier

Spine 7

Spine 5

Spine 4

Spine

 6

Spine 3

1980–86 Dome

Deformed 1980s

dome and crater floor

Spine 7

Spine 5

Spine 4

Spine 3

Deformed 1980s

dome and crater floor

East

Crater Glacier East Crater Glacier

562000 562500 563000 563500

5116500

5116000

5115500

0 100 200 300 400 500 METERS

Figure 8. Map of the lava-dome spine complex in Mount St. Helens crater drawn on a hillshaderelief map from the October 24, 2005, DEM. Spines are numbered according to sequence of 

extrusion events, as discussed by Vallance and others (this volume, chap. 9). Coordinates 

referable to UTM zone 10, North American datum 1983. 

Figure 9. Upwarped firn and ice around margin of new 

whaleback lava spine, November 20, 2004. View to east. 

USGS photograph by S.P. Schilling.

1980–86

dome

Deformed

ice

Spine 3 Spine 4

Figure 10. The new lava dome of Mount St. Helens (dominated by 

spines 3 and 4) and the by-then morphologically distinct east Crater 

Glacier (in foreground) on January 14, 2005. The bulge indicated by 

the arrow is not a kinematic wave but was instead formed when 

upwarped ice around the spine margins (see figs. 7, 8) was rotated 

as dome growth proceeded to east. View to southwest. 

USGS photograph by J.W. Vallance.
262 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 11. Upwarped surface of east Crater Glacier on February 

16, 2005. View to north. Crevasses are oriented roughly east-west, 

paralleling direction of squeeze by new lava dome (at left). 1980s 

dome in left-center distance. USGS photograph by J.S. Walder.

Figure 12. Migration of the contact between rock of the new lava dome and ice of east Crater Glacier during the period from November 

29, 2004, to April 19, 2005. Contact position was determined from DEMs, with a probable error of about 5 m. Background image is 

hillshade-relief map for November 29, 2004. Coordinates are UTM zone 10 easting and northing, North American datum 1983. Eastward 

migration of rock-glacier contact for northing between about 5115500 and 5116000 reflects growth of new lava dome, which caused the 

glacier to thicken locally. The resulting enhanced ice flow to the north caused ice to encroach upon the margin of the old (1980–86) lava 

dome north of about northing 5116050. Also indicated are positions of four GPS stations deployed on the glacier at various times.

C

J K

H I

B

A

O

N

T U

R

S

M

Q

P

L

F G

D E

C

J K

H I

B

A

O

N

T U

R

S

M

Q

P

L

F G

D E

A

#

#

#

#

# #

#

#

#

HIE5

ICY4

ICY5

ELE4

ELE4

WES6

DAVF

JOEA

JOEC

HIE5

HIE5

ICY4

ICY5

ELE4

ELE4

WES6

DAVF

JOEA

JOEC

HIE5

B

Figure 13. Hillshade-relief maps of Mount St. Helens crater constructed from photogrammetric analysis of aerial photographs dated 

October 24, 2005. A, Lines of section for which we calculated changes in glacier-surface altitude. B, Positions of GPS stations. Note that 

any individual station may not have been on the glacier on the date of the photographs. ELE4 appears twice because it was shifted from 

east Crater Glacier to west Crater Glacier during the course of the eruption.

562640

5116200

5115800

563040

N

Nov. 29, 2004

Dec. 11, 2004

Jan. 03, 2005

Feb. 01, 2005

Feb. 21, 2005

Mar. 10, 2005

Apr. 19, 2005

0 150 300 450 METERS

GPS station

Position of

rock-ice contact
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 263

DISTANCE ALONG SECTION, IN METERS

0 200 400 600 800 1,000 1,200 1,400

SURFACE ALTITUDE, IN METERS

1,800

2,000

2,200

2,400

X

X

X

ELE4

ICY5

ICY4

X

HIE5

M N

O

L

A

April 19, 2005

July 14, 2005

December 15, 2005

Surface of east Crater Glacier

October/November 1986

September 2003

January 3, 2005

February 21, 2005

DISTANCE ALONG SECTION, IN METERS

0 100 200 300 400

SURFACE ALTITUDE, IN METERS

2,000

2,100

2,200

2,300

2,400

Q

P

B

EAST CRATER WALL

April 19, 2005

July 14, 2005

December 15, 2005

Surface of east Crater Glacier

October/November 1986

September 2003

January 3, 2005

February 21, 2005

DISTANCE ALONG SECTION, IN METERS

0 100 200 300 400

SURFACE ALTITUDE, IN METERS

1,900

2,000

2,100

2,200

2,300

C

S

R

EAST CRATER WALL

April 19, 2005

July 14, 2005

December 15, 2005

Surface of east Crater Glacier

October/November 1986

September 2003

January 3, 2005

February 21, 2005

DISTANCE ALONG SECTION, IN METERS

0 100 200 300 400 500

SURFACE ALTITUDE, IN METERS

1,800

1,900

2,000

2,100

2,200

2,300

U

T

D

EAST CRATER WALL

April 19, 2005

July 14, 2005

December 15, 2005

Surface of east Crater Glacier

October/November 1986

September 2003

January 3, 2005

February 21, 2005

Figure 14. Cross sections showing changes in surface altitude 

of east Crater Glacier during the course of current eruption, 

based on sequential DEMs. Lines of section and GPS station 

locations shown in figure 13. (Note that GPS station ICY4 was 

adjacent to 1980–86 lava dome and thus north of the part of east 

Crater Glacier that was squeezed.) The 1986 profile represents 

the ground surface at the end of the 1980–86 dome-building 

episode and approximates the glacier bed. The 2003 profile 

should be within a few meters altitude of the glacier surface 

at beginning of current eruption. Not all DEM coverages 

extend to glacier terminus. A, Longitudinal section L–M–N–O 

approximately parallel to ice flow. B, Transverse section P–Q. 

C, Transverse section R–S. D, Transverse section T–U.
264 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

B

Figure 15. Crevasses formed by lava-dome growth at Mount St. 

Helens. A, Longitudinal crevasses on east Crater Glacier cutting 

across transverse crevasses that had formed during eastward 

lava-dome growth (compare fig. 9), as seen on May 12, 2005. 

View to southwest. USGS photograph by M. Logan. B, Part of 

east Crater Glacier on July 26, 2005. View to south. Longitudinal 

crevasse growth by this date had effectively chopped the glacier 

surface into a field of seracs. USGS photograph by S.P. Schilling.

West Crater Glacier

Growth of spine 6 (Vallance and others, this volume, 

chap. 9) adjacent to west Crater Glacier (WCG) became 

noticeable in early August 2005. Surface bulging and crevassing of the glacier proceeded in much the same way as with 

ECG (fig. 17). Spine 6 quit growing, and spine 7 began growing and overriding spine 6, in early to mid-October 2005, but 

WCG continued to be squeezed owing to the push exerted by 

spine 7 on spine 6. Events unfolded much as with ECG: The 

dome-WCG contact migrated locally by >200 m (fig. 18), and 

the glacier locally doubled in thickness (fig. 19). A distinct 

bulge in the WCG surface began propagating downglacier (fig. 

19A) and impinged upon the rather flat, mostly rock-covered 

terminus region, which arguably originated as a separate mass 

shed from the west crater wall (compare fig. 1). In summer 

2006, it became clear that advance of the bulge was being 

accommodated by development of a shear zone within the flat 

terminus region (fig. 20).

Change in Ice Volume During the Eruption

The change in glacier volume during the course of the 

eruption can be determined by comparing DEMs prepared for 

different dates. The method is discussed in appendix 2, and 

results are summarized in figure 21. The estimated volume 

decrease from the start of the eruption (October 2004) until 

October 2005—meaning (approximately) from the end of one 

ablation season to the end of the next ablation season—was 

6.7±3.7×106 m3, corresponding to an average rate of loss of 

0.21±0.12 m3/s. The eruption has clearly not been marked by 

a process commonly associated with volcano-glacier interactions, namely, rapid meltwater generation (Major and Newhall, 

1989). In retrospect, this is unsurprising—the eruption has 

been predominantly quiescent, not explosive, so scouring of 

the glacier surface by hot fragmental flows has been negligible; moreover, the spines have been extruded in a solid 

state, with surface temperature well below the solidus, and the 

glacier is well insulated from them by rubble (Schneider and 

others, this volume, chap. 17).

Ice Dynamics

Given the radical morphological changes to Crater Glacier 

during the eruption, described above, we should not be surprised if the glacier’s dynamics were also significantly affected. 

Unfortunately, our complete lack of data on glacier-surface 

speed before the 2004 eruption complicates an assessment of 

how the eruption affected glacier dynamics. To try to infer a 

rough baseline for preeruption dynamics, we use mass-balance 

considerations to estimate the so-called balance velocity Ub

, 

which is the cross-sectionally averaged speed that a glacier 

would have if it were in steady state (Paterson, 1994, p. 250):

0

1 ( ) () () () ()

x

Ux b W d b W xH x=   ∫  , (1)

where W(x) is glacier width at distance x from the “headwall” 

or upstream end (in this case, the south crater wall), H(x)

is average depth at a cross section, and b(x)  is the local mass 
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 265

Figure 16. Terminus (lower center) of east Crater 

Glacier on June 15, 2005. View to south. Compare to 

indistinct terminus as seen about 5 years earlier (fig. 

1). Arrow indicates bulge similarly indicated in figure 

10. The glacier is only about 100 to 150 m wide where 

it passes between crater wall and old lava dome 

(right center) USGS photograph by S.P. Schilling.

C

A

B

WES6

HIE5

ELE4

West

Crater

Glacier

East

Crater

Glacier

0 100 200 300 400 500 METERS

562000 562500 563000 563500

5116500

5116000

5115500

Dec. 15

Oct. 24

Sept. 1

Aug. 10

Sept. 20

July 14

June 15

GPS station

Feb. 9

2006

Position of dome-

 glacier contact

2005

Figure 17. West Crater Glacier adjacent to westward-growing 

lava dome as seen on September 2, 2005. View to southeast. 

The crevasse pattern in the glacier is complicated and reflects 

shifting directions of dome growth, but those crevasses normal 

to the dome-glacier margin are the youngest. 

USGS photograph by M. Logan.

Figure 18. Migration of 

contact between the new 

lava dome and west Crater 

Glacier during the period 

June 15, 2005, to February 9, 

2006. Contact position was 

determined from DEMs, with 

a probable error of about 

5 m. Background image is 

hillshade-relief map for June 

15, 2005. Coordinates are 

UTM zone 10 easting and 

northing, North American 

datum 1983. As the new 

dome grew, the glacier 

encroached upon margin of 

the old (1980–86) lava dome. 

Also shown are positions of 

three GPS stations that were 

deployed on the glacier at 

various times in 2005.

balance expressed as a thickness per unit time. Equation 1 is 

simply a mathematical statement of the steady-state assumption, 

namely, that the glacier is neither thickening nor thinning. We 

apply equation 1 to what would become (during the eruption) 

east Crater Glacier and estimate Ub near the terminus (at a 

position we denote by x = L) by taking L ≈ 1.2 km, H ≈ 60 m

(fig. 14A), and an average value b

 ≈ 4 m/yr (from a total ice 

accumulation of about 80×106 m3 over an area of about 1 km2
266 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 19. Change in surface altitude of west Crater Glacier during course of ongoing eruption, based on sequential DEMs. Lines 

of sections shown in figure 13. The 1986 profile represents the ground surface at end of the 1980–86 dome-building episode and is 

approximately the glacier bed. The 2003 profile should be within a few meters altitude of glacier surface at beginning of current 

eruption. A, Longitudinal section A–B–C, approximately following the thickest ice. The points labeled A, B, and C match those in figure 

13A. B, Transverse section D–E. C, Transverse section F–G. D, Transverse section H–I. E, Transverse section J–K.

DISTANCE ALONG SECTION, IN METERS

0 50 100 150 200 250

SURFACE ALTITUDE, IN METERS

2,000

2,050

2,100

2,150

2,200

2,250

E

D

B

December 15, 2005

Surface of west Crater Glacier

August 10, 2005

September 20, 2005

October 24, 2005

October/November 1986

September 2003

April 19, 2005

January 3, 2005

F

G

DISTANCE ALONG SECTION, IN METERS

0 50 100 150 200 250 300

SURFACE ALTITUDE, IN METERS

2,000

1,950

2,050

2,100

2,150

2,200

2,250

C

December 15, 2005

Surface of west Crater Glacier

August 10, 2005

September 20, 2005

October 24, 2005

October/November 1986

September 2003

April 19, 2005

January 3, 2005

DISTANCE ALONG SECTION, IN METERS

SURFACE ALTITUDE, IN METERS

1,600

0 200 400 600 800 1,000 1,200 1,400 1,600 1,800

1,800

2,000

2,200

2,400

C

B

A

A

October 24, 2005

December 15, 2005

April 18, 2006

Surface of west Crater Glacier

October/November 1986

September 2003

September 20, 2005

July 14, 2005
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 267

H

I

DISTANCE ALONG SECTION, IN METERS

0 50 100 150 200 250 300

SURFACE ALTITUDE, IN METERS

2,000

1,950

2,050

2,100

2,150

2,200

1,900

D

December 15, 2005

Surface of west Crater Glacier

August 10, 2005

September 20, 2005

October 24, 2005

October/November 1986

September 2003

April 19, 2005

J

K

DISTANCE ALONG SECTION, IN METERS

0 50 100 150 200 250

SURFACE ALTITUDE, IN METERS

2,000

1,950

2,050

1,850

1,800

1,900

E

December 15, 2005

Surface of west Crater Glacier

August 10, 2005

September 20, 2005

October 24, 2005

October/November 1986

September 2003

April 19, 2005

JOEA

JOEC

Figure 20. The glacier in Mount St. Helens crater as seen on 

September 12, 2006. View to south. As bulge on west Crater Glacier 

advanced and impinged on relatively flat terminus area, a shear 

zone delineated by echelon fractures developed (solid red curve). 

The shear zone at its northern end took on the character of a zone 

of compression, with crevasses parallel to direction of maximum 

compression (dotted red lines). Positions of GPS stations JOEA and 

JOEC are indicated. USGS photograph by W.E. Scott.

DATE

1/1/03 1/1/04 1/1/05 1/1/06

CRATER ICE VOLUME,

IN MILLION CUBIC METERS

50

55

60

65

70

75

80

Relative to 1986 DEM

With estimated 2003–04 accumulation

Nominal value at

start of eruption

Figure 21. Total volume of glacier in Mount St. Helens crater as a 

Figure 19—Continued. function of time, with error bars (±1σ).

in 20 years), and by treating W as a constant. We find Ub ≈ 0.24 

m/d, which corresponds to a surface speed of about 0.29 m/d 

for ice with the flow-law exponent n = 3 (van der Veen, 1999, 

p. 103–106). We emphasize that this is at best a rough baseline 

for thinking about the preeruption surface speed, because the 

glacier was manifestly not in a steady state but rather growing.

East Crater Glacier

The GPS station positions during 2005 are shown in figure 13B; measured displacement rates are shown in figure 22. 

Interestingly, the balance velocity estimated above is comparable to the speed of station ICY4, which was downglacier of the 
268 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

domain squeezed by the lava dome, on ice about 70 m thick 

and not far from the ECG terminus (fig. 10). In comparison, 

ICY5, about 300 m upglacier of ICY4, within the glacier reach 

being squeezed by dome growth (fig. 12) and on ice about 150 

m thick, moved about 1.3 m/d, or about four times as fast as 

ICY4. This comparison presents a conundrum if deformation 

is only by simple shear and reflects a balance between gravitational driving stress and drag on the glacier bed and sides, 

in which case the difference in surface velocity between ICY4 

and ICY5 should have been a factor of about (150/70)n+1≈21 

for a flow-law exponent n = 3 (van der Veen, 1999, p. 103–

104). Moreover, owing to the nonlinear rheology of glacier ice 

(van der Veen, 1999, p. 13–15), the squeeze exerted on east 

Crater Glacier by the growing lava dome should have reduced 

the effective viscosity of the ice near ICY5 and made the 

difference in speed from ICY4 to ICY5 even greater. Resolution of the conundrum involves recognizing that gravitational 

driving stress is in fact resisted not only by drag but also by 

gradients in stress along the flow (van der Veen, 1999). A useful mechanical analogy is to think of east Crater Glacier, during the squeezing episode, as a tube of toothpaste with the cap 

removed. If the entire tube were squeezed uniformly, toothpaste would squirt out rapidly, but if squeezing is applied only 

to the part of the tube farthest from the opening, the toothpaste 

nearer the opening acts as a dam. Computational modeling by 

Price and Walder (2007) has confirmed the existence of a very 

strong longitudinal stress gradient.

Strain rates associated with ECG deformation can be 

estimated, in part, by considering the rate of eastward migration of the dome-glacier contact and the rate of glacier-surface 

uplift. Dividing the rate of eastward migration of the domeglacier contact near ELE4 (fig. 12) by the glacier width (about 

300 m), the average rate of contact migration for the period 

December 1, 2004, to January 3, 2005, corresponds to a 

squeeze strain rate of about −0.006/d; for the period January 

3, 2005, to April 16, 2005, the squeeze strain rate was about 

DAY OF YEAR

40

0

0.2

0.4

0.6

0.8

1.0

1.2

1.4

1.6

ICY4

ICY5

ELE4

HIE5

Kennicott Glacier

error

60 80 100 120 140 160 180 200 220 240

HORIZONTAL SPEED, IN METERS PER DAY

Figure 22 Horizontal speed of east Crater Glacier 

GPS stations. Locations of the stations shown in figure 

12. Raw position data were filtered to remove spurious 

spikes and interpolated to 0.2-d intervals. Estimated 

error is 0.05 m/d. ICY4 and ICY5 were on glacier in 

mid- to late winter 2005 while the new lava dome was 

expanding eastward. ELE4 was fortuitously placed on 

glacier about the time that dome growth to east stopped, 

and it stayed on the glacier until early summer 2005. 

HIE5 was on the glacier in mid-summer 2005. Azimuth 

of motion for all stations was within 18° of north. Shown 

for comparison are surface-speed data (adapted from 

Anderson and others, 2005) for a target on Kennicott 

Glacier, a temperate valley glacier in Alaska, during the 

year 2000. The record for Kennicott Glacier shows largeamplitude, commonly diurnal fluctuations not seen at 

east Crater Glacier.

−0.0036/d. Elongational strain rate in the downglacier direction cannot be estimated directly owing to the fact that there 

were never simultaneously two GPS units on the reach being 

squeezed. The strain rate associated with glacier thickening for 

the period January 3, 2005, to April 16, 2005, can be roughly 

estimated (see fig. 14A) at about (0.6 m/d)/(100 m) ≈ 0.006/d 

near the centerline of east Crater Glacier.

To put the ECG strain-rate values in perspective, consider 

ice moving through a valley constriction at a rate of 100 m/y, 

with the valley narrowing by 25 percent over a length of 1 

km—arguably a rather severe constriction. The lateral strain 

rate in this case would be −0.0001/d, or about 1–3 percent of 

the lateral strain rate associated with squeezing of the ECG. 

Thickening strain rate as large as that measured at ECG is 

known only from surge fronts (Kamb and others, 1985; Raymond and others, 1987), although in such cases the maximum 

compression is oriented along the normal ice-flow direction, 

whereas with ECG, maximum compression was transverse to 

the normal ice-flow direction.

West Crater Glacier

The GPS stations on west Crater Glacier during the summers of 2005 and 2006 (fig. 13B) recorded the response of the 

glacier to westward dome growth. We discuss results for 2005 

and 2006 separately.

In 2005 (fig. 23), the peak in speed of ELE4 at about day 

273 (September 30) occurred a few days before the appearance of spine 7 just east of spine 6 (fig. 8; Vallance and others, 

this volume, chap. 9). This peak in speed probably reflects 

a change in the stresses applied to WCG by the dome. During the 23-day period when the GPS records overlapped, all 

three stations on WCG accelerated rather smoothly (fig. 23B); 

differences in azimuth of motion reflect the local direction 

of dome growth. The displacement records for the overlap 
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 269

period were analyzed to determine direction and magnitude 

of the principal strain rates within the (approximately horizontal) plane determined by the three stations. Unsurprisingly, 

the direction of principal compression lined up closely with 

the trend of crevasses that formed during westward dome 

growth (fig. 17). Magnitudes of principal horizontal strain 

EASTING, IN METERS (zone 10, NAD83)

 NORTHING, IN METERS (zone 10, NAD 83)

231.62

313.19

231.75

257.37

209.62

257.42

ELE4

HIE5

A WES6

562,010 50 90 130 170 562,210

5,116,040

80

120

160

5,116,200

DAY OF YEAR 2005

190 210 230 250 270

GPS Station

290 310

HORIZONTAL SPEED, IN METERS PER DAY

0.0

0.5

1.0

1.5

2.0

2.5

WES6

HIE5

ELE4

Error B

DAY OF YEAR 2005

230 235 240 245 250 255 260

STRAIN RATE, IN 1/DAY

-0.010

-0.008

-0.006

-0.004

-0.002

0.000

0.002

0.004

0.006

0.008

.

.

. .

ε11

ε22

ε11+ε22

Error C

Figure 23. GPS-derived motion data for west Crater Glacier in 

2005. A, Displacement trajectories projected into horizontal plane, 

with day of year at beginning and end indicated. B, Horizontal 

speeds, calculated by filtering raw position data to remove 

spurious spikes, interpolating to 0.2-d intervals, and applying 

centered difference. Estimated error, shown by error bar, is 0.05 

m/d. As with the east Crater Glacier record (fig. 22), diurnal speed 

fluctuations are not seen at west Crater Glacier. C, Principal strain 

rates in horizontal plane. Orientation of maximum extension ( 11 ε ) is 

N. 10° E.–S. 10° W.; orientation of maximum compression ( 22 ε ) is N. 

80° W.–S. 80° E. For strictly incompressible ice, the sum 11 22 ε ε+  

would be zero.

rates increased slowly over time, with their sum consistently 

negative at about −0.002/d. Making the plausible interpretation that surface uplift represents thickening of the glacier, 

vertical strain rate can be estimated as the average uplift rate 

divided by the glacier thickness, or about (0.25 m/d)/(120 m) 

= 0.002/d. The sum of the three principal strain rates was thus 

locally near zero, consistent with bulk incompressibility.

In 2006, we had motion data for three GPS stations located 

on WCG downglacier of the region being squeezed by the 

lava dome: DAVF, which operated for several months (during 

which time the station was relocated three times to prevent it 

from toppling into a crevasse), and JOEA and JOEC, which 

operated for about six weeks during summer (fig. 24). Station DAVF was slightly upglacier of the cross section H–I (fig. 

13A), on ice that thickened steadily as west Crater Glacier was 

squeezed (fig. 19D). Although the motion record for DAVF (fig. 

24A) is complicated by the effect of crevasse growth and the 

need to move the instrument, there is again an absence of the 

diurnal speed variation we would expect if glacier sliding were 

occurring. The motion records for JOEA and JOEC (fig. 24B), 

located only about 150 m apart, nicely document deformation associated with the shear zone shown in figure 20. Station 

JOEA (east of the shear zone), on ice being shoved as the bulge 

in the WCG surface propagates downglacier and impinges upon 

the terminus region, moved nearly three times as fast as JOEC 

(west of the shear zone). The difference in azimuth of motion 

between JOEA and JOEC almost certainly reflects the opening 

of roughly north-south-striking crevasses (fig. 20).

Inferences About Glacier and Volcano 

Hydrology

Temperate glaciers (those with ice at the melting point) 

move by a combination of internal creep and sliding of the 

ice over the bed (van der Veen, 1999). The creep component 

reflects the internal stress state of the glacier rather than 

conditions at the bed, whereas the sliding component reflects 

the boundary condition at the bed—in particular, how much 

friction there is against the bed. Measurements at many 
270 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

DAY OF YEAR 2006

160 180 200 220 240 260

HORIZONTAL SPEED, IN METERS PER DAY

0.0

0.2

0.4

0.6

0.8

1.0

1.2

AZIMUTH, IN DEGREES

-45

-30

-15

0

15

30

45

Speed

Azimuth

A

DAY OF YEAR 2006

200 210 220 230 240 250 260

HORIZONTAL SPEED, IN METERS PER DAY

0.0

0.2

0.4

0.6

0.8

1.0

AZIMUTH, IN DEGREES

0

30

60

90

JOEA speed

JOEC speed

JOEA azimuth

JOEC azimuth

B

Figure 24. Motion data for GPS stations on the surface of west 

Crater Glacier in 2006. Estimated errors about 0.05 m/d for speed 

and 2 degrees for azimuth. See figure 13B for locations of the 

GPS stations. A, Horizontal speed and azimuth of GPS station 

DAVF. Dashed lines indicate breaks in data when instrument was 

moved to keep it from toppling into crevasses that formed during 

its stay. B, Horizontal speed and azimuth of GPS stations JOEA 

and JOEC. 

glaciers have shown systematic differences between surface 

speed during the ablation (melt) season and during winter. 

For example, pulses of increased surface speed are commonly 

observed as the melt season begins (Anderson and others, 

2004). More generally, surface speed in summer is higher than 

in winter, and large diurnal variations in surface speed are 

common (Fountain and Walder, 1998). As the creep component 

of glacier motion should be reasonably constant, variations 

in surface speed reflect variations in sliding speed, which is 

modulated by meltwater at the bed (see, for example, Harper 

and others, 2002). Our 2005 data for east Crater Glacier (fig. 

22), however, show neither acceleration with the onset of the 

melt season nor a clear diurnal signal; data for west Crater 

Glacier from the summers of 2005 and 2006 (figs. 23, 24) 

similarly lack any diurnal signal. We propose as an explanation 

that there simply is no pressurized drainage system conveying water along the bed. Crater Glacier grew atop several tens 

of meters of rubble (mainly rock-avalanche debris) that had 

accumulated on the crater floor following the eruption of May 

18, 1980 (Mills, 1992). As argued in appendix 1, much of this 

rubble is likely to be ice free because geothermal heat flow 

will have melted interstitial ice, and flow of the overlying ice 

downward into the rubble will have been slow. The volcanic 

edifice beneath this rubble is geologically complex, consisting of multiple lava flows, pyroclastic and lahar deposits, and 

other fragmental deposits (Crandell, 1987). Thus, water that 

reaches the glacier bed probably flows out of the crater through 

the rubble layer or downward into the volcano’s groundwater 

system, rather than moving along the glacier bed. In support 

of this hypothesis, we note that there are no outlet streams at 

the glacier termini, although there are springs and seeps farther 

downslope. Discharge in Loowit Creek, which heads several 

hundred meters downstream of the WCG and ECG termini and 

drains the crater, is not measured regularly, owing to the impossibility of maintaining a permanent gaging station in the very 

unstable stream channel. However, such occasional discharge 

measurements as have been made (fig. 25) show no evidence 

for systematically elevated streamflow during the eruption.

Inferences from Glacier Dynamics 

Bearing on Lava-Dome Mechanics

We envisage outward push on Crater Glacier by the 

expanding Mount St. Helens dome as involving not glacier sliding, as usually considered by glaciologists, but low-angle thrust 

faulting. In our view, the glacier is being pushed over the underlying unconsolidated rock debris, with the décollement probably near the glacier bed (glacier bed being a rather ill-defined 

concept in the present case, as discussed earlier). Our conception of the process is sketched in figure 26. Glacier deformation 

DATE

1/1/90 1/1/94 1/1/98 1/1/02 1/1/06

DISCHARGE, IN CUBIC METERS PER SECOND

0.0

0.1

0.2

0.3

0.4

0.5

Figure 25. Discharge in Loowit Creek measured above Loowit 

Falls, from unpublished streamflow data collected by hydrologic 

surveillance staff at Cascades Volcano Observatory. Probable 

error in measurements is about 10 percent.
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 271

is for all intents and purposes quasistatic—accelerations can be 

ignored—so conservation of momentum reduces to a force balance. The force exerted by the lava dome will be balanced by the 

sum of resisting forces within the ice and at the glacier bed,

nn ( ) b p WH ≈  +  , (2)

where p is the pressure (force per unit area) exerted by the lava 

dome, nn  is the deviatoric stress within the ice normal to the 

dome-glacier contact, W is the width of the glacier (that is, the 

distance from the dome to the crater wall), H is a typical value 

of ice thickness, and b  is the magnitude of the shear stress 

opposing displacement of the glacier in a direction normal 

to the dome-ice margin (fig. 26). If motion on the décollement involves essentially Coulomb friction (that is, frictional 

resistance proportional to the normal load), then b i   ≈ gH , 

where  is the coefficient of friction, i is the density of ice, 

and g is acceleration due to gravity, and our estimate for p

becomes

nn i p gW ≈   + . (3)

We have taken the normal stress on the décollement to be 

equal to the ice-overburden pressure. Thus we are supposing that water pressure on the décollement is negligible, as is 

reasonable, because, as noted above, water at the glacier bed 

apparently flows downward into the volcano rather than in a 

pressurized drainage system along the bed.

We now estimate the magnitude of the two terms on the 

right-hand side of equation 3. The deviatoric stress within the 

ice normal to the dome-glacier contact, nn  , can be estimated if 

we take into account the rheological behavior of glacier ice as 

(see appendix 3):

2/3

 nn e nn B   −=   , (4)

where nn  is the strain rate normal to the dome-glacier contact 

and the so-called effective strain rate e  (equal to one-half the 

second invariant of the strain-rate tensor) is in this case given by

2 2 2 2 2 e nn tt zz      = ++   , (5)

where tt  is the strain rate tangential to the dome-glacier 

contact and zz  is the vertical strain rate. In writing equation 

5, we assume that the directions normal and tangential to the 

dome-glacier contact are the directions of principal strain rates, 

an assumption that is supported by the available data. Using

B = 5.3×107 Pa·s1/3 (Paterson, 1994; van der Veen, 1999) and 

the strain-rate calculations given above, we estimate nn ≈

0.16–0.21 MPa. Taking  ≈ 0.5 (consistent with there being 

considerable debris within the ice and thus much rock-to-rock 

friction at the décollement), 3 i = 900 kg/m (corresponding to 

glacier ice, not snow or firn), g = 9.8 m/s2, and W ≈ 250 m, the 

frictional term on the right-hand side of equation 3 has a magnitude of about 1.1 MPa. Frictional resistance on the décollement therefore dominates the force balance, with the estimated 

nn p

b

W

H τ

τ

Figure 26. Sketch 

illustrating conception of 

the force balance involved 

in outward displacement of 

Crater Glacier by growing 

lava dome. Symbols 

defined in text.

value of p being about 1.3 MPa. This estimate is admittedly 

rough, as we have not factored in the complicated geometry of 

the real system.

Summary and Outlook for the Future

The eruption of Mount St. Helens that began in fall 2004 

has presented us with the first-ever opportunity to observe 

and document emplacement of a lava dome through glacier 

ice. The eruption has not caused any rapid melting of Crater 

Glacier, but the effects on the glacier have nonetheless been 

striking. Dome growth cut the glacier in two and then successively squeezed the two parts. Measurements using both 

specialized, helicopter-deployed GPS stations and photogrammetrically derived DEMs showed that the two glaciers 

underwent deformation of an extreme variety, with strain rates 

of extraordinary magnitude as compared to those in normal 

alpine temperate glaciers. Moreover, the GPS-derived motion 

records make clear that Crater Glacier is fundamentally unlike 

normal alpine glaciers, in that there is no evidence that it slides 

over its bed. The most reasonable explanation for this anomaly 

is that meltwater reaching the glacier bed enters the volcano’s 

groundwater system rather than flowing toward the glacier 

terminus through a drainage network along the bed.

The part of east Crater Glacier that underwent thickening has been thinning since dome growth shifted to the west 

in April 2005, and normal ice flow has moved mass downstream. Terminus advance is likely to continue unless eruptive processes remove substantial glacier mass. West Crater 

Glacier is likely to evolve similarly in the short term, with 

terminus dynamics complicated by the formation of the shear 

zone shown in figure 20. The pattern of snow accumulation 

in the crater has been radically perturbed, with heat from the 

new lava dome locally preventing accumulation. Sufficiently 

prolonged dome growth could, of course, completely eliminate 

ice from the crater (and indeed completely eliminate the crater 

itself). Glaciers at Mount St. Helens come and go, modulated 

by the style and rhythm of eruptive behavior.
272 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Acknowledgments

R.M. Iverson and S.P. Anderson made helpful comments 

on an earlier version of this manuscript. DEMs were prepared 

by J. Messerich of the U.S. Geological Survey Photogrammetric Lab, Denver, Colorado, using aerial photographs taken by 

Bergman Photographic Services.

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Rotterdam, A.A. Balkema, 462 p.

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S.P., 2005, Crater glaciers on active volcanoes; hydrological 

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doi:10.3189/172756407782282426.
274 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Interstitial Ice in CraterFloor Rock Debris

As shown by Mills (1992) and noted above, until about 

1986, material accumulating on the Mount St. Helens crater floor 

consisted primarily of rock-avalanche material with interstitial 

snow. After 1986, the volumetric rate of snow accumulation 

exceeded the accumulation rate of rock debris. By the time the 

2004 eruption began, the crater-fill material was locally as thick 

as 200 m, and it is hard to envisage that any interstitial snow 

within the lowermost fill would not have transformed to glacier 

ice (Paterson, 1994). However, there is reason to believe that 

some of the deepest fill may in fact be ice free, because interstitial ice within the rock framework will be melted by geothermal 

heat and not replaced by ice from above. If all heat flux from 

below causes melting, then the melt rate m , expressed as thickness per unit time, will be given by the ratio of the geothermal 

heat flux, Gq , to the energy required to melt a unit volume of ice,

G

i

q m

 L  = , (6)

where  is porosity of the avalanche debris, i is the density of 

ice, and L is the heat of fusion. If one considers a glacier in a 

nonvolcanic setting, then taking = 0 , L = 3.35 × 105 J/kg, and 

qG = 0.05 W/m2, one finds m ≈ 5 mm/yr. In a volcanic setting,

qG could easily be one hundred times greater (Murav’ev and 

Salamatin, 1990; Salamatin and Murav’ev, 1992), and taking 

 ≈ 0.4 for the crater-fill avalanche debris, one finds m ≈ 1 m/

yr. Clearly melting can proceed rapidly, although we stress that 

these estimates for m are upper bounds, because ground water 

could carry away some of the geothermal heat flux.

The rate at which overlying glacier ice can flow into the avalanche debris is very low. The ice intrusion rate V is proportional to the gradient of ice pressure across the debris layer 

(Iverson and Semmens, 1995):

V KP = s g , (7)

where Pg is the gradient of ice pressure across the debris 

layer, and the proportionality constant Ks is analogous to 

hydraulic conductivity for ground-water flow. The constant Ks can be determined on theoretical grounds (Philip, 

1980) in the case that the debris grain size is small enough 

that ice flow is dominantly by regelation, with plastic creep 

negligible; experimental results of Iverson and Semmens 

(1995) support Philip’s theory. The crater-fill debris is coarse 

enough that plastic creep is necessary for the ice to flow 

through the pore space (Hallet, 1979), so the regelation-only 

value KS ≈ 3 × 10−15 m2

/Pa·s will give an overestimate of V. 

The ice pressure gradient Pg obviously depends upon the 

thickness of the ice-filled debris layer and the pressure of the 

overlying ice. For present purposes, suppose that the overburden pressure is 1 MPa (corresponding to an ice thickness 

of about 110 m) and the thickness of the ice-filled debris 

layer is 10 m. We then find from equation 7 that an upperbound estimate of V is about 0.01 m/yr. A balance between

m and V can exist only if the ice-filled debris layer is very 

thin—a few centimeters at most. We conclude that the ice 

within the deepest crater fill ought to, over time, melt out and 

not be replenished. It seems likely that the deepest crater fill 

will act as an aquifer conveying water along the crater floor 

toward the glacier terminus.
13. Effects of Lava-Dome Growth on the Crater Glacier of Mount St. Helens, Washington 275

Appendix 2. Calculating Glacier Volume

The change in total glacier volume within the crater of 

Mount St. Helens, over the course of the eruption, was determined by GIS methods. Details of the method can be understood with reference to figure 27. The area covered by the 

glacier before the eruption was broken into three parts: part A1 

includes the area within which dome rock was emplaced, and 

parts A2 and A3 are the east and west glacier arms that were 

not disrupted directly by dome growth. The glacier volumes in 

A2 and A3 were determined by differencing DEMs for various dates with the October–November 1986 DEM, the latter 

representing, as we argued in the main text, approximately the 

bed of the crater glacier.

As we are only trying to track the change in glacier volume 

with time, rather than total glacier volume, the exact choice of 

A2

A3

A1

0 400 800

METERS

5117000

562000 563000

5116000

Figure 27. Separation of crater DEM coverages into three 

sections used in calculating total glacier volume as given 

in figure 21. Area A1 overlaps new lava dome, whereas 

areas A2 and A3 correspond to the downstream arms of 

west Crater Glacier and east Crater Glacier, respectively. 

Method of volume calculation is described in appendix 

2. Background is hillshade-relief map based on DEM of 

October 24, 2005. Coodinates are UTM zone 10 easting and 

northing, North America datum 1983.

datum for the bed is not critical. (The ambiguity in determining the bed, upon which we commented in the main text, is 

thus not a problem.) The glacier volume in A1 is calculated as 

follows: Using the 1986 DEM as the datum, let the difference 

between the total volume above this surface at some date  be 

denoted by V , and the volume of extruded dome rock within 

A1 be given by Vd . (The calculation of Vd . is described by 

Schilling and others, this volume, chap. 8.) The glacier volume 

within A1 is then V −Vd . This volume is added to the volumes 

in A2 and A3 to get the total glacier volume. The error in this 

total volume can be estimated as the total glacier surface area 

(1.0 km2) times the root-mean-square error in the elevationdifferencing procedure, which we take as 2.5 m. (This value 

follows from the 2.5 m error on the 1986 DEM, which was 

produced from a topographic map with contour interval of 

5 m, and the 0.1 to 0.2 m error on later DEMs, which were 

produced directly from aerial photographs.
276 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 3. Glacier Flow Dynamics

The flow law of ice is an empirical relation between 

stress and strain rate. For isotropic ice, the flow law is customarily written as the tensor relation (van der Veen, 1999)

2 ij ij   =  , (8)

where ij  are deviatoric stresses, ij  are strain rates, and  is 

an effective viscosity that depends on the overall strain-rate 

field:

(1/ ) 1 ( / 2) n  B e  − =  . (9)

B is a material property that depends on temperature, n ≈ 3 for 

glacier ice (as compared to n = 1 for a Newtonian-viscous fluid 

like water), and e  is the effective strain rate, defined by the 

relation

22 2 2 2 22 2 2( ) e xx yy zz xy xz yz       =+++ ++ . (10)

Here x, y, and z are arbitrary orthogonal coordinates. In the 

simple case of unidirectional slab flow—that is, flow driven by 

gravity and resisted by drag on the base (see, for example, van 

der Veen, 1999)—one could choose x as the downglacier coordinate, y as the cross-glacier coordinate, and z as normal to the 

glacier surface. The only nonzero strain-rate component would 

then be xz  , in which case e xz  =  and the flow law becomes 

a simple relation between shear stress and shear strain rate,

xz xz xz = B   (1/ ) 1 n −

. (11)

We argued in the main text above that the crater glacier probably does not slide over its bed. The average strain rate xz  may 

therefore be estimated simply as the surface speed U divided by 

the glacier thickness H. Taking U ≈ 1 m/d and H ≈ 150 m in the 

part of east Crater Glacier experiencing lateral squeeze (see, for 

example, station ICY5 on fig. 14), the magnitude of the average 

strain rate xz  is then about 6.7×10−3/d, comparable to the magnitude of the lateral strain rate yy  and vertical strain rate zz  , 

which, as noted in the main text, averaged about 3×10−3/d over 

the period of squeezing. The value e  is not well approximated 

by xz  in this case, and the nonzero (and in fact relatively 

large) values of yy  and zz  substantially reduce the effective 

viscosity (see equations 9 and 10).

In the slab-flow model, shear stress is simply proportional 

to depth and surface slope; that is, sin xz i  =  gz , where 

is surface slope and z increases downward from the glacier 

surface. One then finds that U depends upon flow ice thickness 

and slope according to the expression

2 sin

1

n

i H gH U n B

    =   +  

. (12)

With the usual value n = 3, the surface speed then varies as H4.
Heading 1 277

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278 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 279

Geodesy and Remote Sensing

The renewed eruption of Mount St. Helens provided many opportunities 

to bring modern geodetic tools—such as radar interferometry and GPS—to 

bear on monitoring what was, from 2004 to 2008, the only erupting volcano in the conterminous United States. The volume of lava extruded, about 

0.1 km3

, proved small enough that geodetic signals associated with ground 

deformation were subtle and often difficult to interpret.

Surprises? The days and weeks leading to the seismic unrest of September 23, 2004, lacked demonstrable ground deformation. Presumably the volcano was already pressurized, its magma having moved into a position favorable for eruption as early as the late 1980s or early 1990s and awaiting only 

some incremental change that would initiate the final ascent. The amount of 

lava extruded between 2004 and 2006 was substantially greater than could be 

accounted for by reservoir withdrawal, according to models based on the GPS 

data. Instead, the discrepancy must be accounted for by some combination of 

magma recharge from greater depth and magma expansion by vesiculation in 

the shallow magma reservoir and its conduit.

It is this question of recharge by new magma—how much, from how 

deep, and of what composition?—that remains the inestimable puzzle, 

addressed partly by chapters in this section but raised again in sections 

describing modeling and mechanics, gas geochemistry, and petrology.
280 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

View southwest across spine 3 on November 29, 2004. 

USGS photo by J.W. Vallance.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 14

Constraints and Conundrums Resulting from GroundDeformation Measurements Made During the 2004–2005 

Dome-Building Eruption of Mount St. Helens, Washington

By Daniel Dzurisin1, Michael Lisowski1, Michael P. Poland2, David R. Sherrod1, and Richard G. LaHusen1

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, PO Box 51, Hawaii National Park, HI 96718

Abstract

A prolonged period of dome growth at Mount St. Helens 

starting in September–October 2004 provides an opportunity 

to study how the volcano deforms before, during, and after an 

eruption by using modern instruments and techniques, such as 

global positioning system (GPS) receivers and interferometric 

synthetic aperture radar (InSAR), together with more traditional ones, including tiltmeters, triangulation, photogrammetry, and time-lapse photography. No precursory ground 

deformation was detected by campaign GPS measurements 

made in 2000 and 2003, nor by a continuous GPS station 

(JRO1) operating ~9 km to the north-northwest of the vent 

area since May 1997. However, JRO1 abruptly began moving downward and southward, toward a source centered about 

8 km beneath the volcano, concurrently with the start of a 

shallow earthquake swarm on September 23, 2004. The JRO1 

velocity slowed from ~0.5 millimeters per day (mm/d) in late 

September–early October 2004 until spring 2005. Thereafter, 

it was essentially constant at ~0.04 mm/d through December 

2005. In similar fashion, the growth rate of the welt on the 

south crater floor slowed from 8.9 m3

/s during October 4–11 

to 6.4 m3/s during October 11–13, 2004; this trend continued 

after emergence of the first lava spine on October 11. The 

volumetric extrusion rate decreased from 5.9 m3/s during 

October 13–November 4, 2004, to 2.5 m3/s during December 

11, 2004–January 3, 2005, and for the remainder of 2005, it 

was in the range 2.0–0.7 m3/s. Fifteen continuous GPS stations, installed soon after the eruption began, showed radially 

inward and downward ground motions through December 

2005. Likewise, InSAR observations spanning the first year of 

the eruption indicate broad subsidence centered near the vent. 

Model-derived estimates of source-volume decrease from 

September 23, 2004, to October 31, 2006, are 16–24×106

 m3, 

substantially less than the volume erupted during the same 

period (87×106 m3 through October 21, 2006). The discrepancy can be explained by a combination of magma expansion 

and recharge in the source region.

Lack of precursory deformation at JRO1 suggests that 

the conduit is poorly coupled to the rest of the edifice, so the 

rising magma column was able to push ahead older conduit material rather than intruding it. Constraints on conduit 

length and radius require that reservoir magma (as opposed 

to conduit-filling magma) reached the surface early during 

the eruption, probably soon after CO2 emission rates peaked 

in early October 2004. If rapid emergence of spine 3 (the first 

whaleback-shaped extrusion) in late October 2004 marked 

the arrival of reservoir magma, then the volume of conduit 

material flushed from the system was about 20×106 m3—the 

volume of surface deformation plus spines on November 4, 

2004. The corresponding radius for a cylinder extending from 

the surface to depth d = 5 km is 35.7 m, or 28.2 m for d = 8 

km. The average ascent rate through the conduit, assuming 

reservoir magma began its rise on September 23, 2004, was 

120 m/d for d = 5 km, or 190 m/d for d = 8 km. Observed lineal extrusion rates were 2–10 m/d, so the conduit must widen 

considerably near the surface. Equating magma flux through 

the conduit to that at the surface, we obtain a vent radius of 

125 m and an extrusion rate of 5.7 m3

/s—both values representative of the early part of the eruption.

Lack of precursory inflation suggests that the volcano was poised to erupt magma already stored in a crustal 

reservoir when JRO1 was installed in 1997. Trilateration 

and campaign GPS data indicate surface dilatation, presumably caused by reservoir expansion between 1982 and 1991, 

but no measurable deformation between 1991 and 2003. We 

conclude that all three of the traditionally reliable eruption 
282 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

precursors (seismicity, ground deformation, and volcanic 

gas emission) failed to provide warning that an eruption was 

imminent until a few days before a visible welt appeared at 

the surface—a situation reminiscent of the 1980 north-flank 

bulge at Mount St. Helens.

Introduction

The deformational behavior of active basaltic shield 

volcanoes and large silicic magmatic systems (restless calderas) can be characterized in a general way by, respectively, 

(1) repeated inflation-deflation cycles in response to changes 

in crustal magma storage (for example, Kïlauea, Hawai‘i, 

U.S.A., Dvorak and Okamura, 1987; Krafla, Iceland, Tryggvason, 1994), and (2) episodes of more gradual surface uplift 

or subsidence (bradyseisms) caused by magmatic inflationdeflation or hydrothermal system pressurization-depressurization (for example, Long Valley Caldera, Calif., U.S.A., Hill 

and others, 1985; Yellowstone caldera3

, Wyo., U.S.A., Dzurisin and others, 1999; Wicks and others, 2006; Phlegraean 

Fields caldera, Italy, Caputo, 1979). Deformation of stratovolcanoes of intermediate composition (andesite–dacite) is more 

varied and therefore more difficult to understand. Nonetheless, 

volcanoes of this type are numerous along the Pacific margin 

and, on short to intermediate time scales, they pose a greater 

threat to human populations than do basaltic shields or large 

silicic systems.

The reawakening of Mount St. Helens in 1980 following 

123 years of quiescence provided a modern opportunity to 

study pre-, syn-, and post-eruptive deformation at an accessible and relatively well-studied stratovolcano. However, little 

was learned about any deep-seated deformation (source depth 

more than ~2 km) that might have occurred before the eruption because magma rose to shallow depth—as evidenced by 

the appearance of a bulge on the volcano’s north flank by early 

April 1980 (fig. 1A)—before or soon after March 20, 1980, 

when a swarm of earthquakes caught volcanologists’ attention and spurred the start of intensive monitoring (Lipman and 

others, 1981).

Renewed dome growth at Mount St. Helens starting in 

September–October 2004 (fig. 1B) provides a second opportunity to study how this volcano deforms before, during, and 

after an eruption—this time by using modern geodetic instruments and techniques, such as Global Positioning System 

(GPS) receivers (Lisowski and others, this volume, chap. 15; 

LaHusen and others, this volume, chap. 16) and interferometric synthetic aperture radar (InSAR, Poland and Lu, this 

volume, chap. 18), together with more traditional instruments 

3

 Capitalization of “Caldera” or “Volcano” indicates adoption of the word as 

part of the formal geographic name by the host country, as listed in the Geographic Names Information System, a database maintained by the U.S. Board 

on Geographic Names. Noncapitalized “caldera” or “volcano” is applied 

informally—eds.

and techniques, including tiltmeters, triangulation, photogrammetry, and time-lapse photography (Poland and others, 

this volume, chap. 11; Major and others, this volume, chap. 

12). Deformation data from the first 15 months of the eruption (October 2004–December 2005) help to constrain (1) the 

depth and geometry of the magma plumbing system and (2) 

the time history of magma outflow and possible recharge to 

a crustal magma reservoir (Lisowski and others, this volume, 

chap. 15; Mastin and others, this volume, chap. 22). The 

2004–5 deformation data, especially when considered together 

with other datasets, also raise several new questions.

1980–1986 dome

Oct 1, 2004, ashfall

Oct 1, 2004, vent

Sep−Oct 2004 welt

Crater

Bulge

A

B

Figure 1. Examples of precursory surface deformation before 

eruptions at Mount St. Helens, Washington. A, Upper north flank 

on April 7, 1980. Series of phreatic eruptions that began on March 

27, 1980, deposited ash and formed summit crater, which had 

grown to about 500 m west-east and 300 m north-south. Intensely 

cracked area below and north of crater shows distension of north 

flank. USGS photograph by R.B. Waitt. B, Crater floor as seen from 

above southwest crater rim on October 3, 2004. Vent for October 

1 ash is visible at the west base of the welt of uplifted crater floor 

and glacial ice first recognized as deforming on September 27, 

2004. USGS photograph by S. Konfal.
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 283

This paper draws freely on datasets and analyses presented elsewhere in this volume. Its main objectives are (1) to 

present an integrated overview of ground deformation observations, measurements, and results through December 2005 and 

(2) to explore some implications of these results, together with 

those from other disciplines, for the mechanism of the eruption 

and its future course.

Geodetic Techniques Used to Study the 

Eruption

This section summarizes the techniques used, successfully or unsuccessfully, to measure ground deformation during 

the 2004–5 eruption. The reader is referred to other papers 

in this volume for more thorough treatments of some of the 

techniques, results, and interpretations.

Regional Trilateration and Campaign-Style GPS 

Measurements

The Helens high-precision trilateration network was 

established in 1982 to track regional strain accumulation following the 1980 eruption at Mount St. Helens. Line lengths 

were measured with a Geodolite in 1982 and 1991 and with 

GPS in 2000. Lisowski and others (this volume, chap. 15) 

report that line lengths consistently increased by 1–3 cm 

between 1982 and 1991 but did not change significantly from 

1991 to 2000. The line-length increases correspond to areal 

dilatation at a rate of 144±39 nanostrain/yr, which is an order 

of magnitude greater and distinctly different than the background tectonic strain. Lisowski and others (this volume, chap. 

15) discount possible systematic error in the Geodolite surveys 

as the sole reason for the observed dilatation and invoke “...

some other phenomenon, like recharge of the magma reservoir...” to account for the observations.

During summer 2000, the USGS Cascades Volcano 

Observatory (CVO) installed and made initial GPS observations at more than 40 benchmarks on and around Mount St. 

Helens (herein referred to as the large-aperture GPS network). 

The network is concentrated within 10 km of the volcano, but 

it extends more than 30 km and covers an area of more than 

7,400 km2

 (fig. 2; Dzurisin, 2003). The first repeat survey was 

made in summer 2003. With effects of rigid-block rotation and 

regional strain removed, no significant strain was detected, 

with one exception: a station (DMSH) on the 1980–86 lava 

dome (fig. 2) moved down and east-northeastward, toward the 

center of the dome, at average rates of 9.0 cm/yr and 2.9 cm/

yr, respectively (Lisowski and others, this volume, chap. 15).

The motion of DMSH is consistent with that seen at 

several nearby stations on the 1980–86 dome, which were 

observed each summer from 2000 to 2004 to provide geodetic 

control for vertical air photos and digital elevation models 

(Schilling and others, this volume, chap. 8). These stations 

generally moved down and toward the center of the dome at 

rates of a few centimeters per year, while stations on the crater 

floor moved very little (Lisowski and others, this volume, 

chap. 15). The motion of the stations is attributed to cooling 

and compaction of the uppermost conduit and dome. The latter 

consists mainly of stubby, highly deformed, dacite lava flows 

(lobes) and bouldery talus (Swanson and others, 1987). There 

is no evidence from the August 2004 control-point survey data 

of any change in trend that could be associated with the start 

of the current eruption.

The regional trilateration and campaign GPS data indicate that most of the surface deformation prior to the 2004–5 

eruption took place in the decade following the 1980 eruption, 

and that little or no surface deformation occurred between 

1991 and 2003. This is consistent with data from continuous GPS (CGPS) station JRO1, 8 km north-northwest of the 

volcano, which recorded no volcano-centric deformation from 

May 1997 to the beginning of seismic unrest on September 23, 

2004 (see below). However, there is seismic evidence for pressurization of the magma reservoir and upward fluid intrusion 

starting in late 1987 and continuing at least until November 

2001 (Moran, 1994; Moran and others, this volume, chap. 2). 

This apparent contradiction and other questions are addressed 

in the later Discussion section.

Continuous GPS (CGPS)

Johnston Ridge Observatory (JRO1)

Included in the large-aperture GPS network are 17 CGPS 

stations on or close to the volcano. The first station, JRO1, 

was installed in May 1997 at the Mount St. Helens National 

Volcanic Monument’s Johnston Ridge Observatory. Lisowski 

and others (this volume, chap. 15) conclude that, within uncertainties, (1) the velocity of JRO1 was constant from May 1997 

through mid-September 2004, and (2) JRO1 abruptly started 

moving south-southeast and down, toward a source beneath 

Mount St. Helens, on or about September 23, 2004. The 

anomalous motion was essentially concurrent (±1 day) with 

the start of a shallow earthquake swarm beneath the crater 

floor (Moran and others, this volume, chap. 2). Seismicity was 

accompanied by the growth of a large welt (fig. 1B), which 

was first visible in photographs taken on September 26. The 

first of several lava spines emerged from the welt on October 

11, and extrusion of a dacite dome continued through the end 

of 2005 (Dzurisin and others, 2005; Vallance and others, this 

volume, chap. 9).

Fresh crevasses in glacier ice on the south crater floor, 

indicating intense deformation of the area that was to become 

the welt, were first recognized a few days after the start of 

shallow seismicity and anomalous motion at JRO1. New 

crevassing and uplift were confirmed by direct observation 

on September 29, 2004. In hindsight, they were recognized in 

photographs taken on September 26 by J.S. Pallister during 
284 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

PACK

MORT

VANS

APSA

4100

WOOD

ARIL

YMSH

REDD

MORG

JACK SMCR

BGRM

PAN1 VMSH

HPAD

LVCY

NMAR

BIVO

MITC

JRO1

ELKK

TUBE

WGOT

5

5

12

Cowlitz River

N.Fork Toutle River

S.Fork Toutle River

Lewis River

0 10 20 30 40 50 KILOMETERS

WASHINGTON

123°00' 122°30' 122°00'

46°30'

46°15'

46°00'

A

R100

LOOW

STUD

NWDO

SOFO

WRAD

DMSH NELS

EADM

APEC

WIWE CAFE

WRID

SRID NEBU

SESL

SERI

B MOUNT ST. HELENS C

AREA OF

MAP

30

Figure 2. Large-aperture GPS network at Mount St. Helens, Washington. A, Location map 

showing 40 benchmarks observed during summers of 2000 and 2003. Base from shaded relief 

version of USGS digital elevation model. B, Benchmarks on or near the volcano. C, GPS receiver 

deployed at NELS (Nelson Ridge) on upper east flank of Mount St. Helens in October 2004.

a fixed-wing observation flight. Helicopter pilot Jeff Linscott (JL Aviation, Inc.) first noticed new cracks in the ice on 

September 28 while flying with members of the news media. 

He pointed out the cracks to us (D. Dzurisin and M.P. Poland) 

the next day, September 29. A photograph taken by a hiker 

from the south crater rim on September 25 shows no obvious 

disturbance in the area that became the welt; thus, the first 

surface cracks probably formed 2–3 days after the start of a 

shallow earthquake swarm at about 0200 PDT (0900 UTC) on 

September 23 and within 1 day of an uptick in seismicity on 

September 25 that culminated with the onset of the eruption. 

On August 27, 2004, Linscott landed in the southern part of 

the crater with USGS–CVO staff; fresh cracks in glacier ice 

at their customary landing site forced them to use an alternate 

site nearby. Any significance of those cracks in terms of the 

impending eruption is undocumented. At the time, the cracks 

were attributed to summer snowmelt revealing crevasses in 

the advancing glacier. Similar features had been seen before 

and were visible at the time on the west arm of the glacier. 

Any relation to the recognized onset of volcanic unrest nearly 

a month later is speculative and not supported by the hiker’s 

photograph taken on September 25.
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 285

Following a period of rapid motion toward a source 

beneath the volcano at a peak rate of ~0.5 mm/d in late September–early October 2004 (while the growth rate of the welt 

was as high as 8.9 m3/s), the movement of JRO1 gradually 

slowed through December 2004. Meanwhile, the extrusion rate 

slowed from 5.9 m3/s (October 13–November 4) to 2.5 m3/s 

(December 11, 2004–January 3, 2005; Schilling and others, 

this volume, chap. 8). During May–December 2005, the velocity of JRO1 was essentially constant at ~0.04 mm/d toward the 

volcano. During the same interval, the rate of lava extrusion 

declined from 1.3 m3/s (April 19–June 15) to 0.7 m3/s (October 24–December 15), with somewhat higher rates during July 

14–August 10 (2.0 m3/s) and August 10–September 20 (1.6 

m3/s; Schilling and others, this volume, chap. 8).

Mount St. Helens CGPS Cluster

In response to the current eruption, CVO and the EarthScope Plate Boundary Observatory (PBO) installed 13 additional CGPS stations during October–November 2004. Most 

of the new stations are on the volcano, mounted 3–4 m high 

on steel tripods that served as EDM stations during the 1980s 

(fig. 3). Data from all 14 CGPS stations at Mount St. Helens 

are downloaded automatically to CVO and to the UNAVCO 

Boulder Facility Archive. The data are processed daily at the 

USGS Menlo Park facility to produce a 24-hour solution. Five 

additional stations, mounted on smaller tripods and without 

telemetry links, have operated intermittently since November 

2004. Data from those stations are downloaded manually.

Radio whip antenna

Solar

panels

Spirit Lake

Batteries

and

GPS receiver

GPS antenna

Figure 3. CGPS station WIWE (Windy West) on lower northeast 

flank of Mount St. Helens, Washington (fig. 2). Station is typical 

of several installed by CVO. View is to north-northeast. Mount 

Rainier is visible 77 km in distance.

Time-series plots of CGPS data show that the rate of 

motion toward the volcano peaked in late September or early 

October 2004, while the welt was growing rapidly but before 

the first lava emerged on October 11, 2004 (see figs. 10–13 in 

Lisowski and others, this volume, chap. 15). Both the far-field 

(beyond the crater floor) deformation rate and the extrusion 

rate declined progressively from mid-October through at least 

December 2004. Perhaps as early as January 2005, and surely 

by May 2005, the deformation rate and extrusion rate stabilized; 

both rates were relatively steady through December 2005.

Single-Frequency (L1-Only) GPS “Spiders”

It was clear from the early days of the eruption that 

extreme deformation of the south crater floor and glacier must 

attenuate rapidly with distance, because the 1980–86 dome 

and crater floor to the north remained essentially undisturbed, 

while the welt, just a few hundred meters to the south, grew 

to 500 m west-east, 600 m north-south, and 110–120 m high 

by October 11, 2004. LaHusen and others (this volume, chap. 

16) report that a single-frequency CGPS station at SEP, on 

the 1980–86 dome less than 500 m north of the 2004–5 vent, 

moved anomalously about 20 cm north, 8 cm west, and 12 cm 

up during the 9-month period ending on September 27, 2004. 

This displacement was much less than we expected, given 

the large size and proximity of the welt. Likewise, all 13 of 

the CGPS stations that were installed on the volcano’s flanks 

during October–November 2004 moved less than 2 cm during 

the ensuing 15 months. Therefore, it was necessary to install 

deformation sensors within a few hundred meters of the vent 

to monitor the course of the eruption. Conventional CGPS 

installations were not a reasonable option, because instrumentation in proximal areas had an expected lifetime of only days 

or weeks during the first few months of the eruption. Instead, 

we deployed and, in some cases, retrieved and redeployed a 

series of single-frequency (L1-only) CGPS stations aboard 

“spiders”—self-contained sensor platforms slung from a helicopter to otherwise inaccessible or hazardous sites (LaHusen 

and others, this volume, chap. 16). Here we summarize results 

of three GPS spider deployments to the growing dome, which 

provided reliable measurements of lineal extrusion rates for 

spines 3 and 4. See LaHusen and others (this volume, chap. 

16) for discussion of additional deployments to the 2004–5 

dome, 1980–86 dome, and Crater Glacier.

On November 20, 2004, GPS spider ELEA was set down 

astride spine 3, which was advancing south-southeastward 

across the crater floor. During the ensuing six days, before 

the instrument was destroyed in a rockfall, ELEA traveled 

67 m southeast and 8 m up at an average rate of ~10 m/d and 

an average slope of 5.6° (LaHusen and others, this volume, 

chap. 16). The GPS spider CDAN was destroyed by a small 

explosive event starting at 0318 PST on January 16, 2005 

(Scott and others, this volume, chap. 1), about 36 hours after 

it was deployed to spine 4. In the preceding 24 hours, CDAN 

moved 6.5 m south, 2.2 m east, and 4.7 m up, corresponding 
286 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

to an average velocity of 8.3 m/d at an average slope of 34.5°. 

The GPS spider AHAD was deployed to spine 4 on February 

8, 2005. By the time it was retrieved on February 16, AHAD 

had moved 29 m south, 17 m east, and 9 m up, for an average 

velocity of 6.2 m/day at an average slope of 15.5° (LaHusen 

and others, this volume, chap. 16).

Velocities measured by ELEA, CDAN, and AHAD are 

not directly comparable for two reasons: (1) spines 3 and 4 

differed in morphology (4 was, for the most part, steeper than 

3), and (2) the three spiders occupied different positions with 

respect to the vent. Nonetheless, slowing of the extrusion 

velocity between late November 2004 and mid-February 2005, 

as indicated by the GPS spiders, is consistent with analysis of 

time-lapse images from the Sugar Bowl camera (Major and 

others, this volume, chap. 12) and with declining volumetric extrusion rates derived from analysis of digital elevation 

models (DEMs) for the same period (Schilling and others, this 

volume, chap. 8).

No additional GPS spiders were deployed to active parts 

of the growing dome from March to December 2005 because 

the surface was too steep, too rugged, or too hot. Meanwhile, 

spiders deployed to older parts of the 2004–5 dome, the 

1980–86 dome, and Crater Glacier continued to provide data 

(LaHusen and others, this volume, chap. 16).

Interferometric Synthetic Aperture Radar (InSAR)

InSAR has been used successfully to study surface 

deformation at dozens of volcanoes worldwide (for example, 

Amelung and others, 2000; Zebker and others, 2000; Lu and 

others, 2007), but at Mount St. Helens three factors have limited its utility. First, EDM and GPS measurements show that 

any surface displacements that occurred at Mount St. Helens 

between InSAR’s emergence in 1993 (Massonnet and others, 

1993) and the start of the 2004–5 eruption were small compared 

to the resolution of InSAR (2.83 cm/fringe for C-band images; 

Lisowski and others, this volume, chap. 15). Second, persistent 

winter snow pack and frequent precipitation destroy coherence, 

except in small areas above the tree line in summer-to-summer 

interferograms. Third, the orbital repeat intervals for radar satellites ERS (35 days), RADARSAT (24 days), and ENVISAT (35 

days) are too short to capture any slow, preeruption deformation 

and too long to capture rapid near-vent deformation while an 

eruption is underway. In the latter case, extreme deformation of 

the growing dome and displaced glacier make the in-crater part 

of all interferograms incoherent.

These difficulties with InSAR can be overcome to some 

extent by “stacking” (averaging) several interferograms for 

relatively short time periods (so coherence is maintained), 

which collectively span a period long enough for any deformation signal to emerge from the noise. Stacking works on the 

premise that a small deformation signal from a fixed source 

will accumulate with time, while random atmospheric noise 

sources will tend to cancel. Poland and Lu (this volume, chap. 

18) stacked and analyzed hundreds of interferograms of Mount 

St. Helens acquired by ERS-1/2, ENVISAT, and RADARSAT 

satellites from 1992 to 2005. Stacks of preeruptive interferograms indicate no sign of volcanowide deformation, which 

is consistent with trilateration, campaign GPS, and CGPS 

data for the period from 1992 to 2005 (Lisowski and others, 

this volume, chap. 15). In contrast, the most reliable stack of 

coeruptive interferograms shows line-of-sight (LOS) increases 

(that is, surface subsidence) on the volcano’s flanks at rates of 

40–50 mm/yr, which is consistent with CGPS results for the 

coeruptive period.

Time-Lapse Oblique Photography

Several factors contributed to the widespread use of photography, both from aircraft and the ground, to record various 

aspects of the eruption including (1) growth, deformation, and 

geomorphic evolution of the lava dome, (2) perturbed motion 

and deformation of Crater Glacier, and (3) emergence of the 

lava column from the vent and resulting movement of spines 

across the crater floor. First, digital cameras are capable of 

producing high-quality images in near-real time (an advantage over film cameras) that can be shared among colleagues 

with ease. Second, the perceived level of risk to observers 

in the crater, especially during early phases of the eruption, 

was a factor that encouraged remote photo documentation. In 

hindsight, the 2004–5 eruption has been benign by comparison 

to most dacite dome eruptions; nonetheless, risks in proximal 

areas from rockfalls and small explosions were reduced by 

replacing human observers with cameras, especially for investigations that required repeated observations over long periods 

of time (for example, time-lapse images of dome growth). 

Third, it was recognized early in the 2004–5 eruption that 

cameras were providing a unique and valuable record of the 

measured pace, longevity, and character of events that transpired in the vent area. An unusual suite of circumstances—

gas-poor, crystal-rich magma emerging from an inclined vent 

on a relatively flat crater floor beneath a newly formed glacier 

that is contained on three sides by steep crater walls—combined to create an unusual and visually captivating landscape. 

Over time scales ranging from minutes to months (from 

drumbeat earthquakes to gouge-covered, laterally mobile, 

whaleback-shaped spines), the eruption produced a variety of 

features that are suited to photo documentation and analysis. 

The following sections summarize two different approaches to 

time-lapse photo studies of the eruption. More detailed discussions of these and other camera applications can be found in 

Poland and others (this volume, chap. 11) and Major and others (this volume, chap. 12).

Synoptic Observations of Dome Growth and 

Glacier Deformation

Mount St. Helens lies 70 km north-northeast of CVO, 

within sight on a clear day but at a range suitable only for 

reconnaissance observations. The volcano’s 1980 crater opens 
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 287

to the north, so the view of the current eruption site from CVO 

is blocked by the volcano’s south flank. Continuous surveillance of the eruption by observers within sight of the vent was 

impractical for logistical reasons, so remote cameras were 

used instead to track the changing landscape in the crater. 

Time-lapse photography cannot be construed as deformation 

monitoring in a classical geodetic sense, but in the 2004–5 

eruption, the disruption of the crater floor and glacier was so 

extreme that cameras became an effective tool for monitoring 

dramatic changes in the vent area. Elsewhere, CGPS stations 

and tiltmeters are better suited to measuring strains that are 

smaller by several orders of magnitude.

The first remote camera station was installed on October 10, 2004, at Sugar Bowl (altitude 1,859 m) on the east 

side of The Breach (open north end of 1980 crater), 2.3 km 

north-northeast of the vent. The Sugar Bowl camera recorded 

the changing morphology of the welt starting one day before 

emergence of spine 1 on October 11, 2004, followed by 

development, movement, and disintegration of three prominent 

whaleback spines (3, 4, and 5). The Sugar Bowl camera was 

poorly positioned for photographing the disintegration of spine 

5 and westward migration of spine 6, which began in early 

August 2005, because the camera’s view was blocked by the 

1980–86 dome from August through early October 2005. But 

by mid-October 2005, spine 7 had grown into the field of view 

of the Sugar Bowl camera above the profile of the 1980–86 

dome, where it remained through the end of 2005. Because the 

Sugar Bowl camera captured a large part of the dome growth 

that occurred from mid-October 2004 through December 

2005, its images are well suited to systematic quantitative 

analysis. Major and others (this volume, chap. 12) used the 

Sugar Bowl images to estimate lineal extrusion rates as a function of time and to infer short-term (days to weeks) variations 

in the volumetric extrusion rate.

To complement the Sugar Bowl images, two additional 

camera stations were installed higher on the crater rim in 

August 2005. Brutus (BRUT) is located at 2,479 m altitude 

along the northeast rim, 1.1 km east-northeast of the vent. 

South Rim (SRIM) is located at 2,512 m altitude along the 

south-southwest crater rim, 0.7 km from the vent. BRUT sees 

most of the 2004–5 dome and Crater Glacier from the eastnortheast; SRIM sees essentially the same area from a reciprocal vantage point.

Camera station Guacamole (GUAC), located at 1,634 m 

altitude on the floor of The Breach, 2.6 km north of the vent, 

was added in November 2005. Station GUAC provides a view 

of the western part of the 2004–5 dome, including spines 

6 and 7, and of the deformed east and west arms of Crater 

Glacier. The lower altitude of GUAC, relative to the other 

cameras, allows for occasional glimpses under the clouds 

that obscure views from the higher cameras. See Poland and 

others (this volume, chap. 11) for additional description of 

the camera stations and telemetry system, including map 

locations, equipment used, method of storing and retrieving 

images, and capability for controlling camera parameters and 

shooting schedules.

Arguably the greatest value of the remote camera stations lies in the production of time-lapse “movies” (typically one image per day for several days, or all available 

three-minute images for a 24-hour period) that portray the 

changing morphology of the dome, deformation of Crater 

Glacier, rockfall sequences, and patterns of nighttime glow 

in an intuitive and easy-to-understand format. Several of 

these sequences are available on the DVD accompanying 

this volume and on the CVO Web site (http://vulcan.wr.usgs.

gov/Volcanoes/MSH/Images/MSH04/repeat_views.html, last 

accessed January 14, 2008).

High-Resolution, Small Field-of-View 

Observations of Extrusion Site

In an attempt to capture any short-term (minutes to 

hours) variations in the lineal extrusion rate and to assess any 

correlation between these and “drumbeat” earthquakes, we 

deployed digital cameras with long focal-length lenses and 

time-lapse controllers at BRUT and on the September 1984 

lobe (SEP) of the 1980–86 dome. Drumbeat earthquakes are 

repetitive, small earthquakes, many with waveforms similar 

to other drumbeats, that occurred at remarkably constant 

intervals during the 2004–5 eruption (Moran and others, this 

volume, chap. 2; Thelen and others, this volume, chap. 4). A 

correlation between drumbeats and lineal extrusion rate is suggested by a mathematical model that describes the extrusion 

process in terms of stick-slip motion of a rigid plug under the 

combined influences of magma pressure, gravity, and friction on the conduit wall (Iverson and others, 2006; Iverson, 

this volume, chap. 21). To avoid the need for longer-term 

power and protection from the elements, each of these camera 

deployments lasted for only a few hours. From BRUT, a Nikon 

® CoolPix 990 camera coupled to a 12–36X spotting scope 

(fig. 4) produces a field of view at the vent (range ~1100 m) 

of approximately 60×90 m at 12X, or 20×30 m at 36X. From 

SEP (range ~400 m), the same camera coupled to a Questar 

® 700 telephoto lens (3.5-inch diameter, f7.8) fitted with a 

24-mm eyepiece (fig. 5) produces a 1.3×1.0 m field of view. 

Image size of the Nikon ® CoolPix 990 in “fine” resolution 

mode is 2048×1536 pixels (3.3 megapixels), so using it with 

the Questar ® 700 from SEP produces an image in which each 

pixel corresponds to ~0.6 mm on the ground.

Although the Nikon ® CoolPix 990–Questar ® 700 

combination deployed at SEP is theoretically capable of 

millimeter-scale resolution at the vent, two factors combine to 

degrade the image quality in practice. First is the difficulty in 

establishing a sufficiently stable base for an imaging system 

with such high magnification—a problem with two parts. The 

first, coupling the camera system securely to the ground, can 

be overcome by mounting the camera and lens to a rigid plate 

or a short, heavy-duty tripod and by using sandbags to fix the 

plate or tripod firmly to the ground. The second part of the 

problem is ground motion associated with incessant drumbeat earthquakes and occasional larger shocks (Mmax ~3.6). 
288 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Nikon

CoolPix 990

To DigiSnap

controller

To external

battery

Spotting

scope

2004−2005

lava dome

®

Nikon CoolPix 990 ® 

Questar 700 lens ® 

Figure 4. Temporary time-lapse photo station at Brutus (BRUT) 

on northeast crater rim, Mount St. Helens, Washington. Permanent 

camera station BRUT is ~100 m north (to the right) of this site.

Figure 5. Temporary time-lapse photo station at SEP on 

September 1984 lobe of 1980–86 dome, ~400 m north of the 2004–5 

vent, Mount St. Helens, Washington. Sandbags at base of tripod 

are for added stability. Spirit Lake near center of photo. Mount 

Rainier on skyline.

On five occasions between September 2 and November 18, 

2005, we tried frame intervals of 10 s or 30 s for periods of 

1–6 hours at times when drumbeats were occurring every 2–3 

minutes. Ground shaking is obvious in only a small fraction of 

the images, as would be expected from the framing interval, 

shutter speed (0.004–0.5 s), and average time between quakes. 

Each set of images reveals what appears to be relatively steady 

motion of the extrusion out of the vent, although any slight 

displacements of the camera system or underlying ground 

cannot be distinguished from apparent motion of the extrusion. 

We tried repeatedly to include in the images some part of the 

dome or crater wall that was relatively stable, but the camera’s 

limited field of view and depth of field made this difficult and 

ultimately unproductive.

A second factor that contributes to image degradation is 

heat shimmer caused by uneven refraction of light in unstable 

air along the camera’s line of sight. Heat shimmer is exacerbated by numerous localized heat sources on the 1980–86 and 

2004–5 domes and by extensive, hot talus deposits surrounding 

the active vent. The resulting image degradation can be mitigated to some extent by choosing the line of sight carefully, but 

this usually comes at the expense of the most desirable framing.

We have been unsuccessful in attempts to capture photographically any jerkiness in the extrusion process that might be 

associated with drumbeat or larger earthquakes. However, the 

high-resolution time-lapse images have been useful for another 

purpose. Because the images are at a known scale (determined 

empirically by photographing objects of known size with the 

same camera/lens combination at distances measured using 

a laser rangefinder), each set of images can be used to measure the average lineal extrusion velocity in the image plane. 

Motion of the extrusion can be discerned in images acquired 

only a few minutes apart.

The surface velocity of spine 6 was measured using 

high-resolution, time-lapse photography from SEP on three 

occasions from September 2, 2005, to October 18, 2005 (table 

1, fig. 6). SEP was due north of the advancing front of spine 6 

at the time, so the motion of features through the image frame 

can be used to measure west and up components of the velocity 

vector. The resultant velocity (west+up) is a reasonably good 

indicator of the lineal extrusion rate of spine 6, which was also 

advancing northward during this period owing to endogenous 

growth and spreading. The lineal extrusion rate determined in 

this way declined from 4.51 m/d on September 2 to 4.45 m/d 

on September 20 and to 2.51 m/d on October 18. Distances 

from the SEP camera station to the imaged part of spine 6 were 

measured with a laser rangefinder and are reported as the north 

component of velocity in table 1. The value for September 2 

(>3.3 m/d) is a minimum value because there were no previous rangefinder measurements. The rate was greater than 3.3 

m/d, the average value calculated from measurements made on 

September 2 and September 20, because the extrusion rate and 

growth rate of spine 6 generally slowed with time.

Figure 6 is based on a subset of 702 images acquired 

at SEP from 11:14:17 to 17:04:47 PDT (30-s interval) on 

September 20, 2005. Distance from the camera to the steep, 

gouge-covered north face of emerging spine 6, measured with 

a laser rangefinder, was 342±1 m. The camera’s field of view 

included an aluminum target, 46×61 cm in size, for scale. The 

target was hung by steel cables from an anchor slung by helicopter to the top of the spine. For 58 of the sharpest images, the 
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 289

WEST (mm) = 162 mm/hr x time (hr) - 1,489 mm

Average rate = 3.89 m/d WEST

UP (mm) = 91 mm/hr x time (hr) - 1,969 mm

Average rate = 2.17 m/d UP

-1,000

-500

0

500

1,000

1,500

12 13 14 15 16 17

PACIFIC DAYLIGHT TIME, IN HOURS

-150

-100

-50

0

50

100

150

WEST (mm) UP (mm)

WEST residual (mm) UP residual (mm)

M 2.9 M 1.0

Residuals

(mm)

M 2.9 M 1.0

WEST OR UP COORDINATE, IN MILLIMETERS

RESIDUAL FROM BEST FIT LINE, IN MILLIMETERS

Figure 6. Lineal extrusion velocity of spine 6 for ~3.5-hour period on September 20, 2005, based on analysis of 

high magnification, small field-of-view, time-lapse (30 s interval) digital images from temporary camera station SEP, 

Mount St. Helens, Washington. Departures from best-fit lines for west and up components of motion are shown 

with corresponding symbols in middle (“Residuals”).

pixel coordinates of a distinctive feature on the target relative 

to a corner of the image were determined by inspection using 

a PC graphics program. A spreadsheet program was used to 

convert pixel coordinates to millimeters at the target, using the 

known image scale, to determine optimal linear fits to the data, 

to calculate departures from the best-fit line, and to plot the 

results as a function of time. Average velocity components (that 

is, slopes of the best-fit lines) are 3.89 m/d west and 2.17 m/d 

up. Departures from the best-fit lines are irregular and small, 

on the order of ±10 mm (fig. 6, middle). Therefore, the lineal 

extrusion rate during a 3.5-hr period on September 20, 2005, 

was constant within the uncertainty of the measurements. There 

is no evidence of stick-slip motion of the spine greater than 

about 10 mm associated with drumbeat earthquakes, which 

were occurring every 2–3 minutes, nor with two larger shallow 

quakes at 15:47:19 (Md 2.9) and 16:05:26 (Md 1.0) PDT.

Average velocities (west and up) determined from similar 

observations on September 2, 2005, and October 18, 2005, are 

4.51 m/d and 2.51 m/d, respectively (table 1). No comparable 

measurements were possible from the Sugar Bowl camera on 

September 2 or September 20, but measurements on October 

18, 22, and 24 are in the range 3–4 m/d. In all three cases in 

which these types of observations were made, the motion of 

the extruding spine was steady over time scales of a few hours; 

no evidence of short-term rate variations was seen in any of 

the photos. However, such variations likely are occurring in 

the conduit during earthquakes, and variations could also be 

occurring at the surface in millimeter-scale jerks too small to be 

resolved by this technique. During the first year of the eruption, 

drumbeat earthquakes occurred at rates of 0.3–3 per minute, 

while the lineal extrusion velocity was 1–10 m/d. If all of the 

motion were associated with drumbeat earthquakes, this would 

imply an average event size of 2.3 mm/drumbeat at the source. 

This value for average event size would be smaller if, as seems 

likely, extrusion is partly aseismic and seismic dislocations are 

attenuated between the earthquake source area and the surface.
290 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Date

(2005)

West velocity

(m/d)

Up velocity

(m/d)

Extrusion 

velocity

(west + up)

(m/d)

North velocity

(m/d)

Distance to 

target (m)

September 2 4.23 1.55 4.51 >3.3 402

September 20 3.89 2.17 4.45 3.3 342

October 18 2.41 0.69 2.51 1.3 306

Table 1. Velocities determined from high-magnification, small field-of-view, time-lapse images at station 

SEP, Mount St. Helens, Washington.

Trilateration and Triangulation of Points on the 

Growing Dome

Classical trilateration and triangulation techniques were 

adapted with good results to measure changing displacement 

rates on the growing dome during 1980–86 which, together 

with changing patterns of seismicity, were used to predict the 

start of each dome-building episode from a few days to three 

weeks in advance (Swanson and others, 1983). Measurements 

were made by affixing prism reflectors to the dome and tracking their motion by using an EDM and theodolite from fixed 

points on the crater floor or, in some cases, from points on 

stable parts the dome itself. To increase the number of targets 

without continuously exposing a large number of reflectors to 

harsh conditions in the crater, some target points were marked 

with a survey pin so that a person on foot could accurately 

position a reflector at the point for a few minutes while a colleague made EDM and theodolite measurements. In this way, 

more than a dozen survey points distributed across the dome 

could be surveyed in a few hours.

We tried a similar approach to measure displacement 

rates on the 2004–5 dome, with lesser success. One difference from the 1980–86 dome is that continuous extrusion, 

frequent rockfalls, and the deeply crevassed glacier generally 

precluded foot travel on or around the 2004–5 dome. Instead, 

five targets, each consisting of a single prism reflector at the 

center of a painted wooden board attached to a barbell weight 

as a base, were slung from a helicopter to the 2004–5 dome 

and Crater Glacier. Three of the targets were set on spine 4, 

one to the east on rubble of spine 3, and one on the east arm 

of the glacier. None of the targets could be placed on active 

spine 5 because it was too steep and hot at the time. All of the 

targets were observed from BRUT (range ~1,100 m) using a 

Wild ® TC1000 total station tacheometer on April 21, May 

3, and September 20, 2005. The tacheometer can produce 

surveys with angle measurements accurate to 3 arc seconds at 

range 1–2.5 km, depending on viewing conditions. Standard 

deviation of range measurements is 3 mm + 2 ppm, which corresponds to ±5 mm at 1 km range.

Four of the targets survived to the second survey, and 

three survived to the third survey. However, reflections were 

obtained from only two targets in the initial survey and from 

no targets in the second or third surveys, so we lack range 

measurements for the repeat surveys. The lack of reflections 

likely resulted from dust clinging to the prisms, etching of the 

glass by volcanic gases, and (or) changing alignment of the 

prisms with respect to the sight line as a result of earthquake 

shaking or slumping. The longest-surviving targets were near 

the top of spine 4, near the base of spine 3, and on the east arm 

of the glacier.

For the period May 3 to September 20, 2005, minimum 

average horizontal velocities of the three surviving targets 

declined to about one-fourth to one-third of the corresponding 

values for April 21 to May 3, 2005. For example, the average velocity of a target near the top of spine 4 declined from 

0.28 m/d during the early period (April 21 to May 3) to 0.06 

m/d during the latter period (May 3 to September 20). For the 

same periods, the target near the base of spine 3 slowed from 

0.09 m/d to 0.02 m/d, and the eastside glacier target slowed 

from 0.49 m/d to 0.18 m/d. Steep, rugged, and hot terrain 

on the active dome precluded the use of GPS spiders during 

this period, so no data are available for comparison. Regardless, the observed slowing trend is consistent with what was 

happening in the crater at the time. When the first triangulation measurements were made on April 21, 2005, spine 5 was 

extruding southward, shoving the remnants of spines 3 and 4 

southeastward and compressing the east arm of Crater Glacier, which resulted in thickening and increased rate of flow 

northward (Walder and others, this volume, chap. 13). Spine 

6 emerged in early August 2005 and moved westward for the 

duration of the measurements. The movement of spine 6 eased 

the stress on spines 3 and 4 and on the glacier’s east arm, 

which caused their velocities to slow.

Total-station surveys made by using helicopter-deployed 

targets are a feasible, albeit costly and manpower-intensive 

means to monitor surface displacements on a growing dome. 

This approach might be cost effective in some situations, but 

in our case the measurements were abandoned owing to the 

difficulty of maintaining targets on the dome and the cost of 

helicopter operations required for each survey. The long sight 

lines and consequent inability to obtain reflections from the 

prisms greatly weakened the value of the survey. Although 

the targets costs less than $200 each, their deployment and 

initial survey required about one hour of helicopter flight time 

(~$750/hr), as did each of two repeat surveys spaced 12 days 

and ~4.5 months apart, resulting in ~1 m accuracy at three 
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 291

surviving targets. For comparison, a GPS spider costs about 

$2,500, can be deployed to the dome in about one helicopter 

flight hour and later retrieved if desired. A spider provides 

three-dimensional positioning information every few minutes, 

with an accuracy of a few centimeters, without the need for 

field personnel, except for deployment and retrieval.

Borehole Tiltmeters

Two Pinnacle 5000 series borehole tiltmeters with tilt 

resolution of 5×10-9 radian (3×10-7 angular degree) were 

installed on the 1980–86 dome to record any transient (seconds to days) ground tilt that might accompany extrusion of 

the 2004–5 dome and consequent deformation of the crater 

floor and glacier. The installations were motivated in part by 

a dynamical model in which stick-slip motion of the plug produces repetitive, reversible strain in host rock near the conduit 

(Iverson and others, 2006; Iverson, this volume, chap. 21). If 

each drumbeat earthquake represents an upward “slip” of the 

plug, followed by a few minutes of increasing magma pressure 

while the plug remains “stuck,” drumbeats should be associated with sudden inward tilt followed by gradual outward tilt 

between beats.

To test this hypothesis, the first tiltmeter was installed in 

a 2.3-m-deep hole bored into the September 1984 lobe of the 

1980–86 dome (SEP), 480 m N. 16° W. of the 2004–5 vent, on 

August 5, 2005. A second tiltmeter (REM) was installed in a 

1.8-m-deep hole on the northeast sector of the 1980–86 dome, 

480 m N. 34° E. of the vent, on November 17, 2004. Both 

tiltmeters were modified by the manufacturer to enable onesecond sampling. Data are telemetered to CVO by way of the 

U.S. Forest Service’s Coldwater Ridge Visitor Center.

Analysis of the tiltmeter records is ongoing, and a complete discussion of results is beyond the scope of this paper. 

To date, we have seen no evidence of repetitive, reversible tilt 

events associated with drumbeat earthquakes. It is possible 

that tilts occur but are too small to be resolved by the tiltmeters, or that the tiltmeters are too far away from the source of 

the drumbeats. Fractured host rock surrounding the conduit 

wall or the rubbly interior of the 1980–86 dome might accommodate strains produced by stick-slip motion of the plug without transmitting them to the tiltmeter sites. Without further 

analysis, this null result should not be interpreted as negative 

evidence for the occurrence of stick-slip motion as predicted 

by the model.

Both tiltmeters have recorded tilt changes over time 

scales of a few minutes to weeks that are likely related to the 

extrusion process. The azimuths of tilt vectors for many of 

these events point either toward or away from the extrusion 

site within a few degrees. Consequently, the vectors from 

two tiltmeters tend to intersect close to the vent, indicating a 

shallow pressure source there. It seems plausible that some of 

the tilt signals are caused by relatively subtle changes in the 

shallow supply rate or extrusion rate, perhaps related to continual reorganization of the uppermost part of the conduit as 

the locus of surface activity shifts and the shape of the dome 

changes. Some of the larger earthquakes (Md

 ~2) are associated with tilt events lasting several minutes, but most are not.

A tiltmeter and GPS spider (DOM1), located within 30 m 

of each other on the September 1984 lobe, responded to some 

of the same events. For example, the radial tilt component and 

DOM1 northings increased during September 1–23, 2005, by 

about 30 microradians and 12 cm, respectively. The instruments are nearly due north of the vent, and both measurements 

suggest inflation of the vent area. The simplest explanation is 

that increasing pressure beneath the vent caused the September 

1984 lobe to tilt and move northward. At the time, spine 6 was 

extruding southwestward and migrating westward prior to the 

emergence of spine 7 in mid-October 2005. Northward tilting 

and displacement of the September 1984 lobe presaged the 

emergence of spine 7, possibly during a period of increased 

extrusion rate. Unfortunately, measurements of the lineal 

extrusion rate from Sugar Bowl were impossible in September 

2005 because the camera’s view of spine 6 was blocked by the 

1980–86 dome (see earlier section titled Synoptic Observations of Dome Growth and Glacier Deformation).

Discussion

Although some of the techniques discussed above were 

more successful than others, all of them contributed to our 

evolving understanding of how Mount St. Helens deformed 

before and during eruptive activity in 2004–5. Rather than 

discussing the deformation results separately here, we also 

consider information from other disciplines (seismology, 

gas geochemistry, and petrology) and address a few broader 

issues. Our intent is to characterize current thinking about the 

eruption and to draw attention to some unresolved questions.

Where, and for how long, was the 2004–5 magma stored 

before the beginning of unrest on September 23, 2004? What 

is the nature of the magma reservoir and conduit system? How 

quickly did the 2004–5 magma rise to the surface? Why was 

there no inflation at JRO1 prior to September 23, 2004, given 

that the station clearly responded to deflation and shallow 

seismicity starting on that date? What are the implications for 

the future course of this eruption and for anticipating eruptions 

at similar volcanoes? None of these questions can be answered 

definitively from the geodetic data alone; however, when combined with other datasets, the geodetic data help to place useful constraints on Mount St. Helens’ magma plumbing system 

and to highlight areas for additional research.

Source Models

The details of various source models fit to the geodetic data are discussed elsewhere in this volume, including 

Lisowski and others (chap. 15), Poland and Lu (chap. 18), 

and Mastin and others (chap. 22). Common elements of these 

source models are (1) depth in the range of 5–12 km (bot-
292 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

tom possibly as deep as 20 km), and (2) volume loss through 

December 2005 in the range 15–30×106 m3, compared to 

extruded volume of 73×106 m3. Lisowski and others (this 

volume, chap. 15) modeled and removed the effects of plate 

motion, regional tectonics, and gravitational loading by the 

new dome before inverting CGPS data for a best-fit volcanic source model—a point prolate spheroid with its long 

axis vertical centered near 8 km depth, with a volume loss 

of 16–24×106 m3 through October 31, 2006. Poland and Lu 

(this volume, chap. 18) modeled LOS changes from stacked 

co-eruptive interferograms by using a point source, which best 

fit the data with a depth of 12 km and a volume loss rate of 

27×106 m3/yr. They concluded that a more complicated model 

is not justified by the InSAR data alone, but they agreed with 

other authors in the volume that a vertically elongate source 

is probably more realistic and a better fit to the geodetic 

dataset as a whole. Mastin and others (this volume, chap. 22) 

used a vertical ellipsoidal source to fit surface displacements 

from CGPS data from November 8, 2004, to July 14, 2005. 

They concluded that the top of the source is 5±1 km deep, 

but the basal depth is poorly constrained—possibly 10–20 

km. Their estimate of source-volume loss during that period 

is 11–15×106

 m3, increasing as the bottom of the source gets 

deeper. This corresponds to an average volume-loss rate of 

16–22×106 m3/yr, which we might double to account for the 

period of rapid motion at JRO1 from September 23, 2004, to 

November 8, 2004, which was not included in the model.

These geodetic models generally are consistent with the 

amount of CO2 and other volcanic gases emitted during the 

eruption, which indicate that the erupting magma last equilibrated at 850°C and 130 MPa—conditions corresponding to a 

reservoir 5.2 km deep (Gerlach and others, this volume, chap. 

26). Likewise, samples of the 2004–5 dome indicate that the 

pressure of last phenocryst growth corresponds to that near the 

apex of the magma reservoir at a depth of about 5 km (Pallister 

and others, this volume, chap. 30; Rutherford and Devine, this 

volume, chap. 31). Pallister and others (this volume, chap. 30) 

noted the possibility that the 2004–5 magma might be residual 

from the 1980–86 reservoir and went on to write: “Viewed in 

the context of seismic, deformation, and gas emission data, the 

petrologic and geochemical data can be explained by ascent of 

a geochemically distinct batch of magma into the apex of the 

[1980–86] reservoir during the period 1987–1997, followed by 

upward movement of magma into a new conduit, beginning in 

late September 2004.”

In our opinion, the geodetic, gas-emission, and geochemical data from the 2004–5 eruption all point to involvement of 

magma that was stored in a vertically elongate reservoir centered near 8 km depth and extending upward to about 5 km. 

We believe this is the same reservoir that fed the plinian eruption of May 18, 1980, smaller explosive eruptions during the 

summer of 1980, and dome-building eruptions during 1980–

86, as inferred by Scandone and Malone (1985) from seismic 

data (source centered at 9 km, extending from 7 km to 14 

km) and Rutherford and others (1985) from the mineral phase 

assemblage found in May 18 pumice (7.2±1 km). It seems 

likely that the reservoir was replenished prior to (Moran, 1994; 

Moran and others, this volume, chap. 2) and possibly during 

the current eruption (Mastin and others, this volume, chap. 22; 

Pallister and others, this volume, chap. 30). Data from GPS 

and InSAR show that most of the surface deformation caused 

by pre-eruption reservoir inflation occurred prior to 1991, but 

there is seismic evidence for pressurization of the reservoir 

and upward fluid intrusion starting in 1987 and continuing 

at least until 2001. Apparently, reservoir replenishment and 

surface inflation occurred aseismically until shortly after the 

extended pause in dome building that began in 1986. Thereafter, increasing strain on rock hosting the reservoir and episodic 

intrusion of fluids induced microseismicity but no measurable 

surface deformation.

Implications for the Magma Plumbing System

On September 22, 2004, two things were clear: (1) 

seismicity at Mount St. Helens was at a low background level 

characteristic of long periods between sporadic earthquake 

swarms, and (2) JRO1 was moving northeastward relative to 

stable North America as part of a tectonic block undergoing 

clockwise rotation about a pole located either in south-central 

Oregon (Savage and others, 2000) or along the western part 

of the Oregon-Washington border where it intersects the 

Olympic-Wallowa lineament (McCaffrey and others, 2000). 

The emission rates of volcanic gases probably were low, as 

they had been each time they were measured after a temporary increase in CO2

 emission was noted during increased 

seismicity in spring–summer 1998. On four occasions from 

September 27 to October 2, 2004, measured emission rates of 

magmatic CO2 and SO2 were less than detection thresholds; 

the first notable increase did not occur until October 3, 2004 

(Gerlach and others, this volume, chap. 26).

In hindsight, the September 22 calm belied a volcano 

poised to erupt. A swarm of small earthquakes, all within 2 

km of the crater floor, began around 0200 PDT September 

23. By midday on September 28, cracks large enough to be 

seen from a helicopter flying in the crater appeared in glacier 

ice on the south crater floor. Rapid uplift and intense seismicity continued for several days, and by October 7, a large welt 

growing on the south crater floor had approached the height of 

the 1980–86 dome (fig. 7). The first of several spines emerged 

from the welt on October 11, less than three weeks after the 

peaceful dawn of September 22.

Model-derived estimates of source-volume decrease 

during the first 15 months of the eruption are in the range 

15–30×106

 m3, compared to an erupted volume of 73±4×106

m3

 through December 15, 2005. This suggests, but does not 

require, that the magma reservoir was partially recharged 

during the eruption. The lava extrusion rate declined from 5.9 

m3/s to 2.5 m3/s during the 2.7 months from October 13, 2004, 

to January 3, 2005, but it varied only between 2.4 m3/s and 0.7 

m3/s during the ensuing 14 months through the end of 2005 

(Schilling and others, this volume, chap. 8). Meanwhile, the 
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 293

1980−1986 dome Sep−Oct 2004 welt

Opus area

Crater Glacier

Uplifted

glacier ice

Crater Glacier

Figure 7. Welt growing on south crater floor, October 7, 2004, as seen from Brutus (BRUT) on northeast crater 

rim, Mount St. Helens, Washington. South (left) half of the welt is mantled by uplifted glacial ice and ash from small 

explosions on October 1 and October 5. The north (right) half comprises uplifted crater-floor material and part of the 

south flank of the 1980–86 dome. The latter was dubbed “Opus area” after 1980s-vintage EDM station Opus, which was 

located on south side of graben that formed during May 1985 dome-building episode (Swanson and others, 1987).

JRO1 displacement rate slowed from a peak value of 0.5 mm/d 

in late September–early October 2004 to a steady 0.04 mm/d 

from May through December 2005. These trends suggest that 

the eruption might be approaching a steady-state condition 

in which the extrusion rate equals the recharge rate. If so, the 

situation could persist for years to decades, limited only by the 

volume of magma available for recharge.

Constraints and Conundrums

The geodetic data help to constrain the depth and geometry of the magma plumbing system and the time history of 

magma outflow and possible recharge to the reservoir. However, 

additional questions arise when the geodetic data are considered 

together with other datasets. For example, why did JRO1 immediately respond to the onset of shallow seismicity, presumably 

when magma began its ascent, but did not detect any precursory surface deformation during the preceding seven years? 

The most straightforward answer is that no such deformation 

occurred because the system was poised to erupt magma stored 

in the reservoir or conduit system since the 1980–86 eruption, long before JRO1 was installed in May 1997. This idea is 

consistent with data from trilateration surveys that show surface 

dilatation, presumably caused by magma accumulation beneath 

the volcano, between 1982 and 1991. It is also possible that 

the reservoir inflated during the series of earthquake swarms 

between 1987 and 2001 by an amount that was too small to 

cause measurable surface deformation. Mastin and others (this 

volume, chap. 22) make the point that magma compressibility 

could have significantly dampened any geodetic signal at Mount 

St. Helens. Perhaps reservoir magma is compressible enough 

and the host rock stiff enough that the amount of magma added 

to the reservoir since 1987 was accommodated without measurably deforming the surface.

How did material that resulted in rapid development of 

a welt on the south crater floor and subsequent extrusion of 

highly crystalline, gas-poor magma as gouge-covered spines 

(Cashman and others, this volume, chap. 19) move through or 

past the 1980–86 conduit system without substantially deforming more of the volcano and its surroundings? Station JRO1, 9 

km north-northwest of the 2004–5 vent, recorded no anoma-
294 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

lous motion at millimeter scale before the eruption. A singlefrequency GPS station at SEP on the 1980–86 dome, less than 

500 m from the vent, moved anomalously about 20 cm north, 

8 cm west, and 12 cm up during the 9-month period ending 

on September 27, 2004 (LaHusen and others, this volume, 

chap. 16)—seemingly not enough to accommodate the shallow 

intrusion of a meters-thick dacite magma body. One possibility is that the conduit is so poorly coupled to the rest of the 

volcano that the rising magma column was able to push ahead 

older conduit material, rather than intruding it. Otherwise, 

the conduit walls and surrounding host rock would have to be 

implausibly compliant to accommodate an intrusion without 

causing more widespread deformation.

If the former contents of the conduit were pushed ahead 

of reservoir magma, when did the latter first reach the surface? 

If rapid emergence of the first whaleback-shaped extrusion 

(spine 3) in late October 2004, following the more-labored 

extrusions of spines 1 and 2, marked the first arrival of reservoir magma, then the volume of conduit material flushed 

from the system was about 20×106

 m3—the volume of surface 

deformation plus spines on November 4, 2004. The corresponding radius rc

 for a cylinder extending from the surface to 

depth d = 5 km is 35.7 m, or 28.2 m for d = 8 km. Using the 

volume of material extruded by November 4, 2004 (11.8×106

m3), instead of the volume of surface deformation, we obtain 

rc = 27.4 m for d = 5 km, or 21.7 m for d = 8 km. The average ascent rate through the conduit, vc

, assuming reservoir 

magma began its rise on September 23, 2004, is 120 m/d for d

= 5 km, or 190 m/d for d = 8 km. If, instead, reservoir magma 

arrived at the surface as spine 1 on October 11, 2004, when the 

volume of the welt was 10.1×106 m3, then rc = 25.4 m for d = 

5 km, or 20.0 m for d = 8 km; and the average ascent rate was 

280 m/d (d = 5 km), or 440 m/d (d = 8 km). Thus, for all plausible dimensions of the conduit system, reservoir magma must 

have reached the surface early in the eruption, probably about 

the time CO2 emission rates peaked in October 2004 (Gerlach 

and others, this volume, chap. 26).

Calculated ascent rates in the conduit exceed observed 

lineal extrusion rates, vs = 2–10 m/d, by more than an order 

of magnitude, so the conduit must widen considerably near 

the surface. Thornber and others (this volume, chap. 32) reach 

a similar conclusion (that is, a “wineglass shape” for the 

conduit) based on their interpretation of amphibole-rim thicknesses in post-November 2004 magma. Equating magma flux 

through the conduit to that at the surface for d = 8 km, rc

 = 

28.2 m, vc = 190 m/d, and vs = 10 m/d, we obtain a vent radius, 

rv, = 125 m and an extrusion rate of 5.7 m3/s. Using d = 5 km, 

rc = 35.7 m, vc = 120 m/d, and vs = 10 m/d produces essentially 

the same results, which are representative of the early part 

of the eruption. For comparison, Schilling and others (this 

volume, chap. 8) estimated the vent diameter to be 150–230 

m east-west based on partial exposures of the basal perimeters 

of extruding spines. The north-south vent diameter is unconstrained owing to lack of adequate exposure.

Slowing of the JRO1 displacement rate suggests a corresponding reduction in the net rate at which magma is being 

withdrawn from the reservoir to a value that might be sustainable for the foreseeable future (0.5–1 m3

/s). If the lesser 

withdrawal rate is a result of recharge, the eruption could 

continue indefinitely. On the other hand, in the absence of sufficient recharge, the eruption will end when (1) the reservoir 

is depleted of eruptible magma, (2) the pressure difference 

between reservoir and surface falls below some threshold, or 

(3) friction in the upper part of the conduit chokes off magma 

flow (that is, stick-slip motion of the plug ceases—the plug 

stays stuck). In the latter case, increasing magma pressure 

beneath the plug could eventually result in an explosion; 

however, the low gas content of the 2004–5 magma makes an 

explosion less likely than it would be for gas-rich magma.

When gas-saturated magma rises, decreasing lithostatic 

pressure results in bubble formation. If some of the bubbles 

rise buoyantly through the magma and escape toward the surface, we might expect an eventual uptick in gas emission rates 

if, as we suspect, the Mount St. Helens magma reservoir has 

been recharged during the current eruption. The fact that no 

such increase was observed through the end of 2005 requires 

explanation. Emission rates of CO2

, SO2, and H2S were negligible during September 27–30, 2004, a period characterized by 

scrubbing or sealing-in of gases. Several days of wet degassing 

ensued, when scrubbing by the ground-water system dominated degassing. October 5–6, 2004, marked the beginning 

of a period of dry degassing when emission rates increased to 

800–2,400 metric tons per day (t/d) CO2, 40–250 t/d SO2, and 

0–10 t/d H2S (Dzurisin and others, 2005). Throughout 2005, 

the emission rates of all three gases were near or below the 

low end of the October 2004 ranges (Gerlach and others, this 

volume, chap. 26).

It is possible that the 2004–5 eruption was triggered by 

a recharge event in September 2004 and that gases exsolved 

from the rising magma made their way to the surface by early 

October 2004, causing the observed uptick in emission rates; 

however, this seems unlikely for two reasons. First, the total 

amount of CO2 emitted during the eruption is consistent with 

degassing the volume of magma that reached the surface, 

assuming the magma last equilibrated at about 5 km depth. In 

other words, there is no need to invoke recharge to account for 

the amount of CO2

 emitted. On the contrary, low CO2 emission 

rates imply that the reservoir magma is “flat,” that is, depleted 

of any significant exsolved-gas phase that might have accumulated in its upper part (Gerlach and others, this volume, chap. 

26). Perhaps the catastrophic landslide and explosive eruptions 

of 1980 depleted a volatile-rich cap that formed since the end 

of the Goat Rocks eruptive period in 1857, and any exsolved 

gases that accumulated since then escaped before the current 

eruption began.

The second factor that weighs against a recharge event 

in September 2004 as the cause of increased CO2

 emission 

in early October is the short time interval involved. If, for 

example, recharge at 8 km depth beginning on September 23 

was responsible for newly-exsolved CO2 reaching the surface by October 1, the average ascent rate of the gas was ~1 

km/d. This ascent rate is far greater than, for example, the 
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 295

bubble ascent rate of 1.7 m/hr inferred for the 1975 eruption 

of Mauna Loa volcano, Hawai‘i (Ryan, 1995), even though 

bubble ascent rates in basalt are likely to be much greater than 

in dacite. It is possible that the existence of fracture permeability along the margins of the reservoir could allow this constraint to be bypassed. However, it is more straightforward and 

seems more likely that the source of CO2 reaching the surface 

in early October 2004 was the same magma from the conduit 

or upper part of the reservoir that began extruding shortly 

thereafter. If so, any volatiles exsolved from magma entering 

the lower part of the reservoir since the eruption began have 

not yet fully traversed the length of the reservoir and conduit 

system. This implies an average rise rate less than about 10 km 

in 15 months, corresponding to 0.9 m/hr.

The foregoing paragraphs are not meant to imply that a 

reservoir recharge event could not have occurred during September and October 2004, but rather that such an event, if it 

did occur, might not have been the source of CO2

 that reached 

the surface in the next few days. The CO2 that did reach the 

surface in early October 2004 could have come from magma 

previously stored in the conduit or reservoir. Furthermore, the 

lack of an uptick in CO2 emission rates since peak values were 

measured in October 2004 suggests that the path for CO2 from 

the base of the reservoir to the surface might be complex and 

time-consuming to the extent that the CO2 signature of any 

recharge that occurred since the 2004–5 eruption began has yet 

to reach the surface.

Thus, all three of the traditionally reliable eruption 

precursors (seismicity, ground deformation, and volcanic 

gas emission) failed to provide warning that an eruption was 

imminent at Mount St. Helens until a few days before a visible 

welt appeared at the surface—both in September 2004 and 

March–April 1980 (the north-flank bulge). Sporadic earthquake swarms beneath the volcano from 1987 to 2003, at 

least one of which (1998) was accompanied by increased CO2

emission, provided longer-term but nonspecific forewarning 

of the current eruption. Data from a single-frequency CGPS 

station on the 1980–86 dome might have provided several 

months warning that something anomalous was occurring, 

but the station was not operational during most of that period, 

and its data were not analyzed until after the eruption began. 

Crater-floor deformation might have been hidden beneath 

thick glacial ice for some time before the first surface cracks 

were noticed on September 28.

Implications for Volcano Monitoring

One obvious lesson from the 2004–5 eruption is the value 

of a dense network of CGPS stations operating long before the 

onset of unrest. The need for CGPS data from proximal stations 

was addressed at Mount St. Helens only after the 2004–5 eruption began. Most of the ground deformation that accompanied 

reservoir replenishment following the 1980–86 eruptions had 

ceased by 1991, before installation of the first CGPS station, 

JRO1, in 1997. Playing catch-up with a restless volcano is short 

sighted and dangerous. The best and safest time to install a 

comprehensive, integrated monitoring system is before unrest 

begins, when the need for such a system is less apparent.

There is much we do not understand about the 2004–5 

eruption of Mount St. Helens, but the challenge does not end 

there. On January 11, 2006, Augustine Volcano in Cook Inlet, 

Alaska, began erupting for the first time since 1986. Augustine is a central dome and lava-flow complex surrounded by 

pyroclastic debris—not unlike Mount St. Helens in several 

respects. Both volcanoes are relatively young and frequently 

active. Mount St. Helens is notorious for its debris avalanche 

and eruption of May 18, 1980; Augustine has produced at least 

11 debris avalanches in the past 2,000 years as a result of summit dome collapses (Waitt and Beget, 1996; Beget and Kienle, 

1992). Mount St. Helens’ eruptive products range from basalt 

through dacite (dominantly dacite). Augustine’s products are 

dominantly andesite with small amounts of basaltic andesite 

and dacite. Before their current eruptions, both volcanoes 

last erupted in 1986, and both eruptions were dome-building 

events (Mount St. Helens, 1980–86; Augustine, March–September, 1986).

Given these similarities, we might expect that the precursors to Mount St. Helens’ and Augustine’s most recent 

eruptions were similar, too, and in several respects they were; 

however, there were some striking differences. On November 

29, 2005, the Alaska Volcano Observatory (AVO) released 

the following statement (http://www.avo.alaska.edu/activity/

avoreport.php?view=info&id=342&type=info&month=Novem

ber&year=2005, last accessed January 14, 2008):

Beginning in May 2005, there has been a slow increase in the 

number of earthquakes located under Augustine Volcano. The earthquakes are generally small (less than magnitude 1.0) and concentrate roughly 1 km below the volcano’s summit. These earthquakes 

have slowly increased from 4–8 earthquakes/day to 20–35 earthquakes/day. Additionally, data from a 6-station Global Positioning 

System (GPS) network on Augustine Volcano indicate that a slow, 

steady inflation of the volcano started in mid-summer 2005 and continues at present. The GPS benchmark located nearest the summit 

has moved a total of 2.5 cm (1 inch). This motion is consistent with 

a source of inflation or pressure change centered under the volcano. 

This is the first such deformation detected at Augustine Volcano 

since measurements began just prior to the 1986 eruption.

Small explosions began on December 9, 2005, and there were 

reports of unusual “steaming” and sulfur smell downwind 

of the volcano, as well as a visible, condensed-steam plume 

extending at least 75 km from the volcano on December 12. 

The eruption began in earnest on January 11, 2006, characterized by magmatic explosions, pyroclastic flows and lahars, 

and extrusion of a new lava dome and flow that was first 

observed on January 16, 2006. Emission rates for SO2

 were 

about 1,000 t/d in early January and several thousand t/d during the first week of the eruption. Corresponding CO2

 emission rates were 2–3 times greater (Cervelli and others, 2006; 

Power and others, 2006).

The Augustine and Mount St. Helens eruptions were 

similar in the following respects:
296 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

 1. No precursory inflation of a deep (>5 km) source 

 was detected; modeling of GPS data indicates the preerup-

 tive inflation source at Augustine was 1–2 km beneath the

 volcano’s summit, at a depth of approximately sea level

 (Cervelli and others, 2006),

 2. Deep deflation was associated with extrusion,

 3. Phreatic explosions preceded the beginning of

 extrusion,

 4. Energetic swarms of shallow earthquakes occurred 

 in the days prior to the first explosion, and

 5. Erupted products were predominantly effusive, 

 with only minor tephra.

The eruptions differed in the following ways:

 6. There were eight months of elevated seismicity at 

 Augustine versus a few days of elevated seismicity at

 Mount St. Helens,

 7. There were six months of precursory inflation at 

 Augustine versus no precursory inflation at 

 Mount St. Helens (but see below), and

 8. There were several thousand tons per day of SO2

 and CO2 recorded at Augustine versus 40–250 t/d SO2

 and 800–2,400 t/d CO2 recorded at Mount St. Helens 

 (Gerlach and others, this volume, chap. 26).

The differences in items 6 and 8 are supported by unambiguous seismic and gas emission-rate data, but the apparent lack of 

precursory inflation at Mount St. Helens (item 7) could be an 

artifact of inadequate geodetic monitoring. Station JRO1 might 

have been too distant (9 km) to detect inflation of a shallow source 

comparable to the one that inflated for months at Augustine. Even 

so, relatively deep sources at both volcanoes deflated during the 

eruptions but did not inflate in the years or months beforehand—a 

pattern unlikely to be sustainable over several eruption cycles. A 

possible explanation for this observation is that inflation of deep 

magma sources beneath arc volcanoes occurs mainly early in the 

eruption cycle, soon after an eruption ends, and slows as the next 

eruption approaches. This pattern has been observed in InSAR 

studies of Westdahl and Okmok volcanoes in the Aleutian arc by 

Lu and others (2003, 2005). Lu and his colleagues suggest that the 

magma supply rate is governed by the pressure gradient between 

a deep source and shallow reservoir. The pressure gradient and 

hence the flow rate are greatest immediately after eruptions. Pressurization of the reservoir decreases both the pressure gradient 

and flow rate, but eventually the reservoir ruptures, an eruption 

or intrusion occurs, and the cycle starts anew. This scenario is 

consistent with trilateration and campaign GPS results for Mount 

St. Helens that showed dilatation between 1982 and 1991, but no 

measurable deformation from 1991 to 2000 or from 2000 to 2003.

Given that the latest repose periods at Mount St. Helens 

and Augustine were similar, and that the magma erupted at 

Mount St. Helens is more viscous (greater SiO2

 content and 

crystallinity, lower temperature), why did Mount St. Helens 

erupt with less precursory seismicity and (perhaps) surface 

deformation? An important difference between the two eruptions is the higher gas-emission rates at Augustine. Gassy 

magma suggests a deep source, and the prolonged period of 

elevated seismicity and surface inflation at Augustine probably reflect the accumulation of rising magma in a shallow 

reservoir. In contrast, there are indications that gas-poor 

2004–5 magma had already accumulated in a crustal reservoir 

long before the September 2004 earthquake swarm at Mount 

St. Helens. Inflation might have occurred anytime between 

1982 and 1991, as suggested by trilateration and GPS results, 

episodically during repeated earthquake swarms between 1987 

and 2001, or continuously and mostly aseismically.

Conjecture Regarding the Link Between 1980–

1986 and 2004–2005 Eruptions

Continued low gas-emission rates at Mount St. Helens 

more than a year after the eruption began, and after extrusion of more than 73×106

 m3 of dacite, suggest that the 

eruption is being fed from a reservoir of mostly degassed 

magma. The most likely source is magma left over from the 

1980s. The May 18, 1980, eruption was unusual in that it 

was triggered by a large landslide at 0832 PDT that unloaded 

the magmatic system suddenly and tapped gas-rich magma 

from a reservoir at 8 km depth. A devastating lateral blast 

and debris avalanche produced the most compelling images 

from the eruption, but an important change in the eruption’s 

character did not occur until midday. Rowley and others 

(1981, p. 492, 489) reported: “At about noon, the eruption 

cloud lightened in color from medium gray to dirty white * 

* * (D.A. Swanson, oral commun., 1980).” and “Pyroclastic flows were first observed being emplaced at 1217 PDT 

(Pacific Daylight time) on May 18, nearly 4 hr after the start 

of the eruption, and successive pyroclastic flows continued to 

form intermittently for about 5 hr thereafter.” It seems likely 

that the change in the eruption cloud at midday corresponded 

to the first arrival at the surface of gas-rich magma from a 

deeper source, presumably the 7-km-deep reservoir proposed 

by Rutherford and others (1985). This idea is consistent with 

the onset of vigorous harmonic tremor, indicative of magma 

movement, just before noon, and with its decline at about 

1730 PDT (Malone and others, 1981).

If this scenario is correct, the hours-long delay between 

the initiation of the landslide and the first arrival of reservoir 

magma at the surface suggests that the latter might not have 

occurred without the former, that is, that the landslide interrupted a sequence of events that otherwise might not have 

tapped the reservoir on May 18. Without the landslide trigger, 

magma intruding the volcano to cause the famous north-flank 

bulge might have reached the surface in less spectacular fash-
14. Constraints and Conundrums Resulting from Ground-Deformation Measurements Made During 2004–2005 297

ion and might not have been followed by gas-rich magma from 

the reservoir. We speculate that catastrophic unloading of the 

magmatic system by the landslide decreased the reservoir-tosurface distance, thus increasing the pressure gradient abruptly. 

As a consequence of the greater pressure gradient, rapidly 

rising magma created and maintained an open conduit to the 

surface for 9 hours—enough time to release any separate gas 

phase that might have accumulated near the top of the reservoir. 

There would have been ample time, since the end of the previous eruption in 1857, for a volatile-rich cap to form by vesiculation and bubble-rise in the reservoir. Magma that rose into 

the lower-pressure environment of the reservoir prior to 1980 

would have been temporarily oversaturated in volatiles. The 

“excess” volatiles would have exsolved to form bubbles that 

tend to aggregate and migrate toward the top of the reservoir. 

Given enough time and low enough permeability to the surface, 

a volatile-rich cap develops on the reservoir. The physics of this 

process is beyond the scope of this paper. We suggest only that 

the May 18, 1980, landslide might have initiated a cascade of 

events that resulted in, among the more obvious consequences, 

catastrophic loss of the gas-rich upper part of the reservoir. 

This set the stage for dome-building eruptions that followed in 

1980–86 and 2004–5, which tapped reservoir magma that had 

been separated from its excess volatiles and only partly replenished during the interim.

The presence in the upper crust of magma remnant from 

the 1980s and earlier, which was devoid of its excess volatiles 

as a result of the May 18, 1980, eruption, might explain some 

aspects of the 2004–5 eruption. For example, the absence of 

precursory reservoir inflation is consistent with the idea that the 

magma erupted in 2004–5 was already present in the reservoir and conduit system when JRO1 was installed in 1997. 

Likewise, the short duration of precursory seismicity and the 

low gas emission rates throughout the 2004–5 eruption can be 

explained if the first few days of unrest involved the mobilization of magma that had been stored in the upper part of the 

conduit since the end of the 1980–86 eruption. This was soon 

followed by the rise of magma that arrived in the reservoir 

sometime before 1980, equilibrated to the pressure and temperature conditions in the reservoir, and lost its excess volatiles 

during the eruption on May 18, 1980. This explanation does 

not preclude partial recharge of the reservoir prior to or during 

the 2004–5 eruption. Mastin and others (this volume, chap. 22) 

conclude from their analysis of the geodetic data that “ * * * 

erupted magma has been replaced in increasing proportions by 

recharge, but that the recharge rate remains somewhat less than 

the current effusion rate.” Both the velocity of JRO1 and the 

effusion rate declined throughout 2006–7. During part or all of 

that period, the recharge rate might have been comparable to the 

effusion rate (L. Mastin, oral commun., 2008).

Although mostly speculative, we believe the following scenario is consistent with the current state of knowledge concerning the 1980–86 and 2004–5 eruptions. Both eruptions were fed 

from a reservoir centered near 8 km depth; magma erupted in the 

1980s accumulated in the reservoir during an extended period, 

long enough for gas bubbles to rise and form a volatile-rich cap 

(Gerlach and McGee, 1994). The May 18, 1980, landslide and 

plinian eruption depressurized the reservoir catastrophically, 

releasing the volatile-rich cap and causing additional bubbles to 

form throughout the reservoir. Most of those bubbles managed to 

rise to the top of the reservoir at about 5 km depth and escape via 

the 1980 conduit system prior to the start of the 2004–5 eruption, leaving behind a large volume of gas-poor magma—at least 

several cubic kilometers according to Mastin and others (this volume, chap. 22). The 2004–5 eruption was preceded by and has 

been accompanied by reservoir recharge, which is continuing at 

a rate of ~1 m3

/s (early 2006). Any magmatic gas signature from 

recharge might be delayed and muted by the large vertical extent 

of the reservoir and relatively high viscosity of reservoir magma. 

Factors that will influence the future course of the eruption (for 

example, slowing and eventual cessation of dome growth, or a 

transition to either more explosive activity or effusion of more 

mafic magma) include (1) the recharge rate going forward, (2) 

bubble-rise rates in the reservoir–conduit system, and (3) the 

character of magma entering the base of the reservoir (that is, gas 

content and melt composition).

Volcanoes are complex natural systems. Effective hazards 

mitigation, even at well-monitored volcanoes such as Mount 

St. Helens, requires constant vigilance and better understanding of a wide variety of physical, chemical, and hydrologic 

processes that interact to produce eruptions. The events of late 

2004 and 2005 at Mount St. Helens amazed and confounded us 

as they unfolded, and to some extent they still do. What were 

the perceived odds on September 22, 2004, that an enormous 

welt would rise on the crater floor in a matter of days, without 

producing far-field deformation, deep seismicity, or greatly 

increased gas emission? Or that the top of a dacite magma 

column would emerge from the welt in less than three weeks 

and extrude for more than a year as a series of gouge-covered 

spines, with older spines shunted aside as newer ones emerged? 

Or that during the eruption, Crater Glacier would be split in two 

and crumpled against the crater walls without producing even a 

trickle of water flow at the crater mouth? Clearly, there is much 

left to learn at this volcano before we can understand how, when, 

and why such amazing events occur, and to what extent they 

might be hazardous.

Acknowledgments

This is an overview and ideas paper, mostly devoid of its 

own data. The authors make no claim of originality or exclusivity. We accept responsibility for bad or half-baked ideas 

and credit any good ones to the eruption team at the USGS 

Cascades Volcano Observatory. We thank L.G. Mastin and P.F. 

Cervelli for helpful reviews.
298 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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magma discharge and readjustment of the feeding system 

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 15

1

U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Analysis of GPS-Measured Deformation Associated with 

the 2004–2006 Dome-Building Eruption of 

Mount St. Helens, Washington

By Michael Lisowski1, Daniel Dzurisin1, Roger P. Denlinger1, and Eugene Y. Iwatsubo1

Abstract

Detecting far-field deformation at Mount St. Helens 

since the crater-forming landslide and blast in 1980 has been 

difficult despite frequent volcanic activity and improved 

monitoring techniques. Between 1982 and 1991, the systematic extension of line lengths in a regional GPS trilateration network is consistent with recharge of a deep magma 

chamber during that interval. The rate of extension, however, 

averages only 3 mm/yr, and some of this apparent deformation may result from systematic scale error in the electronic 

distance measurements. Subsequent GPS surveys and data 

from a continuous GPS station, located 9 km north of Mount 

St. Helens and operating since 1997, show no significant 

volcanic deformation until the start of unrest on September 

23, 2004. The current eruption has been accompanied by 

subtle but widespread inward and downward movement of 

GPS monitoring stations, exponentially decreasing with time 

and totaling as much as 30 mm. The observed deformation is 

consistent with the predictions of an elastic half-space model 

of a vertically elongate magma chamber with its center at 

a depth of around 7 to 8 km and with a total cavity-volume 

loss of about 16–24×106

 m3. The discrepancy between the 

estimated cavity-volume loss and the >83×106-m3 volume 

of the erupted dome can be explained, for the most part, by 

exsolution of gas in the stored magma and by minor input of 

new magma during the eruption.

Introduction

The current episode of volcanic unrest at Mount St. 

Helens began suddenly with a shallow earthquake swarm on 

September 23, 2004. Geodetic monitoring of deformation in 

the months before and days after the start of unrest was limited to data from a single, continuous, dual-frequency GPS 

(CGPS) station, JRO1 (installed in 1997), having an antenna 

on the roof of Johnston Ridge Observatory (JRO) (fig. 1). 

Regional deformation in the years before the start of unrest 

was measured using campaign-style GPS surveys of a 43-station network during the summers of 2000 and 2003. Annual 

surveys of background deformation in the crater of Mount 

St. Helens began in 2000 at six GPS stations on and near the 

1980s dome (“old dome”) for the purpose of ground control 

(Schilling and others, this volume, chap. 8). A network of 

three single-frequency (L1) CGPS stations was installed in 

2000 to monitor deformation on the September 1984 lobe of 

the 1980s dome, relative to two stations just outside of the 

crater (LaHusen and others, this volume, chap. 16). Only 

one of these L1 CGPS stations, SFOT (on the west flank), 

remained operational during most of 2004, and there is no 

GPS measurement of deformation in the crater until September 28, 2004. In the 1980s and 1990s, deformation of Mount 

St. Helens was measured by using (1) electronic distance 

meters (EDM) to track line-length changes, (2) theodolites 

to track angle changes, and (3) level arrays and electronic 

tiltmeters to monitor ground tilt (Dvorak and others, 1981; 

Lipman and others, 1981; Swanson and others, 1981; Ewert 

and Swanson, 1992).

The observed patterns and rates of deformation in 

volcanic areas are used to constrain possible sources and 

mechanisms of volcanic unrest. We did not detect a systematic pattern of deformation, other than secular tectonic 

motion, between the 2000 and 2003 campaign GPS surveys; nor were there anomalies in the JRO1 CGPS threedimensional (3D), position-component time series before 

the start of unrest. By the time of the first phreatic explosion 

on October 1, 2004, the L1 CGPS network was restored to 

operational condition (the station on the September 1984 
302 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

lobe of the dome was destroyed by ballistics on October 1) 

and several campaign GPS stations on the volcano’s flanks 

were occupied. Within two weeks, additional CGPS stations 

were installed by the Cascades Volcano Observatory (CVO) 

and the Plate Boundary Observatory (PBO). Displacement 

of these stations during the eruption was small (a few centimeters maximum) and consistently inward toward Mount 

St. Helens and downward, the pattern of movement expected 

if a deep magma chamber were being depleted during the 

eruption. The deformation, however, is small given the volume of lava erupted at the surface. The pressure loss in the 

deep magma chamber that is feeding the eruption may be 

compensated by recharge and by exsolution of gas (Mastin 

and others, this volume, chap. 22).

We begin by discussing our methods of GPS data reduction and analysis, emphasizing the technique used to improve 

measurement precision by removing common-mode noise. 

We then use regional GPS surveys and CGPS data to calculate the pattern and rate of background tectonic deformation 

and to derive a simple tectonic model. Local deformation near 

Mount St. Helens is then calculated from EDM-measured, 

line-length data collected during the 1980s and 1990s and 

from GPS-measured position-change data collected between 

2000 and 2003; these results are compared with the predictions indicated in the tectonic model. The time series of 

deformation at CGPS station JRO1 is examined carefully for 

anomalies in the years before the start of the current episode 

of volcanic unrest at Mount St. Helens. Finally, eruptionrelated deformation is calculated, and an elastic, half-space 

model of a deflating, spheroidal magma chamber is used to fit 

the observed deformation.

GPS Data Collection and Analysis

Campaign GPS and Continuous Dual-Frequency 

GPS Networks

Regional and local deformation around Mount St. Helens 

is measured episodically with campaign GPS surveys and 

continuously with CGPS stations (fig. 2). The following is a 

simplified discussion of how GPS positions are determined. 

The basic principal is that a GPS position can be triangulated 

from distance measurements to several known locations. The 

GPS satellites transmit encoded timing signals (pseudoranges) 

at two carrier frequencies, L1 (1575.42 MHz) and L2 (1227.60 

MHz). The orbits of the satellites are predictable, making 

them reference points from which the ground position can 

be triangulated. The distance is measured directly by using 

the travel time of the coded signal, but it is called a pseudorange because the satellite and ground clocks used to measure 

the travel time of the one-way satellite transmission are not 

synchronized. The most accurate measurement of distance 

is obtained from the relatively short-wavelength, continuous 

L1 and L2 carrier signals and from the so-called “ionospheric 

free” linear combination of L1 and L2 known as LC or L3. 

Generally, the LC observable is used to minimize scale error 

introduced by fluctuations in ionospheric activity. With the 

carrier signals, the range to the satellite is expressed as the 

sum of the observed carrier phase plus an ambiguous integer 

number of carrier wavelengths. Methods have been developed 

to determine or “fix” the phase ambiguities to their correct 

integer values, which removes them as an unknown in the 

solution and improves measurement precision.

The GPS data discussed here are acquired with geodeticquality receivers that record pseudorange and carrier-phase 

data at least every 60 s, and most often at rates of 30, 15, or 

10 s. During campaign GPS surveys, we collect data for a 

minimum of 6 hours (usually 24 hours) on multiple days.

Most of the campaign GPS stations around Mount St. 

Helens were surveyed initially in 2000, with the last complete 

survey of the network in 2003. Several stations were surveyed 

after the start of unrest on September 23, 2004, and again in 2005 

and 2006. The center-punched, permanent tablets marking the 

campaign GPS stations generally are set in bedrock to provide a 

stable, long-term, ground-reference point. The only measure of 

deformation near Mount St. Helens at the start of unrest was provided by CGPS station JRO1. Other CGPS stations, operated by 

the Pacific Northwest Geodetic Array (PANGA) and the National 

Geodetic Survey’s Continuous Operating Reference Stations 

(CORS), are found in the surrounding region and serve as stable 

local reference stations. Following the initial seismic swarm, a 

small number of the campaign GPS stations were reoccupied to 

measure deformation accumulated since 2003 and possible ongoing deformation during the early stages of the eruption.

By mid-October 2004 a new network of CGPS stations, 

installed by the USGS–CVO and the National Science Foundation’s PBO, were tracking deformation associated with the 

Figure 1. Radome-covered antenna for CGPS station JRO1, 

on the roof of the U.S.D.A. Forest Service’s Johnston Ridge 

Observatory, 9 km north of Mount St. Helens, Washington. 

In the weeks before and days after September 23, 2004, 

when seismic unrest began, JRO1 was the only GPS station 

operating within 40 km of the volcano. 

USGS photo by E.Y. Iwatsubo.
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 303

HOWE

GP06

REDD

MOWI

LAR5 A545

TMBL

WHEY

STAA

MORT

PACK FL90

SARG

SPR4

DALL

CAST

RKBU

PRES MARS YALE

KELS

GOBS

GWEN

CVO1

JRO1

MITC

PAN1

ARIL

4100

5040

BGRM

VANS

0 50

KILOMETERS VELOCITY SCALE

10 MM/YEAR

 123 ̊00' 122 ̊30' 122 ̊00' 121 ̊30' 121 ̊00' 120 ̊30'

45 ̊00'

45 ̊30'

46 ̊00'

46 ̊30'

MOUNT ST.

HELENS INSET

MOUNT

HOOD

MOUNT

ADAMS

COLUMBIA R IVER

MACK

STEV

WOOD

HELN

Fig. 4

Figs. 6, 9, 15, 17, 18

Figs. 2, 5

0 300

KILOMETERS

OREGON

WASHINGTON

Figure 2. Map showing GPS stations in southwest Washington and northwest Oregon and their average 

horizontal-velocity vectors relative to stable North America (tipped with 95-percent-confidence error ellipses, 

vector scale at bottom) through September 2004. CGPS station names and velocity vectors are shown in 

blue. Black vectors at CGPS stations are calculated only from data coincident with campaign GPS surveys. 

Velocities at stations marked with yellow triangles are used to calculate background deformation for a tectonic 

model, which includes rigid-block motion and uniform-strain accumulation. The position of many stations was 

measured several times using campaign GPS, with some measurements made as early as 1992. However, the 

positions of most stations near Mount St. Helens (red triangles) were measured only in summers of 2000 and 

2003. The Helens trilateration network, measured in 1982 and 1991, is shown here with faint solid lines between 

stations. The dashed box marks the local area around Mount St. Helens shown in figure 6.

eruption (table 1). Additional PBO and CVO CGPS stations 

were added in 2005 and 2006. In order to expedite obtaining 

site usage permits, and to keep the antennas above the winter 

snow pack, many of the new CGPS antennas were placed at 

the top of existing “winterized” 3- to 3.5-m-high steel towers that originally were installed in the fall of 1980 for EDM 

measurements (Swanson and others, 1981). Burial of the GPS 

antennas by snow has not been a problem, but ice buildup 

around the antenna and tower, particularly for stations high 

on the volcano’s flank, has caused pseudodisplacements, two 

warped towers, and many communication and power failures.

GPS Data Processing: Minimally Constrained 

Daily Solution

We process GPS data with GIPSY/OASIS II software 

(Webb and Zumberge, 1995). Daily point-positioning solutions (Zumberge and others, 1997) use satellite orbit and clock 
304 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Station Agency Date

operational Monument type Latitude(degrees)

Longitude

(degrees)

Ellipsoidal 

elevation

(m)

GOBS PANGA 08/27/1997 Deep-drilled and braced 45.8388 -120.8147 621.49

KELS PANGA 10/09/1997 Deep-drilled and braced 46.1182 -122.8961 -15.92

JRO1 CVO 05/23/1997 Roof mount 46.2751 -122.2176 1275.61

REDM PANGA 07/24/1998 Shallow pier 44.2598 -121.1479 919.85

LIND PANGA 07/28/1998 Roof mount 47.0003 -120.5390 471.93

LINH PANGA 04/04/2002 Roof mount 47.0000 -120.5390 472.28

CVO1 CVO 01/04/2003 Roof mount 45.6109 -122.4961 66.67

P432 PBO 06/20/2004 Deep-drilled and braced 46.6228 -121.6832 318.88

P420 PBO 06/25/2004 Deep-drilled and braced 46.5886 -122.4961 74.11

P421 PBO 12/15/2004 Deep-drilled and braced 46.5318 -122.4292 220.60

TWRI CVO 10/06/2004 EDM reflector tower 46.1979 -122.2119 2070.81

TWIW CVO 10/11/2004 EDM reflector tower 46.2129 -122.1587 1630.55

P697 PBO 10/14/2004 EDM reflector tower 46.1876 -122.1766 2214.61

P693 PBO 10/14/2004 EDM reflector tower 46.2103 -122.2024 2113.00

P695 PBO 10/15/2004 EDM reflector tower 46.1990 -122.1642 2015.68

P696 PBO 10/15/2004 EDM reflector tower 46.1969 -122.1516 1600.03

TSTU CVO 10/15/2004 EDM instrument tower 46.2369 -122.2241 1248.65

P687 PBO 10/16/2004 Shallow-drilled and braced 46.1096 -122.3546 391.10

TGAU CVO 10/21/2004 EDM reflector tower 46.2192 -122.1923 1618.28

P698 PBO 11/04/2004 EDM reflector tower 46.1735 -122.1606 1479.39

P702 PBO 11/05/2004 Shallow-drilled and braced 46.3002 -122.3456 1305.70

P690 PBO 02/01/2005 EDM reflector tower 46.1800 -122.1899 2078.91

P699 PBO 02/02/2005 EDM reflector tower 46.1898 -122.2032 2274.11

P689 PBO 10/23/2005 Shallow-drilled and braced 46.1896 -122.3606 1359.46

THAR CVO 06/28/2006 EDM instrument tower 46.2753 -122.1740 1430.55

P701 PBO 09/22/2006 Shallow-drilled and braced 46.1946 -122.1333 1247.79

P694 PBO 09/24/2006 Shallow-drilled and braced 46.2996 -122.1819 1726.03

P692 PBO 09/25/2006 Shallow-drilled and braced 46.2245 -122.1842 1492.23

P700 PBO 09/26/2006 Shallow-drilled and braced 46.1781 -122.2173 1522.03

P691 PBO 09/29/2006 Shallow-drilled and braced 46.2315 -122.2269 1183.26

P703 PBO 09/30/2006 Shallow-drilled and braced 46.1453 -122.1963 988.75

P705 PBO 09/30/2006 Shallow-drilled and braced 46.1730 -122.3106 1151.23

Table 1. Regional and local continuously transmitting GPS (CGPS) stations near Mount St. Helens, Washington.

[Stations are listed in order of installation. PANGA, Pacific Northwest Geodetic Array; CVO, Cascades Volcano Observatory; PBO, Plate Boundary 

Observatory. See figure 2 for locations of most stations. Station names in bold type are within 20 km of Mount St. Helens.]
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 305

files obtained from the Jet Propulsion Laboratory (Pasadena, 

Calif.). The methods described by Blewitt (1989) are applied 

to the daily solutions of a network of stations, which improves 

measurement precision by estimating carrier-phase ambiguities within overlapping subsets of stations. The resulting 

minimally constrained network solution is in the reference 

frame of the satellite orbits (a nonfiducial reference frame—

for example, Helfin and others, 1992). The GPS data processing to this point is standardized by using gp, a set of UNIX 

scripts and programs developed by Will Prescott and others at 

the USGS in Menlo Park, Calif. (see http://quake.wr.usgs.gov/

research/deformation/gps/gpmanual/index.html).

Transformation of Minimally Constrained 

Solution to a Terrestrial Reference Frame

The next step in the GPS data-reduction process is 

experimental and has changed over time. The daily, minimally 

constrained, fixed-ambiguity network solution is merged with 

a point-position solution of a global network of CGPS stations 

(http://quake.wr.usgs.gov/research/deformation/gps/auto/

Track/) that are in the stable interiors of the Earth’s lithospheric plates; their positions and velocities are defined in the 

ITRF2000 terrestrial reference frame (Altimimi and others, 

2002). No ambiguity resolution is attempted for the global 

network because of the large station separations. Station positions in the ITRF2000 frame station evolve following the nonet-rotation NNR-NUVEL-1A model of DeMets and others 

(1994). We are interested in deformation relative to the North 

American plate (NOAM), and for this purpose we created a 

nominal fixed-NOAM ITRF2000 reference frame using the 

Euler vector for NOAM given by Altimimi and others (2002, 

their table 6). We refer herein to this reference frame as “fixedNOAMgp.”

The combined (global and local), minimally constrained 

daily solution is transformed into ITRF2000 and to the 

fixed-NOAMgp terrestrial reference system by applying a 

seven-parameter Helmert transformation (x,y,z translations + 

rotations about the three axes + a scale change) that minimizes 

the misfit in the observed positions of the global reference stations with the positions predicted for that day. Misfit residuals 

from this adjustment are examined automatically to identify 

outliers, a reference station is removed from the solution if its 

residual exceeds five times the expected error, and the transformation process is iterated. Finally, the minimally constrained network solution and the combined ITRF and nominal 

fixed-NOAMgp solutions are archived, and Web-accessible 

plots of station time series and estimates of station velocities 

are updated.

GPS Networks Near Mount St. Helens

Campaign GPS and CGPS data from stations near Mount 

St. Helens are included in three regional groupings (networks): 

Cascadia includes most regional CGPS stations in the Pacific 

Northwest; Helens includes Mount St. Helens campaign GPS 

stations along with a subset of the regional CGPS stations 

that are processed only on the days campaign GPS data are 

collected; and HelensMonit includes CGPS stations around 

Mount St. Helens, a few regional GPS stations, and campaign GPS stations that provide bedrock local ties to the 

Mount St. Helens CGPS stations (for example, see http://

quake.wr.usgs.gov/research/deformation/gps/auto/Helens/). 

These regional groupings are arbitrary and have resulted 

from the combination of a rapid increase in the number of 

stations over time and from the computational efficiency of 

processing data from networks having a smaller number of 

stations. New analysis techniques that detect and remove 

common-mode error in the station positions are most successful when all Mount St. Helens and related referencestation data are processed in a single group.

Local Terrestrial Reference Frames

The precision of regional GPS solutions can be improved 

by removing spatially correlated noise (Wdowinski and others, 1997). The transformation to a terrestrial reference frame 

is approximate, and it can be refined by using local stations 

(within a few hundred kilometers) with well-defined positions and velocities to estimate common-mode daily bias in 

the local solutions. Common-mode daily bias is estimated 

by either stacking the position component residuals (after 

removal of linear trends) of the local reference stations or by 

the equivalent procedure of estimating origin translations of 

the daily network solutions. We estimate the local network 

translations from a subset of CGPS reference stations included 

in the solutions.

We construct a local reference frame for the Mount St. 

Helens region to define the position of selected CGPS reference stations at any given time. The reference stations have 

long measurement histories, daily position measurements, and 

constant velocities. Starting with the daily combined, fixedNOAMgp solutions, we remove all but the local reference stations and a subset of NOAM global stations. These solutions 

are merged to produce weekly averages of positions, outliers 

are identified and removed, and then the weekly averages for 

multiple years are adjusted to give an average position at a 

particular time and average station velocities (see Smith and 

others, 2004). We include the subset of ITRF2000 NOAM stations to keep the local reference frame aligned with our fixedNOAMgp reference frame during the adjustment. Thereafter, 

most NOAM stations are removed from the final local reference frame because they are far from the region of interest.

The daily local solutions, which include several of the local 

reference stations, are then shifted into the local reference frame 

by applying a three-parameter Helmert transformation (3D 

translation), which minimizes the misfit between the observed 

positions of the local reference stations and those predicted by 

the local reference frame for that date. Occasionally, one or 

more of the local reference stations is poorly fit; it is removed 

from the regional solution and the transformation is iterated.
306 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Removing the common-mode daily bias in the solutions 

can reduce the scatter in the station-component time series by 

a factor of 2 to 3 (Williams and others, 2004). Local-station 

seasonal wander, data offsets, and nonlinear deformation 

remain in the station-component time series after removing 

spatially correlated noise common to the reference stations. 

We use the QOCA (pronounced “coca”) software (QuasiObservation Combination Analysis) (Dong and others, 1998, 

2002, 2006), designed and developed at NASA’s Jet Propulsion Laboratory, for final analysis of the trends and noise in 

the GPS solutions. The QOCA software can be used to align 

the minimally constrained network solutions to any defined 

terrestrial reference frame, manipulate 3D station-position 

time series, and calculate spatial and temporal deformation.

Deformation in the Mount St. Helens 

Region Determined by Using GPS and 

Trilateration Survey Data

Tectonic interactions between the Pacific, Juan de Fuca, 

and North American plates displace and deform the ground in 

the Pacific Northwest (for example, Savage and others, 1991; 

Murray and Lisowski, 2000; McCaffrey and other, 2000; Savage and others, 2000; Miller and others, 2001; Svarc and others, 2002; Mazzotti and others, 2003). The widespread secular 

tectonic deformation can mask the more localized background 

and episodic deformation across the volcanic arc. For example, 

locking of the shallow interface between the Juan de Fuca and 

North American plates compresses the adjacent coastal region 

of the North American plate in the east-northeast direction 

of relative convergence between the plates. More enigmatic 

are the rigid-body movements of large continental regions 

(microplates) and possible strain accumulation or relative 

motion along their boundaries (for example, Wells and others, 

1998; Wells and Simpson, 2001).

We use repeated GPS surveys and data from CGPS stations to derive a model for background tectonic deformation 

in the Mount St. Helens region. We then compare the modeled 

background rates of strain accumulation with that determined 

from line-length changes in an arc-crossing, regional trilateration network that was measured in 1982, 1991, and 2000. 

Finally, the tectonic model is used to adjust observed GPS station movements to determine the rate and pattern of volcanic 

deformation around Mount St. Helens since 2000.

Tectonic Rigid-Body Rotation and Strain 

Accumulation in the Mount St. Helens Region

The Helens GPS network includes several regional campaign stations with well-determined velocities, some occupied 

as early as 1992 (stations marked with yellow triangles in figs. 

2 and 3). In addition, solutions for the Helens and Cascadia 

networks include several PANGA CGPS stations with welldetermined velocities that are situated in the arc and back-arc 

regions (stations with blue names in figs. 2 and 3).

The GPS station velocities in the fixed-NOAMgp terrestrial reference frame reveal what appears to be a regional, 

clockwise, rigid-block rotation (figs. 2, 3). This microplate 

motion is roughly consistent with paleomagnetic studies that 

suggest widespread Cenozoic rotation in the Cascadia forearc 

(for example, Simpson and Cox, 1977; Magill and others, 

1982; Wells and others, 1998) and is similar to widespread 

block rotation observed by using GPS to survey much of the 

Cascadia forearc, arc, and backarc region (McCaffrey and 

other, 2000; Savage and others, 2000; Svarc and other, 2002). 

In addition to the obvious block rotation, there may be a small 

amount of internal deformation (strain accumulation).

We use the method of Savage and others (2001, their 

appendix A) to simultaneously estimate an Euler vector for 

the rigid-body rotation and the rate of strain accumulation that 

best fits the observed regional-station velocities. Our bestfitting tectonic model was obtained by excluding the northern 

PANGA CGPS stations (CHWK, DRAO, SEDR, BREW) from 

the model. The resulting Euler vector, principal strain rates, and 

predicted station velocities are shown in figure 3 and summarized in table 2 (Helens GPS). An independent estimate of 

the rotation rate and strain was computed using QOCA, which 

gave similar results (table 2, Helens GPS QOCA). The computed Euler vector is roughly consistent with previous studies 

of GPS-measured deformation in Cascadia, which show a 

similar rate of clockwise block rotation (McCaffrey and other, 

2000; Savage and others, 2000; Svarc and others, 2002). A 

small amount of uniaxial contraction (–18.0±1.9 nanostrain/

yr, extension reckoned positive) accumulates in the ~N. 60° 

E. direction of relative convergence between the Juan de Fuca 

and North American plates. This rate of strain accumulation 

is equivalent to an average of 3.6 mm/yr of east-northeastdirected contraction across the 200-km-wide zone covered by 

the regional GPS stations and likely represents penetration of 

deformation from the locked subduction thrust into this region. 

The tectonic model does not define possible background vertical deformation, and we assume that there is none because 

regional studies of historical leveling data show little long-term 

elevation change across the Cascades (Verdonck, 2006).

Deformation in the 1980s and 1990s Within a 

Regional High-Precision Trilateration Network

The Helen trilateration network, which straddles the Cascade volcanic arc in southwestern Washington, was established 

to track regional strain accumulation following the 1980 craterforming eruption of Mount St. Helens (fig. 4). Line lengths were 

measured in 1982 and 1991, and they were recovered with GPS 

in 2000 (EDM-measured line lengths can be downloaded from 

http://quake.wr.usgs.gov/research/deformation/gps/geodolite/

index.html). The GPS-recovered line lengths were corrected for 

the –0.283 ppm systematic scale bias relative to Geodolite line 

lengths derived empirically by Savage and others (1996).
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 307

Network Averaging 

interval

e1

(nanostrain/yr)

e2

(nanostrain /yr)

Azimuth e3

(degrees 

clockwise 

from north)

Rotation rate 

(nanoradian/

yr)

Euler Pole 

latitude

(degrees)

Euler Pole 

longitude

(degrees)

Correlation 

latitudelongitude

Standard 

deviation 

unit weight

Helens EDM 1982.7–1991.7 82±28 62±27 44 ± 37 -- -- -- -- --

Helens EDM 1991.7–2000.7 11±29 –10±29 70 ± 35 -- -- -- -- --

Helens GPS 1992–2004 2.5±1.7 –18.0±1.9 63.8 ± 3.5 –15.4±1.2 44.476±0.141 –118.526±0.252 –0.804 1.0

Helens GPS 

(QOCA)

1992–2004 –0.8±2.7 –20.1±2.9 55.0 ± 6.0 –14.3±2.0 45.969±0.31 –117.876±0.540 –0.932 0.84

 125 ̊ 124 ̊ 123 ̊ 122 ̊ 121 ̊ 120 ̊ 119 ̊ 118 ̊ 117 ̊

44 ̊

45 ̊

46 ̊

47 ̊

48 ̊

49 ̊

0 50 100

KILOMETERS

5 MM/YEAR

EULER POLE

OBSERVED

TECTONIC MODEL

OREGON

WASHINGTON

BRITISH COLUMBIA

 NANOSTRAIN/YEAR

ε1 = 2.5 ± 1.7

ε2 = -18.0 ± 1.9

BREW

CHWK

CVO1

DRAO

GOBS

GWEN

JRO1 KELS

LINH

REDM

SEDR

Table 2. EDM and GPS network strain and rotation rates.

[Strain rates within entire network assuming uniform strain accumulation. --, data not calculated.]

Figure 3. Map of GPS station velocities in Oregon, Washington, and British Columbia used to 

define the tectonic model of background rigid-block rotation and uniform strain accumulation. 

Observed station velocity vectors (black or blue arrows tipped with 95-percent-confidence 

error ellipses) are relative to stable North America. Velocity vectors predicted using the model 

are shown in red. Region for 95-percent confidence of the computed Euler pole is indicated by 

dark ellipse in eastern Oregon, rotation rate is indicated by dark wedges, and one-standarddeviation uncertainty in rotation rate is indicated by a light wedge. Large black arrows 

represent computed directions and rates of principal strains. The four most northerly CGPS 

stations (CHWK, DRAO, SEDR, BREW) are not included in the tectonic model calculation. Large 

orange triangles mark locations of major Cascade Range volcanoes.
308 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Line lengths increased consistently by 1 to 3 cm between 

1982 and 1991 but showed no significant change between 

1991 and 2000. The uncertainty in a Helen line length 

measured with a Geodolite is between 4 and 8 mm, with the 

larger value for the longest lines in the network (see Savage 

and Prescott, 1973). By converting the observed line-length 

changes to extension rates and assuming that strain accumulates uniformly over the network, we can map variation in 

the extension rates as a function of line azimuth into surface 

tensor-strain-rate components. Between 1982 and 1991, areal 

dilatation (ε1+ ε2, where ε1 and ε2 are the principal strains 

and extension is reckoned positive) accumulated at a rate 

of 144±39 nanostrain/yr, which is equivalent to a ~3 mm/yr 

(~70 nanostrain/yr) average increase in the length across the 

~40-km-aperture network in any direction. This rate of strain 

accumulation is distinctly different than the background areal 

dilatation rate (at –15±3 nanostrain/yr) estimated with the 

tectonic model (table 2 – note areal dilatation = ε1

+ ε2). No 

significant strain accumulated between 1991 and 2000.

Surficial areal dilatation is expected to accompany 

recharge of a deep spheroidal magma reservoir. Areal dilatation 

 122°40' 122°20' 122°00' 121°40'

46°00'

46°20'

ELK

(ELKK)

VIEW (VMSH)

VANSON (VANS)

MORGLE

(MORG)

BADGER

(BGRM)

MITCHELL

(MITC)

RED

(REDD)

15.7

24.5

20.7 3.7

8.4

19.6

30.1

3.0

26.3

18.3

1.5

8.5

9.5 -8.5

6.2

8.2

-1.4

-11.2

-9.3

-1.9

MOUNT ST.

HELENS

0 5 10

KILOMETERS

15 20

LINE-LENGH CHANGE

1982 to 1991

1991 to 2000

 NANOSTRAIN/YEAR

ε1

 = 82 ± 28, 11 ± 29

ε2 = 62 ± 27, -10± 29

AREA

OF MAP

OREGON

WASHINGTON

Figure 4. Map of the Helen trilateration network. Measured length changes 

and calculated uniform principal strain rates from 1982 to 1991 are shown in 

red. Corresponding values from 1991 to 2000 are shown in blue. EDM station 

names are given, with four-character GPS station name in parentheses. Line 

lengths in 2000 were measured with GPS. Directions and relative magnitudes 

of principal strain rates are shown with large arrows; their values are given 

in nanostrain per year.

is also a sensitive measure of possible systematic scale error in 

the individual EDM surveys because such error is proportional 

to distance measured and is independent of the direction of the 

line (Savage and others, 1986, p. 7471–7472). Systematic error 

in Geodolite line-length measurements is estimated to contribute about 0.14 ppm to the error in a line length, whereas a total 

systematic error of about 0.6 ppm between the two surveys is 

needed to reproduce the observed dilatation. The contribution 

of possible systematic error is included in the uncertainty in 

the rate of areal dilatation, and we conclude that the observed 

areal dilatation between 1982 and 1991 is too large for systematic error alone. Some other phenomenon, like recharge of the 

magma reservoir, contributes to this deformation.

Deformation in the Mount St. Helens GPS 

Network in the Years before September 23, 2004

The 43-station Helens campaign GPS network, surveyed 

in 2000 and 2003, extends for more than 50 km around Mount 
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 309

St. Helens (stations marked with red triangles in fig. 2). Network coverage is increased in space and time by incorporating 

surveys of 13 nearby campaign GPS stations, some of which 

were measured as early as 1992, in the QOCA adjustment for 

average station velocities. Several PANGA CGPS stations, most 

in service since 1997 (table 1), are included in the solutions, and 

one of these, JRO1, is 9 km north of Mount St. Helens.

We use the predictions of the tectonic model given in 

table 2 (network Helens GPS) to remove regional secular 

deformation from the computed average GPS station velocities. We did not propagate uncertainty in the tectonic model 

into error in the corrected station velocities shown in figure 5 

and listed in table 3. At the stations with well-defined velocities—those stations included in the tectonic-model calculation 

(stations marked with yellow triangles in fig. 5)—we see no 

significant motion. A few stations near Mount St. Helens have 

significant velocities, but their movements lack a consistent 

pattern, so we suspect they result from local instabilities or 

survey errors.

HOWE

A545

TMBL

WHEY

STAA

MORT

PACK FL90

SARG

SPR4

DALL

CAST

RKBU

GP06

MARS

REDD

MOWI

LAR5

PRES YALE

KELS

GOBS

GWEN

CVO1

JRO1

MITC

PAN1

ARIL

4100

5040

BGRM

VANS

0 50

KILOMETERS

 123 ̊00' 122 ̊30' 122 ̊00' 121 ̊30' 121 ̊00' 120 ̊30'

45 ̊00'

45 ̊30'

46 ̊00'

46 ̊30'

MOUNT ST.

HELENS INSET

MOUNT

HOOD

MOUNT

ADAMS

COLUMBIA R IVER

VELOCITY SCALE

10 MM/YEAR

AREA

OF

MAP

OREGON

WASHINGTON

Figure 5. Map 

showing average GPS 

station horizontalvelocity vectors (tipped 

with 95-percentconfidence error 

ellipses, scale at 

bottom) for southwest 

Washington and 

northwest Oregon 

through September 

2004, corrected for 

background motion 

predicted by a tectonic 

model. None of stations 

marked with yellow 

triangles, which were 

included in calculations 

for the tectonic model, 

have significant residual 

velocity. CGPS station 

names and velocity 

vectors are shown in 

blue. Most stations 

indicated with red 

triangles were surveyed 

only in the summers of 

2000 and 2003.

Examining the station velocities in the local network 

around Mount St. Helens (fig. 6), we observe no consistent 

pattern of deformation, but roughly half of the velocities 

exceed the 95-percent-confidence error ellipses. The highest velocity is found in the crater at station DMSH on the 

September 1984 lobe of the old dome, which moves to the 

east-northeast and down. This lobe is west of the old dome’s 

center, and the observed motion is toward the center of the 

old dome, consistent with the motion expected as the old 

dome continues to cool and shrink. The westward velocity 

of stations MORG and WGOT, west of Mount St. Helens, is 

greater than 5 mm/yr, but other nearby stations show little or 

no significant motion. With surveys only in 2000 and 2003, it 

is difficult eliminate local instabilities or survey error, such as 

instrument setup, as a cause for the anomalous velocities. The 

daily measurements of position change at CGPS station JRO1 

provide the strongest evidence for the lack of anomalous 

deformation in the years before September 2004. We examine 

the JRO1 data in the next section.
310 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Station Longitude

(degrees)

Latitude

(degrees)

East velocity

 ± 1r error

(mm/yr)

North velocity

 ± 1r error

(mm/yr)

E-N 

correlation

Up velocity

 ± 1r error

(mm/yr)

4100 -122.4706 46.2316 -2.65 ± 1.34 -0.50 ± 1.65 -0.04 -0.49 ± 6.64

5040 -122.5819 46.3695 1.86 ± 1.25 -0.95 ± 1.49 -0.06 13.87 ± 5.70

A545 -120.7435 45.4745 0.27 ± 0.82 1.08 ± 0.92 0.01 -1.76 ± 3.60

APEC -122.1330 46.1950 -1.92 ± 1.06 4.00 ± 1.13 -0.01 -1.77 ± 4.93

ARIL -122.5608 45.9561 1.16 ± 1.18 -2.10 ± 1.35 -0.05 4.29 ± 5.47

BGRM -121.8943 46.2834 -4.32 ± 1.02 2.05 ± 1.12 0.00 0.14 ± 4.66

BIVO -122.1962 46.1453 -2.37 ± 1.20 -4.60 ± 1.41 -0.03 11.24 ± 5.63

CAFE -122.1349 46.2135 -2.56 ± 1.08 -0.39 ± 1.17 -0.03 0.36 ± 5.05

CANN -123.9603 45.8617 4.29 ± 0.72 0.11 ± 0.77 -0.04 -0.94 ± 2.93

CAST -122.9188 46.2833 -0.09 ± 0.85 -0.11 ± 0.92 -0.08 3.42 ± 3.30

CVO1 -122.4961 45.6109 -10.76 ± 6.44 1.92 ± 7.12 0.00 17.10 ± 29.68

DALL -121.1752 45.6178 0.18 ± 0.67 0.76 ± 0.75 0.00 -3.40 ± 2.81

DMSH -122.1910 46.2000 27.69 ± 1.04 9.39 ± 1.15 -0.01 -90.29 ± 4.86

DRAO -119.6250 49.3226 -1.78 ± 0.43 1.46 ± 0.44 0.00 1.25 ± 1.39

EADM -122.1521 46.1973 -1.18 ± 1.04 0.90 ± 1.12 -0.02 -3.54 ± 4.84

ELKK -122.3452 46.3009 -0.65 ± 1.09 0.94 ± 1.18 -0.02 -0.24 ± 4.95

FL90 -121.1333 46.6288 -0.14 ± 0.64 0.36 ± 0.74 0.05 -6.53 ± 2.90

GOBS -120.8147 45.8388 0.69 ± 0.59 0.61 ± 0.61 -0.02 0.33 ± 1.68

GP06 -122.5472 45.7810 1.76 ± 1.32 -1.80 ± 1.21 0.06 -5.07 ± 4.49

GWEN -121.3276 45.7826 1.28 ± 0.52 -0.53 ± 0.54 -0.02 0.51 ± 1.58

HOWE -122.8797 44.9900 3.50 ± 1.18 -0.29 ± 1.25 -0.07 -4.69 ± 4.57

HPAD -122.0315 46.2520 -4.97 ± 1.03 1.09 ± 1.16 0.00 2.94 ± 4.70

JARD -122.2973 46.2469 -1.89 ± 1.10 0.71 ± 1.19 -0.01 -28.35 ± 5.13

JRO1 -122.2176 46.2751 -0.16 ± 0.78 1.03 ± 0.83 -0.01 -2.93 ± 3.53

KELS -122.8961 46.1182 -0.01 ± 0.55 0.02 ± 0.58 -0.02 -0.05 ± 1.64

LAR5 -122.0876 45.5327 -1.95 ± 0.90 -4.10 ± 0.98 -0.01 -1.77 ± 3.96

LOOW -122.1843 46.2248 4.03 ± 1.05 0.77 ± 1.13 -0.01 1.29 ± 4.88

LVCY -122.0912 46.1638 3.52 ± 3.69 2.02 ± 2.82 0.11 2.91 ± 8.44

MARS -121.5419 46.0001 0.66 ± 0.64 0.28 ± 0.69 -0.02 1.29 ± 2.72

MITC -122.1931 46.0309 -2.05 ± 1.05 -3.21 ± 1.18 -0.01 4.56 ± 4.73

MORG -122.3572 46.1893 -7.42 ± 1.00 2.70 ± 1.11 -0.01 -0.66 ± 4.61

MORT -122.2704 46.5506 0.80 ± 0.58 0.56 ± 0.64 -0.01 -1.76 ± 2.51

MOWI -122.0308 45.7998 -4.45 ± 2.56 -1.17 ± 2.06 -0.12 -0.15 ± 6.61

NEBU -122.2173 46.1782 -3.56 ± 1.08 2.41 ± 1.17 0.01 3.49 ± 5.02

NELR -122.1649 46.1993 -1.33 ± 1.06 -1.07 ± 1.13 -0.03 1.99 ± 4.88

NESK -123.9662 45.1343 5.07 ± 0.79 1.86 ± 0.83 -0.05 -1.51 ± 3.10

NMAR -122.1280 46.1112 -1.23 ± 1.06 0.81 ± 1.20 0.03 5.58 ± 4.87

Table 3. Adjusted Helens campaign GPS network average station velocities through September 23, 2004.

[Velocities are in nominal ITRF2000 fixed North American Plate terrestrial reference frame and are corrected for regional rigid-body 

rotation and strain.]
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 311

Station Longitude

(degrees)

Latitude

(degrees)

East velocity

 ± 1r error

(mm/yr)

North velocity

 ± 1r error

(mm/yr)

E-N 

correlation

Up velocity

 ± 1r error

(mm/yr)

NWDO -122.2037 46.2096 2.72 ± 1.04 0.10 ± 1.13 -0.01 0.95 ± 4.79

PACK -121.6752 46.6068 0.23 ± 0.64 0.80 ± 0.69 0.00 -5.60 ± 2.71

PAN1 -121.7685 46.3078 -0.02 ± 1.05 -1.04 ± 1.14 0.00 0.22 ± 4.84

PRES -122.9245 46.0401 1.54 ± 0.78 2.06 ± 0.86 -0.04 -4.45 ± 3.23

R100 -122.1538 46.2322 1.16 ± 1.01 -1.01 ± 1.12 0.00 -0.64 ± 4.60

REDD -121.8210 45.9349 -2.01 ± 1.53 -5.29 ± 1.38 -0.04 15.11 ± 5.71

RKBU -122.5660 45.5467 0.49 ± 0.72 1.41 ± 0.78 -0.03 -5.58 ± 2.96

RS26 -123.4611 45.7968 3.37 ± 0.84 1.31 ± 0.94 -0.04 -2.90 ± 3.46

SARG -120.4772 46.6024 -0.04 ± 0.71 0.53 ± 0.77 -0.01 -1.67 ± 2.83

SERI -122.1765 46.1877 -2.81 ± 1.10 -4.41 ± 1.22 -0.05 -5.52 ± 5.07

SESL -122.1615 46.1737 -1.37 ± 1.06 1.71 ± 1.14 -0.02 0.40 ± 4.94

SMCR -122.1117 46.2586 -3.83 ± 0.96 2.37 ± 1.05 0.00 0.78 ± 4.33

SOFO -122.2329 46.2153 0.83 ± 1.09 5.43 ± 1.18 -0.01 -1.88 ± 5.04

SPR4 -121.1371 46.2269 0.46 ± 0.59 -0.11 ± 0.63 -0.02 0.52 ± 2.47

SRID -122.1899 46.1801 -1.65 ± 1.03 0.36 ± 1.12 -0.01 -5.21 ± 4.78

STAA -122.9853 46.6708 0.15 ± 0.55 -0.72 ± 0.58 -0.03 1.60 ± 2.33

STUD -122.2269 46.2314 2.99 ± 1.08 1.89 ± 1.18 0.00 -0.12 ± 5.03

TMBL -121.7119 45.3340 0.35 ± 0.81 0.98 ± 0.80 -0.07 0.85 ± 3.22

TUBE -122.3547 46.1095 2.22 ± 1.25 3.59 ± 1.41 -0.06 -2.24 ± 5.88

VANS -122.1516 46.4101 3.54 ± 1.05 1.75 ± 1.16 0.01 0.47 ± 4.77

VMSH -122.1820 46.2998 -1.54 ± 1.04 3.07 ± 1.12 -0.01 2.91 ± 4.81

WGOT -122.3106 46.1730 -8.49 ± 0.88 0.80 ± 0.97 -0.01 -4.27 ± 3.94

WHEY -122.0349 45.3762 -0.43 ± 0.83 2.02 ± 0.96 0.03 -4.67 ± 3.69

WIWE -122.1587 46.2129 1.20 ± 1.05 -4.18 ± 1.13 0.00 2.21 ± 4.85

WRAD -122.2035 46.1897 0.25 ± 1.07 7.46 ± 1.18 0.02 3.44 ± 5.00

WRID -122.2107 46.1978 1.05 ± 1.05 -2.31 ± 1.13 0.00 0.23 ± 4.87

YALE -122.3181 46.0262 0.26 ± 0.64 -0.47 ± 0.69 -0.07 1.18 ± 2.61

Table 3. Adjusted Helens campaign GPS network average station velocities through September 23, 2004. —Continued

[Velocities are in nominal ITRF2000 fixed North American Plate terrestrial reference frame and are corrected for regional rigid-body rotation 

and strain.]
312 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Time Series of Deformation from CGPS Data 

Collected at Johnston Ridge Observatory

A detailed history of deformation near Mount St. Helens 

since 1997 is found in the daily displacement components of 

CGPS station JRO1. The Cascadia network includes daily 

solutions in the fixed-NOAMgp reference frame and in a 

local reference frame that is defined by a subset of 11 CGPS 

stations in the network (table 4). QOCA’s time-series analysis 

is used to simultaneously compute the component velocities, 

offsets, and repetitive seasonal displacements of JRO1. We do 

not include the anomalous data after September 23, 2004, in 

this analysis for background deformation. Outliers in the time 

series, found mainly in the winter when ice and snow accumulate around the GPS antenna, are defined as observations 

where the residual east or north displacement is larger than 5 

mm, and the residual vertical displacement is larger than 15 

mm. These residual outliers are about three times the weighted 

root-mean-square (WRMS) scatter observed in the JRO1 time-

 122 ̊20' 122 ̊15' 122 ̊10' 122 ̊05' 122 ̊00'

46 ̊05'

46 ̊10'

46 ̊15'

46 ̊20'

0 5 10

KILOMETERS

JRO1

NMAR

LVCY

BIVO

MORG

TUBE

ELKK VMSH

JARD SMCR

HPAD

WGOT

R100

LOOW

SOFO

STUD

CAFE

APEC

SESL NEBU

WRID

NWDO WIWE

EADM

NELR

WRAD SERI

SRID

DMSH

AREA

OF MAP

WASHINGTON

VELOCITY SCALE

10 MM/YEAR

Figure 6. Map showing 2000–2003 corrected average GPS station horizontalvelocity vectors (tipped with 95-percent-confidence error ellipses, scale at bottom) 

for Mount St. Helens area. Although many stations show significant velocities, 

there is no systematic pattern that might suggest premonitory deformation before 

the start of the 2004 eruption. GPS station DMSH, on the 1980–86 lava dome, moved 

downward 87 mm/yr and east-northeastward 35 mm/yr, presumably as a result of 

cooling and compaction of the dome.

series analysis (table 4). Of the 2,267 daily observations, 130 

(6 percent) were removed from the analysis because one or 

more of the components was identified as an outlier.

Small vertical offsets (~13 mm) in the time series 

were introduced on July 8, 1998, when a UNAVCO radome 

replaced the original Snow radome, and on July 16, 2002, 

when a SCIGN radome replaced the UNAVCO radome. 

The vertical offsets are calculated in QOCA by first setting 

assumed offset values based on estimates from a short span of 

data around the offset (we use 0.2 year), which is then given a 

weight of 0.25 times the offset; the final values are estimated 

simultaneously with other parameters. In our final analysis, 

only vertical offsets are estimated because initial analysis 

resulted in horizontal offsets that were small and insignificant. 

Nevertheless, the long-term trends could be biased slightly by 

error in the applied vertical offsets and possible small, horizontal offsets related to the antenna-radome changes.

The estimated JRO1 component velocities, offsets, and 

repetitive seasonal movements are removed to produce a resid-
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 313

Station JRO1 KELS KELS BREW CVO1 GOBS REDM

Start date (year month day) 19970523 19971010 20040923 19971010 20030104 19970827 19980725

End date (year month day) 20040922 20040922 20061013 20061013 20061013 20061013 20061013

Number of daily positions (post-1996) 2,137 2,977 -- 1,669 1,308 3,149 2,889

Number of outliers 130 201 -- 1 4 30 31

East WRMS (mm) 1.70 1.62 -- 0.82 1.13 1.13 1.32

North WRMS (mm) 1.75 1.77 -- 0.91 1.19 1.23 1.12

Up WRMS (mm) 5.01 4.62 -- 3.23 3.50 3.36 3.25

East velocity (mm/yr) 2.76 3.07 6.14 1.33 3.26 1.52 0.16

North velocity (mm/yr) 5.82 5.84 5.03 0.75 5.63 2.09 2.80

Vertical velocity (mm/yr) -1.99 -0.45 0.11 0.18 -1.54 -0.73 0.02

East annual amplitude (mm) 2.02 0.84 -- 0.35 0.51 1.83 0.35

East annual phase (degrees) 52.35 37.32 -- 239.69 288.89 139.06 351.70

North annual amplitude (mm) 1.96 0.98 -- 0.10 0.06 1.30 0.62

North annual phase (degrees) 137.86 160.71 -- 39.96 161.07 326.93 144.47

Vertical annual amplitude (mm) 1.70 1.10 -- 0.35 1.66 0.49 0.97

Vertical annual phase (degrees) 25.41 354.79 -- 284.29 356.08 277.76 50.14

East semiannual amplitude (mm) 0.34 0.34 -- 0.06 0.12 0.25 0.20

East semiannual phase (degrees) 244.99 109.28 -- 316.72 176.88 267.05 220.03

North semiannual amplitude (mm) 0.80 0.22 -- 0.11 0.19 0.30 0.27

North semiannual phase (degrees) 263.56 236.34 -- 199.17 87.40 85.48 330.50

Vertical semiannual amplitude (mm) 0.64 0.31 -- 0.01 0.84 0.40 0.19

Vertical semiannual phase (degrees) 276.03 122.27 -- 22.97 84.91 140.26 145.93

Station LIND LINH GWEN DRAO CHWK SEDR

Start date (year month day) 19980728 20020506 19970827 19970827 19981119 19971007

End date (year month day) 20061013 20061013 20061013 20061013 20061013 20061013

Number of daily positions (post 1996) 1,145 1,531 3,138 3,255 2,543 3,209

Number of outliers 13 16 72 12 11 7

East WRMS (mm) 1.3 1.11 1.55 1.00 1.10 1.14

North WRMS (mm) 1.19 1.15 1.50 1.16 1.42 1.33

Up WRMS (mm) 3.95 3.90 4.32 2.97 4.24 3.55

East velocity (mm/yr) 1.70 1.70 2.10 0.93 2.43 2.97

North velocity (mm/yr) 1.69 1.63 2.56 0.41 1.60 2.11

Vertical velocity (mm/yr) -0.53 0.01 -0.85 1.61 0.25 -0.23

East annual amplitude (mm) 0.26 0.33 0.87 0.42 0.44 0.46

East annual phase (degrees) 90.92 75.29 278.58 73.03 322.68 305.77

North annual amplitude (mm) 0.81 0.48 0.39 0.24 0.37 0.23

North annual phase (degrees) 218.14 209.76 138.47 258.20 85.11 204.89

Vertical annual amplitude (mm) 0.47 1.64 1.69 0.71 0.47 0.09

Vertical annual phase (degrees) 83.08 25.46 229.45 98.87 160.64 279.40

East semiannual amplitude (mm) 0.11 0.37 0.15 0.12 0.02 0.28

East semiannual phase (degrees) 137.32 45.34 349.51 250.97 149.82 77.91

North semiannual amplitude (mm) 0.05 0.25 0.26 0.03 0.17 0.11

North semiannual phase (degrees) 306.59 269.95 232.06 278.87 168.30 30.14

Vertical semiannual amplitude (mm) 0.39 1.33 0.62 0.13 0.61 0.05

Vertical semiannual phase (degrees) 115.22 186.33 357.05 18.87 252.14 117.52

Table 4. Time-series analysis of Cascadia network local reference frame stations.

[WRMS is weighted root-mean-squared deviation (experimental standard deviation). Phases are relative to start of the calendar year. Data for station 

KELS are divided into two intervals to calculate velocities; other KELS data, equivalent between columns, are shown by dashes in second KELS column.]
314 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ual-component time series. The horizontal components are 

then rotated so that one component represents motion radial 

(N. 13.5o W.), and the other component represents tangential 

motion (N. 76.5o E.) to Mount St. Helens (fig. 7A). Previously 

identified outliers are included in these plots, but they will be 

outside the 5-mm (horizontal) and 15-mm (vertical) datarejection criteria indicated by the yellow bands on figure 7.

Station JRO1 moved at the background rate (no slope 

in fig. 7A) until the start of unrest on September 23, 2004, 

when anomalous subsidence and movement toward Mount St. 

Helens began at the station. Intervals of low-level systematic 

1998 2000 2002 2004 2006

YEAR

- 20

0

20

40

60

DETRENDED DISPLACEMENT. IN MILLIMETERS

1998 2000 2002 2004 2006

YEAR

0

20

40

60

80

TANGENTIAL (N10°W)

RADIAL (S80°W)

VERTICAL

DETRENDED DISPLACEMENT. IN MILLIMETERS

JRO1

A

B

KELS

TANGENTIAL (N76.5°E)

RADIAL (N13.5°W)

VERTICAL

Figure 7. Stacked plots of 

detrended tangential, radial, 

and vertical daily average 

displacement of CGPS 

stations (A) JRO1, located 

9 km north-northwest of 

Mount St. Helens, and (B) 

KELS, located 50 km westsouthwest. Displacements 

are residuals from fitting 

a station velocity, mean 

value, cyclical annual 

and semiannual drift, and 

offsets from GPS antenna 

changes as determined 

by data collected through 

September 23, 2004. Vertical 

blue lines mark times that 

GPS antenna radomes 

were changed; vertical red 

line marks start of seismic 

unrest on September 23, 

2004. Yellow horizontal lines 

are ±5 mm from pre-event 

tangential and radial mean 

values (outlier rejection 

limit) and ±15 mm from preevent vertical mean value 

(outlier rejection limit). For 

station JRO1, one or more 

displacements were rejected 

for 130 of a total 2,267 days 

of data before September 

23, 2004. For station KELS, 

one or more displacement 

components were rejected 

for 201 of 2,927 days of data.

noise remain in the residual-component time series, as do short 

periods of drift when winter snow and ice accumulate on the 

GPS antenna and distort its phase center.

The average velocity of station JRO1 derived from the 

time-series analysis is similar to the velocity estimated from 

the local reference-frame calculation, but it is 0.9 mm/yr faster 

in the north component and 0.5 mm/yr slower in the east 

component than the velocity predicted by the tectonic model. 

We later use the tectonic model to correct post-September 23, 

2004, velocities at JRO1 and at other stations around Mount 

St. Helens having unknown secular velocities. The observed 
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 315

discrepancy between the tectonic model and the observed 

motion at JRO1 is considered in the discussion of eruptionrelated deformation source models.

A small inward displacement of KELS, one of the 

PANGA CGPS stations used as a reference station in the local 

frame, also appears to have initiated at the start of volcanic 

unrest (fig. 7B). Station KELS is in Kelso, Wash., about 50 km 

west-southwest of Mount St. Helens, and the GPS antenna is 

mounted on a stable, deep-drilled, braced monument (Langbein and others, 1995) that is anchored at a depth of 10 m and 

isolated from the topsoil. The time-series analysis at KELS 

includes an additional term to resolve average velocities for 

the periods before and after September 23, 2004 (table 4). 

If deformation from the eruption of Mount St. Helens penetrated as far as station KELS, then it likely is the result of 

magma chamber pressure changes in a relatively deep (~10 

km) source. We later examine data from several PBO CGPS 

stations located an equivalent distance to the north of Mount 

St. Helens to determine whether eruption-related deformation 

could be detected there.

2005 2005.5 2006 2006.5

-20

-10

0

10

20

30

40

YEAR

DETRENDED DISPLACEMENT, IN MILLIMETERS

TANGENTIAL (N76.5°E)

RADIAL (N13.5°W)

VERTICAL

2004.74 2004.76 2004.78

20

24

28

32

Figure 8. Stacked daily tangential (N. 76.5° E., top), radial (N. 13.5° W., middle), and vertical (bottom) 

displacements of CGPS station JRO1 from shortly before and during the 2004–2006 dome-building eruption 

of Mount St. Helens. Error bars are ±1 standard deviation. Bold horizontal black line represents projection 

of long-term trend in radial displacement, red vertical line marks start of unrest (seismic swarm beginning 

September 23, 2004), and orange vertical lines show start and end times of a series of phreatic explosions 

between October 1 and 5, 2004, and the time of last explosion on March 8, 2005. Vertical displacement 

appears to have ceased shortly after the March 8, 2005, explosion, whereas radial displacement has decayed 

more slowly. Inset shows a 23-day period at the start of unrest. Winter months show periods of anomalous 

displacements because of snow and ice accumulation on the GPS antenna radome.

GPS Station Displacements Early in the 2004–

2006 Eruption

The only direct measure of deformation associated 

with the initiation of the eruption was made at JRO1, where 

slow, relatively steady motion toward Mount St. Helens was 

observed (fig. 8). The movement of station JRO1 was fastest in 

the two weeks after September 23, 2004, but even then it averaged only about 0.5 mm/day (inset, fig. 8). The inward movement continued at a nearly linear rate until a few days after the 

vigorous steam-and-ash explosion on October 5, 2004, when 

movement began to slow at a time-decaying rate. We later 

conclude that deformation has gradually slowed at other CGPS 

stations around Mount St. Helens.

The conduit that fed the eruptions in the 1980s is buried 

under the dome that developed in the years after the crater 

formed and that presumably was plugged with cooled magma. 

Was this plug pushed out at the start of the new domebuilding eruption or was it bypassed? If a bypass formed, 

most likely it would appear as a dike that propagated to the 
316 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

surface, and we would expect associated deformation to be 

fairly widespread. The propagation of a shallow dike might 

not produce much deformation at station JRO1, located 9 km 

from the center of the crater, and we examine data from other 

stations closer to the newly formed vent for possible offsets at 

the start of the eruption.

An indirect measure of deformation associated with 

initial movement of magma to the surface is obtained at the 

subset of campaign GPS stations that were occupied within a 

week or two after the start of unrest. Displacement offset of 

these stations is estimated by projecting the observed 2000 to 

2003 position-change trends to the start of unrest (September 

23, 2004). We include data from an L1 CGPS station, DOM1, 

located on the September 1984 lobe of the old dome (see 

LaHusen and others, this volume, chap. 16). A power failure 

occurred at DOM1 in January 2004, and the station was not 

reactivated until September 27, 2004.

The large northward and upward offset at station DOM1 

(in the same location as station DMSH) contrasts with the 

smaller and seemingly random pattern of offsets at surrounding stations (horizontal-displacement offset are shown in 

fig. 9, and component offsets are listed in table 5). Many of 

these offsets, though statistically significant on the basis of 

assumed error estimates, likely result from systematic error 

P690

P699

P702

 122 ̊20' 122 ̊15' 122 ̊10' 122 ̊05'

46 ̊05'

46 ̊10'

46 ̊15'

46 ̊20'

0 5 10

KILOMETERS

JRO1

LVCY

BIVO

TUBE (P687)

VMSH

SMCR

WGOT

LOOW R100

STUD (TSTU)

NEBU

WRID

(TWRI)

NWDO

(P693) WIWE (TWIW)

EADM (P696)

NELR (P695)

SERI (P697)

DMSH

P698

HORIZONTAL

DISPLACEMENT SCALE

50 MM

AREA

OF MAP

WASHINGTON

in the projected trends. Anomalous 2000 to 2003 velocities (fig. 6) were observed at many of the stations having 

relatively large event offsets (for example, stations WGOT, 

BIVO, and SERI, fig.9). These offsets are opposite of the 

anomalous trends, as might be expected if there were errors 

in the estimated station velocities. Event deformation appears 

to be limited to DOM1, which was about 300 m from the 

newly formed vent. The partially congealed magma plugging the conduit appears to have been pushed to the surface 

on the south side of the old dome, displacing the old dome 

slightly to the north and up. Most of the movement at station 

DOM1 had occurred by the time the station was reactivated 

(September 27, 2004). In its final days, the station continued moving north until it was destroyed by the first phreatic 

explosion on October 1, 2004 (LaHusen and others, this 

volume, chap. 16).

Displacement of GPS Stations During the 2004–

2006 Eruption

The inward displacement of CGPS station JRO1 appears 

to start at about the same time as the first seismic swarm on 

Figure 9. Map showing early 

event GPS station horizontaldisplacement vectors (tipped 

with 95-percent-confidence 

error ellipses, vector scale 

on bottom) in the Mount St. 

Helens area. Displacements 

calculated by projecting station 

velocities to September 23, 

2004, and estimating the offset 

between pre- and post-event 

positions. Many stations show 

significant movement, but 

there is no systematic pattern 

to suggest intrusion of a new 

dike to transport magma to the 

surface. Large displacement on 

the 1980–86 lava dome is that 

determined from measurements 

at DOM1 (same location as 

DMSH), an L1 CGPS station that 

was resurrected on September 

27, 2004 (LaHusen and other, this 

volume, chap. 16).
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 317

Station

Measurement 

date

(2004)

Longitude

(degrees)

Latitude

(degrees)

East offset

(mm)

North offset 

(mm)

East-North 

correlation

Vertical offset

(mm)

JRO1 09/28 -122.218 46.275 -1 ± 1.8 -3 ± 1.9 0.004 6 ± 8.7

DOM1 09/28 -122.191 46.200 -16 ± 15.3 235 ± 15.7 0.014 90 ± 40.1

EADM 10/01-10/07 -122.152 46.197 3 ± 2.6 -7 ± 2.9 -0.007 -4 ± 12.2

LOOW 09/28-10/07 -122.184 46.225 -12 ± 2.5 -5 ± 2.8 0.012 -13 ± 11.8

NELR 09/29-10/07 -122.165 46.199 9 ± 2.7 7 ± 2.9 -0.038 3 ± 12.6

NEBU 09/28-10/07 -122.217 46.178 10 ± 2.5 -6 ± 2.8 0.043 -14 ± 12.0

NWDO 09/28-10/07 -122.204 46.210 -13 ± 2.5 9 ± 2.9 -0.002 -3 ± 12.0

SERI 09/30-10/05 -122.177 46.188 17 ± 3.0 20 ± 3.4 -0.070 24 ± 13.7

STUD 09/30-10/07 -122.227 46.231 -15 ± 2.6 -7 ± 2.9 0.016 0 ± 12.1

WIWE 09/29-10/07 -122.159 46.213 -9 ± 2.5 19 ± 2.8 0.021 -7 ± 11.8

WRID 10/01-10/05 -122.211 46.198 -11 ± 2.7 15 ± 3.0 0.024 2 ± 12.9

BIVO 11/03-11/09 -122.196 46.145 10 ± 3.4 27 ± 3.9 0.019 -41 ± 16.1

LVCY 11/04-11/09 -122.091 46.164 -22 ± 5.4 2 ± 4.6 0.057 -3 ± 17.1

R100 11/06 and 11/09 -122.154 46.232 -14 ± 3.0 8 ± 3.4 0.020 -3 ± 14.2

SMCR 11/06 and 11/09 -122.112 46.259 5 ± 3.2 -8 ± 3.6 0.010 -4 ± 14.7

TUBE 11/03-11/07 -122.355 46.110 -9 ± 3.0 -3 ± 3.4 -0.033 -13 ± 13.9

VMSH 11/04-11/09 -122.182 46.300 1 ± 2.8 -9 ± 3.2 0.009 -6 ± 13.3

WGOT 11/04 and 11/05 -122.311 46.173 37 ± 3.3 2 ± 3.7 0.014 -8 ± 15.1

Table 5. Estimated GPS station offsets at start of 2004 seismic unrest.

[Offsets are estimated by projecting 2000–2003 trends to September 23, 2004. Uncertainties are ±1 standard deviation.]

September 23, 2004, and the displacement has continued at 

a decreasing rate throughout the eruption (fig. 8). The rate of 

subsidence, however, appears to have leveled off shortly after 

the March 8, 2005, explosion. Several short periods of anomalous vertical offsets in the winter of 2005, probably the result 

of snow and ice accumulation on the GPS antenna, make the 

subsidence appear jerky during the first few months of the eruption. The horizontal components of position change at JRO1 are 

affected less by winter anomalies, but outliers and short-term 

variations are apparent in the time series. At the measurement 

uncertainty of about ±1.7 mm in the horizontal components 

and ±5 mm in the vertical component of displacement (table 4), 

there is no apparent deformation anomaly in the months or days 

before the start of unrest.

The pattern and rate of deformation around Mount St. 

Helens during all but the first few weeks of the eruption are 

measured by a network of CGPS stations. Eleven new CGPS 

stations (seven from PBO and four from CVO) were installed 

by November 5, 2004 (see table 1 for installation dates and 

additional information). The PBO stations are part of a planned 

network of 20 CGPS stations (four with collocated borehole 

tiltmeters) and four strainmeter stations (each consisting of a 

borehole tensor strainmeter, a three-component seismometer, 

and a tiltmeter) in the NSF-funded PBO Mount St. Helens 

volcanic cluster. The data from new CGPS stations, along with 

data from JRO1, are processed by CVO as part of the HelensMonit network. Network solutions are transformed into fixedNOAMgp, ITRF2000, and a local reference frame. The local 

reference frame for the HelensMonit network consists of five 

reference stations (GOBS, REDM, KELS, LINH, CVO1).

We use QOCA’s time-series analysis to estimate and 

remove cyclical annual and semiannual seasonal noise from 

the daily CGPS solutions. We then remove secular tectonic 

motion predicted by the tectonic model and rotate the resulting 

cleaned, horizontal-component time series from the east and 

north geographic components into components that are radial 

and tangential to Mount St. Helens. At several stations, an 

eccentric offset correction (table 6) is used to transfer the positions measured at the campaign GPS station early in the eruption to the nearby CGPS station. Then, horizontal and vertical 

time-series data are stacked to compare position change over 

time (figs. 10, 11).

The CGPS stations near Mount St. Helens are divided into 

high- and low-altitude groups for the stacked-component timeseries plots. The low-altitude stations have minor winter displacement anomalies, whereas stations at altitudes greater than 2,000 

m experience frequent, and sometimes large, winter excursions. 

The locations of the CGPS stations are shown in figure 9.
318 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

B

2005 2005.5 2006 2006.5

 YEAR

20

40

60

80

100

P687 17.0 KM

P702 16.7 KM

JRO1

8.5 KM

TSTU

5.0 KM

TGAU

2.0 KM

TWIW

2.0 KM

P698

3.7 KM

P696

2.3 KM

2005 2005.5 2006 2006.5

20

40

60

80

100

P687

P702

JRO1

TSTU

TGAU

TWIW

P698

P696

RADIAL DISPLACEMENT, IN MILLIMETERS TANGENTIAL DISPLACEMENT, IN MILLIMETERS

YEAR

A

Figure 10. Plots of detrended and stacked time series of daily radial (A), tangential (B ), and 

vertical (C) displacements at eight CGPS stations around Mount St. Helens at altitudes less 

than 2,000 m. Locations of stations shown in figure 9. Error bars are ±1 standard deviation, and 

gray bands in radial and vertical displacements are best-fitting rates of exponential decay. The 

red vertical line marks the start of unrest on September 23, 2004, and the orange vertical lines 

mark three phreatic explosions. A few of the earliest measurements at TSTU, TWIW, and P696 

were made at a nearby campaign GPS station, and an eccentric correction was used to reduce 

positions to corresponding CGPS station. Tangential displacements at TSTU, TWIW, and TGAU 

are significantly different than zero. Station distances in A and C are computed relative to surface 

projection of modeled source (red cross in fig. 17A).
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 319

2005 2005. 5 2006 2006. 5

50

100

150

200

250

VERTICAL DISPLACEMENT, IN MILLIMETERS

YEAR

P687 

17.0 KM

P702

16.7 KM

JRO1

8.5 KM

TSTU

5.0 KM

TGAU

2.0 KM

TWIW

2.0 KM

P698

3.7 KM

P696

2.3 KM

C

Figure 10—Continued.

The maximum radially inward movement of the lowaltitude stations (fig. 10A) occurs at stations TSTU and JRO1, 

located at distances of 5 and 9 km from Mount St. Helens, 

respectively. Radial movement diminishes at stations located 

farther (P687, P702) and closer (TGAU, TWIW, P698, P696) 

to Mount St. Helens. The tangential component of motion is 

small, except at stations TSTU, TGAU, and TWIW (fig. 10B). 

Subsidence occurs at all stations and reaches its maximum at 

TSTU, but it diminishes at stations closest to and farthest from 

Mount St. Helens (fig. 10C).

During winter months, CGPS stations installed at 

altitudes above 2,000 m on Mount St. Helens (P695, P690, 

TWRI, P693, P697, P699) experience periods of temporary 

pseudo-displacements (fig. 11) from the accumulation of rime 

on the GPS antenna (fig. 12). At two high-altitude stations 

(P690 and P699) where the solar panels were mounted on the 

south side of the GPS antenna mast, asymmetrical ice loading 

permanently warped the towers.

The 2004–2005 winter displacements at station P697 

were unusually large and constant (fig. 11). Initially we 

hypothesized that the apparent displacement of this station 

normal to the south crater wall resulted from the impact of 

the growing lava spine on that wall. With warming weather 

in the spring of 2005, the ice covering the GPS antenna 

melted, and the station returned to its initial position. The 

winter of 2004–2005 was unusually mild, and large, systematic pseudodisplacements were observed only at the highest 

stations (P699 and P697), and only the tower for station P699 

was warped permanently by ice loading. During the more 

typical winter of 2005–2006, ice coated the GPS antennas at 

all high-altitude stations, and the towers at stations P699 and 

P690 were warped by ice loads on the attached solar panels.

Wintertime pseudodisplacements at the high-altitude 

CGPS stations complicate analysis of eruption-related deformation there. The application of suitable outlier-rejection 

criteria in the QOCA time-series analysis minimizes the need 

for hand-editing data outliers. The calculation of tower offsets 

at stations P699 (2005 and 2006) and P690 (2006), however, 

requires careful data editing, and even then the calculated 

horizontal offsets are poorly determined. The warping of the 

3-m-high towers did not appear to cause significant vertical 

offset. A mechanical offset was introduced at station P699 in 

September 2006 during site maintenance when a UNAVCO 

field engineer removed the radome cover and reset the tilted 

GPS antenna. At the same time, the pole-mounted solar panels 

at P690 and P699 were removed and replaced with panels 

mounted on separate frames in an effort to prevent future loadinduced offsets.

Time-Decaying Movement of GPS Stations 

During the 2004–2006 Eruption

The radial and vertical CGPS station displacements have 

been slowing over time, and the wide, gray bands shown in the 

stacked radial and vertical displacement plots are best-fitting 

curves of the form

 , (1) 

where a and b are constants, t is the observed time, and t

0

 is 

the start time. In the data fit we use October 15, 2004, as the 

start time and include only data after that date. Our main inter-
320 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

2005 2005. 5 2006 2006. 5

25

50

75

100

125

150

175

200

P695

1.4 KM

P690

2.7 KM

TWRI

2.5 KM

P693

1.8 KM

P697

1.7 KM

P699

2.1 KM

2005 2005.5 2006 2006.5

25

50

75

100

125

150

175

200

P695

P690

TWRI

P693

P697

P699

RADIAL DISPLACEMENT, IN MILLIMETERS TANGENTIAL DISPLACEMENT, IN MILLIMETERS

YEAR

YEAR

B

A

Figure 11. Plots of detrended and stacked time series of daily radial (A), tangential (B ), and vertical (C ) 

displacements at six CGPS stations around Mount St. Helens at altitudes greater than 2,000 m. Locations 

of stations shown in figure 9. Error bars are ±1 standard deviation, and gray bands in radial and vertical 

displacements are best-fitting rates of exponential decay. Gray dots are outliers (pseudodisplacements) 

that result from ice coating the GPS antennas, and they are not used in analysis. Short red vertical lines 

mark time used to calculate an offset for stations P690 and P699 to account for permanent warping of 

the tower from ice loading. The red vertical line marks the start of unrest on September 23, 2004, and the 

orange vertical lines mark three phreatic explosions. A few of earliest measurements at P695 and P693 

were made at a nearby campaign GPS station, and an eccentric offset correction was used to reduce 

positions to corresponding CGPS station.
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 321

2005 2005.5 2006 2006.5

50

100

150

200

250

300

P695

1.4 KM

P690

2.7 KM

TWRI

2.5 KM

P693

1.8 KM

P697

1.7 KM

P699

2.1 KM

VERTICAL DISPLACEMENT, IN MILLIMETERS

YEAR

C

Bedrock station CGPS station Measurement date X offset

(m)

Y offset

(m)

Z offset

(m)

WIWE TWIW 2004/10/12 -4.2706 -0.2173 2.2044

WRID TWRI 2005/08/11-08/18 -71.1883 71.0514 -10.0691

SERI P697 2005/08/11-08/18 -11.1530 -2.2178 -0.8293

NWDO P693 2004/10/15-10/19 121.0864 -5.9706 49.7039

STUD TSTU 2004/10/15 394.1229 211.2114 465.2217

NELR P695 2004/10/15 36.6088 -37.1623 -30.4480

TUBE P687 2004/11/03-11/07 1.9623 1.2851 4.3265

SESL P698 2005/07/13-07/22 50.0780 -53.8675 -15.3364

ELKK P702 2004/11/05-11/07 -51.9197 -23.0480 -64.4552

EADM P696 2004/10/15 10.7039 -46.8483 -36.9437

WRAD P699 2005/08/19 18.1105 -12.8303 19.8468

Figure 11—Continued.

Table 6. Surveyed offsets between bedrock campaign GPS stations and nearby tower or braced CGPS 

stations.

[Offsets may have been surveyed repeatedly and are given in Cartesian coordinates. CGPS, continuously transmitting GPS.]

est is in b, which gives the rate of time decay in the displacement 

components. We add a constant to the residual-component timeseries data to make all residual components positive; thus intercept a is arbitrary. As discussed previously, the radially inward 

displacement of JRO1 between September 23 and approximately 

October 8 averaged 0.5 mm/day (inset, fig. 8). Equation 1 is 

one of several similar curves derived by Mastin and others (this 

volume, chap. 22) to fit the volume growth of the dome over 

time. Changes in the rate of dome growth might be expected 

to be matched by changes in the rate of volume or mass loss in 

the magma chamber feeding the eruption. We later test whether 

the observed deformation can be explained by depressurization 

of a magma chamber, but first we look for coherent patterns of 

deformation in the CGPS displacement time series.
322 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A B

Figure 12. Examples of heavy winter rime on GPS station 

antennas and tower mounts at altitudes greater than 2,000 m. 

Locations of stations shown in figure 9. A, Station P697. Tower 

stands 3.5 m above ground, which has been blown free of snow. 

USGS photo by W.E. Scott. B, Station P699. Until September 2006, 

station P699 had a solar panel mounted on its GPS antenna mast, 

which exacerbated ice loading and resulted in permanent warping 

of the tower. USGS photo by M. Lisowski.

1 2 3 4 5 6 7 8 9 10 11 12 13 14

DEFORMATION MODE

40

50

60

70

80

90

100

PERCENT OF DATA VARIANCE

NORTH

EAST

UP

Figure 13. Cumulative percentage of data variance in 14 Mount 

St. Helens CGPS residual-component time series accounted 

for by the first 14 PCA deformation modes. North, east, and up 

spatial components are analyzed independently.

Space-Time Correlated Deformation During the 

2004–2006 Eruption

The CGPS station-component time series indicate small 

(total to several centimeters) radially inward and downward 

movements that appear to have slowed during the eruption. To 

determine whether these movements are correlated in time, 

as might be expected if they were in response to pressure loss 

in a magma reservoir, we apply principal component analysis 

(PCA; Preisendorfer, 1988). The PCA decomposes the collective station-displacement time series into a number of deformation modes, each of which consists of a common temporal 

response and a spatial scaling factor that varies between stations 

but remains constant over time at each station. This method of 

decomposing space-time data into the superposition of spaceand time-separable modes is also called empirical orthogonalfunction analysis (Menke, 1984) and eigenanalysis (Aubrey and 

Emery, 1983). Along with a similar method called KarhunenLoeve expansion, Dong and others (2006) used PCA to identify 

and remove common-mode error in daily GPS station-coordinate time series. PCA also has been used to study interseismic 

deformation (for example, Savage, 1988, 1995).

A mathematical description of the method as applied 

in QOCA-based PCA analysis is given by Dong and others 

(2006). Summarizing their discussion, a data matrix Xij is 

constructed such that each column (subscript i in equation 2) 

contains a single residual-displacement component (either 

east, north, or vertical) from a single station in the network. 

The rows of the matrix (subscript j) include displacement 

components from all network stations in each period of measurements (daily averages are used here). PCA allows one to 

represent Xij by

( ) (2) 1

jk

N

k

i j i k X ∑ tCA =

= , (2)

where the N products AikCk(tj) are the individual modes, 

with Aik as the spatially varying scale factor and Ck(tj) as the 

temporally varying time factor for each mode k. If the PCAidentified deformation modes are ordered by their contribution 

to the data variance, spatially correlated signals are contained 

in the first few modes, whereas the higher-order modes usually 

reflect more local effects.

We apply PCA to the cleaned, residual station-component 

time series from the 11 Mount St. Helens CGPS stations with 

the most complete time series (all stations except P690, P697, 

and P699), three PBO CGPS stations (P420, P421, and P432) 

located more than 40 km to the north, and the five distant 

CGPS reference stations (CVO1, KELS, LINH, GOBS, and 

REDM) included in the solutions. Outliers, predicted tectonic 

motion, and cyclical annual and semiannual noise are removed 

from the data. PCA requires observations at all included stations in each epoch. Therefore, Lagrangian interpolation is 

applied to fill small data gaps, and an iterative process that 

uses the predictions of the first three PCA modes is used to fill 

large data gaps (Dong and others, 2006). We examine deformation that occurred after October 15, 2004, when most of the 

new CGPS stations were operational.

The PCA can identify a single coherent mode of time-varying deformation. The first principal mode (mode 1) accounts 

for a large percentage (between 40 and 70 percent) of the data 

variance (fig. 13), particularly in the horizontal components, 

and only mode 1 temporal components (fig. 14A) vary systematically with time. The mode 2 (fig. 14B) and higher-order 

temporal components (not shown) have no long-term changes 

with time and likely represent random noise and systematic 

short-term fluctuations at one or more of the analyzed stations.

Mastin and others (this volume, chap. 22) derive logarithmic and exponential functions to fit the time decay in the 
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 323

2005 2005.5 2006 2006.5

YEAR

10

20

30

40

50

MODE 1 TEMPORAL COMPONENT, IN MILLIMETERS

NORTH

EAST

UP

2005 2005.5 2006 2006.5

YEAR

10

20

30

40

50

NORTH

EAST

UP

MODE 2 TEMPORAL COMPONENT, IN MILLIMETERS

A

B

Figure 14. Temporal components from principal component analysis, in millimeters, from the 

combination of 14 Mount St. Helens CGPS stations and 5 CGPS reference-station residualcomponent time series. Dark lines are the fits to an exponential decay (equation 1) and lightgray lines are 95-percent-confidence limits of that fit. Vertical orange lines as in figures 10 and 

11. A, Mode 1. B, Mode 2.

volumetric rate of dome growth and show that a logarithmic 

decay is expected if flow rate is controlled by frictional 

sliding of a near-surface plug (provided the coefficient 

of friction increases with displacement rate), whereas an 

exponential decay is expected if the effusion rate is a linear 

function of magma-chamber pressure. Ongoing, constant 

recharge adds a linear term to the equations of volumetric 

rate of dome growth.

The mode 1 deformation likely represents the loss of 

pressure in the magma chamber feeding the eruption, and the 

mode 1 temporal component can be fit with curves similar to 

those used by Mastin and others (this volume, chap. 22) to fit 

dome growth. The mode 1 east and north temporal components are best fit by an exponential decay (equation 1) and by 

exponential decay with a linear term (table 7). The relatively 

rapid decrease in the mode 1 vertical component is best fit by 
324 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

an exponential decay with a constant linear increase. The fit 

of the relatively noisy mode 1 temporal component without 

a linear term is only slightly worse. In all cases the inverted 

linear coefficient (d) and the exponential time coefficient 

(b) are highly correlated and, therefore, individually poorly 

determined. The best-fitting curves with the form of equation 

1, along with their 95-percent-confidence limits, are shown in 

figure 14A.

So far we have examined only the PCA temporal components Ck

(tj). Scaling the temporal components by the 

individual station response Aik gives the corresponding station 

displacement. The scatter in the temporal components, when 

Component a

(mm)

b

(per yr)

d

(mm/yr)

Variance

(mm2

)

East 22.057 0.366 1.764

East+linear 22.119 0.381 0.163 1.766

North 26.122 0.544 1.432

North+linear 26.077 0.536 -0.075 1.434

Up 24.333 0.471 15.713

Up+linear 29.207 1.171 4.492 13.112

Table 7. Exponential fits to PCA Mode 1 temporal components.

[Constant coefficients for y=ae –b(t-t0

)

 + d(t–t0) with t

0

 = 2004.853.]

CGPS 

station

Longitude

(degrees)

Latitude

(degrees)

East

displacement (mm)

North

displacement (mm)

Up 

displacement (mm)

JRO1 237.7824 46.2751 5.0 -14.8 -9.6

TGAU 237.8077 46.2192 4.2 -10.3 -15.3

TSTU 237.7759 46.2369 0.5 -17.2 -14.7

P687 237.6454 46.1096 4.3 -0.4 -8.8

P696 237.8484 46.1969 -0.9 -1.5 -10.3

P698 237.8394 46.1735 -2.7 1.2 -10.2

P702 237.6544 46.3002 6.1 -8.6 -5.5

KELS 237.1039 46.1182 3.7 -2.6 0.8

P693 237.7977 46.2103 8.1 -6.4 -9.3

P695 237.8358 46.1990 -3.2 -1.3 -9.2

TWIW 237.8413 46.2129 -11.5 -3.2 -9.3

P420 237.1337 46.5886 1.4 -3.9 -0.1

P421 237.5708 46.5319 2.4 -5.7 -5.0

P432 238.3168 46.6229 0.1 -3.7 -3.3

GOBS 239.1853 45.8388 -0.7 1.3 -0.8

LINH 239.4615 47.0003 0.3 -1.1 0.7

CVO1 237.5039 45.6109 0.1 -1.5 -1.4

REDM 238.8521 44.2598 0.3 -0.1 -0.7

TWRI 237.7881 46.1979 7.8 -4.5 -9.7

scaled, will result in scatter in the displacements. To calculate 

a station displacement for a specified time interval, we need to 

average or smooth the temporal components. The exponential 

fit supplies a smoothing that reproduces the data fairly well, 

and we use it to predict the temporal component and its uncertainty at any given time. The mode 1 horizontal and vertical 

displacements between October 15, 2004, and October 31, 

2006, derived from the exponential fit, are shown in figure 15. 

Displacements are shown only for close-in stations, so that the 

map scale remains large enough to display details of deformation near Mount St. Helens, but all displacement values are 

listed in table 8.

Table 8. PCA Mode 1 station displacements from October 15, 2004, to October 31, 2006.

[Initial and final mode 1 temporal components are scaled by spatial eigenvalue for each station. Estimated 1 

standard deviation uncertainty: east, 1.31 mm; north, 1.18 mm; up, 3.59 mm.]
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 325

 122 ̊20' 122 ̊15' 122 ̊10' 122 ̊05' 122 ̊00'

46 ̊05'

46 ̊10'

46 ̊15'

46 ̊20'

0 5 10

KILOMETERS

AREA

OF MAP

WASHINGTON

WGOT

LVCY

NEBU

LOOW

BIVO

JRO1

TSTU

TWIW

TGAU

TWRI P696

P695

P690 P698

P697

P699

P693

P702

P687 10 MM

VECTOR SCALE -10 MM

HORIZONTAL

VERTICAL

A few general features are apparent in the mode 1 displacements shown in figure 15. Maximum horizontal and vertical 

displacements are roughly equal, and subsidence is observed 

everywhere, with maximum values at distances greater than 

about 3 km and less than about 10 km. Horizontal displacements are toward Mount St. Helens, with maximum values at 

distances greater than about 5 km and less than about 20 km, 

and they are larger to the north of Mount St. Helens than to 

the south. Some of the apparent asymmetry in the deformation 

could result from the fact that there are more CGPS stations 

north of Mount St. Helens. Some other possible explanations for 

the asymmetry include error in the tectonic model that underestimates the background rate of northward motion, a source of 

deformation that is located to the south of the crater, a dipping 

source, and a laterally inhomogeneous crust.

The CGPS station displacements can be supplemented 

with data from a few campaign GPS stations to increase the 

areal coverage. This combination will be used to constrain 

possible sources of deformation. When we have a suitable 

model, we use the PCA mode 1 deformation to estimate the 

change in time.

Analytical Source Models

By using the PCA, a single mode of time-varying 

deformation was identified in the cleaned CGPS time series, 

which consists of an exponentially decaying downward and 

inward collapse toward Mount St. Helens. Such a collapse 

can result from pressure loss in the magma chamber feeding 

the eruption.

A simple model of an erupting volcano includes a magma 

chamber and a conduit to allow magma to escape to the surface (fig. 16). Magma chambers are idealized as fluid-pressurized, ellipsoidal cavities in an elastic half-space (Mogi, 1958; 

Davis and others, 1974; Davis, 1986; Yang and others, 1988; 

Fialko and others, 2001), and conduits are idealized as vertiFigure 15. Map showing PCA mode 1 estimated horizontal (black arrows, tipped with 

95-percent-confidence error ellipses) and vertical (blue arrows tipped with orange 

95-percent-confidence error bars) displacements from October 15, 2004, to October 31, 

2006, for the Mount St. Helens area. CGPS stations P699, P690, and P697 are not included 

in PCA analysis because of offsets and persistent winter excursions.
326 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

cal closed or open pipes (Bonaccorso and Davis, 1999). The 

eruption will continue until overpressure in the magma system 

drops below that of an equilibrated open magmatic system or 

until the conduit is blocked by congealing magma. As magma 

is withdrawn to feed an eruption, the pressure decrease in the 

magma system can be buffered by exsolution of gas (Mastin 

and others, this volume, chap. 22) and by input of new magma 

from depth.

Pressure change in a conduit with a radius of 10–50 m, 

typical for volcanoes (Bonaccorso and Davis, 1999), contributes little to the surface deformation, except near the vent; for 

the same reason, we can ignore deformation produced by shear 

tractions on the wall of the conduit exerted by magma flowing 

through the conduit. We assume that pressure change in the 

magma reservoir beneath the conduit is responsible for most 

of the observed deformation; therefore, we use the data to constrain the approximate limits of source depth, geometry, and 

the cavity volume change of a pressure source that represents 

the magma chamber. We limit our search to spheroidal pressure sources, which include spheres and prolate (elongate) and 

oblate (flattened) spheroids. We assume Poisson’s ratio equals 

0.25, and that the shear modulus equals 30 GPa in all the models. The pressure change that best reproduces the deformation 

is not estimated directly; rather, it is combined with the elastic 

parameters (shear modulus G, and Poisson’s ratio ) and 

the unknown source volume. Paradoxically, except for very 

shallow sources (where the chamber size is a large fraction of 

the depth), a large pressure change in a small source volume 

is indistinguishable from a small pressure change in a large 

source volume (McTigue, 1987). The ratio of pressure change 

MAGMA

INPUT?

OPEN-PIPE

CONDUIT

ERUPTION MODEL

MAGMA

CHAMBER:

SPHERE OR

PROLATE/OBLATE

SPHEROID

ISOTROPIC

ELASTIC

HALF-SPACE

d

a

c

A B

Figure 16. Schematic cross section of erupting volcano (A) and 

pressure sources (B) used to model surface deformation. Source 

depth, d, of magma chamber in half-space is relative to mean 

elevation of volcano.

to shear modulus must remain within the elastic limits of the 

surrounding rock.

Combined CGPS and Campaign GPS Data

Stations in the CGPS network are concentrated around 

the crater and to the north of Mount St. Helens. To improve 

spatial coverage of the deformation, we add data from campaign GPS surveys at five stations located at moderate distances from Mount St. Helens (red triangles in fig. 17). These 

stations were surveyed, at minimum, in late 2004 and again in 

2005. We assume that deformation accumulates at a constant 

rate during the year after October 11, 2004, and obtain average 

station velocities for CGPS and campaign GPS stations (table 

9). Our start and stop dates are arbitrary, but they include 

surveys of all the key campaign GPS stations. The mode 1 

temporal components (fig. 14A) decay at approximately a 

linear rate during this year-long interval, except for the relatively noisy vertical component. The average horizontal (fig. 

17) and vertical (fig. 18) station velocities for this one-year 

period are comparable in direction and relative magnitude to 

those derived from the PCA for a two-year period. Note that 

we include velocities for stations P690 and P697, which were 

excluded from the PCA analysis because of winter anomalies. 

The uncertainties in the station velocities are those from the 

QOCA time-series analysis supplemented in quadrature with a 

random-walk error of 1 mm/yr.

Elastic Half-Space Deformation Source Models

The nearly radial symmetry in the eruption-related 

deformation is consistent with a spheroidal-source geometry, 

and we first fit the data to a spherical-point (also called Mogi) 

pressure source by using a random-cost (Berg, 1993; Murray 

and others, 1996) search algorithm for the optimal location 

and depth of the source. Use of a point-source model is justified because the deformation is widespread, and, therefore, 

consistent with a source that is much deeper than its size 

(McTigue, 1987). The best-fitting spherical point-source depth 

(table 10) is deep (13 km), but, as indicated by the data misfit 

(reduced v2

) of greater than 2, it is a poor choice when the 

maximum vertical and horizontal deformation are roughly 

equal (fig. 19A). A spherical source predicts a maximum 

horizontal deformation that is about 0.4 of the maximum vertical. Some of the misfit of the data to the model may be from 

the tectonic model used to correct the observed velocities. 

To account for such systematic error, we solve for a constant 

horizontal velocity in the data inversion. The point source with 

a systematic velocity shift fits the data slightly better, having a shallower source depth, and the center of deformation 

is shifted slightly to the northwest. The components of the 

translation (1.4 mm/yr west, 2.1 mm/yr north), however, are 

considerably larger than the misfit of the tectonic model to the 

pre-event velocity of JRO1 (0.5 mm/yr west, 0.9 mm/yr north).

A prolate spheroidal (cigar-shaped) pressure source 

predicts roughly equal maximum horizontal and vertical 
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 327

Table 9. Adjusted Helens CGPS average station velocities from October 11, 2004, to 

October 11, 2005.

[Velocities are in nominal ITRF2000 fixed North American Plate terrestrial reference frame and are corrected 

for regional rigid-body rotation and strain.]

Station Longitude

(degrees)

Latitude 

(degrees)

East velocity

 ± 1 r error

(mm/yr)

North velocity

 ± 1 r error

(mm/yr)

Up velocity

 ± 1r error

(mm/yr)

JRO1 -122.2180 46.2750 5.42 ± 1.09 -10.50 ± 1.10 -13.19 ± 1.45

TSTU -122.2240 46.2370 2.92 ± 1.09 -11.02 ± 1.10 -15.48 ± 1.46

TGAU -122.1920 46.2190 3.78 ± 1.10 -5.32 ± 1.11 -12.58 ± 1.51

TWIW -122.1590 46.2130 -5.53 ± 1.10 -2.24 ± 1.11 -8.49 ± 1.50

TWRI -122.2120 46.1980 7.03 ± 1.13 -2.46 ± 1.14 -9.21 ± 1.64

P687 -122.3550 46.1100 4.56 ± 1.10 1.69 ± 1.11 -5.75 ± 1.54

P690 -122.1900 46.1800 0.56 ± 1.43 7.88 ± 1.48 -7.63 ± 2.71

P693 -122.2020 46.2100 7.41 ± 1.14 -2.59 ± 1.16 -8.98 ± 1.67

P695 -122.1640 46.1990 -2.08 ± 1.13 0.50 ± 1.14 -8.41 ± 1.66

P696 -122.1520 46.1970 -1.13 ± 1.10 0.32 ± 1.11 -9.11 ± 1.51

P697 -122.1770 46.1880 1.44 ± 1.33 0.58 ± 1.37 -3.30 ± 2.45

P698 -122.1610 46.1730 -1.72 ± 1.11 2.33 ± 1.12 -10.52 ± 1.56

P702 -122.3460 46.3000 5.44 ± 1.12 -7.10 ± 1.13 -3.45 ± 1.57

NEBU -122.2170 46.1780 5.15 ± 1.77 0.65 ± 1.85 -14.71 ± 3.80

BIVO -122.1960 46.1450 -1.43 ± 1.19 7.55 ± 1.21 -5.01 ± 1.88

LVCY -122.0910 46.1640 -5.56 ± 2.16 3.25 ± 2.23 -11.54 ± 4.69

LOOW -122.1840 46.2250 -0.35 ± 1.84 -7.65 ± 1.91 -10.02 ± 3.91

WGOT -122.3110 46.1730 5.20 ± 2.40 1.90 ± 2.49 -15.80 ± 5.07

P420 -122.8660 46.5890 1.40 ± 1.08 -1.85 ± 1.09 0.94 ± 1.43

P421 -122.4290 46.5320 2.02 ± 1.17 -3.78 ± 1.19 -1.44 ± 1.79

P432 -121.6830 46.6230 -0.65 ± 1.09 -1.94 ± 1.10 -2.39 ± 1.44

KELS -122.8960 46.1180 2.07 ± 1.08 -0.81 ± 1.09 -0.23 ± 1.38

GOBS -120.8150 45.8390 0.86 ± 1.08 -1.55 ± 1.09 -1.10 ± 1.38

CVO1 -122.4960 45.6110 1.41 ± 1.08 0.04 ± 1.09 -0.65 ± 1.38

REDM -121.1480 44.2600 -0.51 ± 1.08 -0.68 ± 1.09 0.65 ± 1.38

LINH -120.5390 47.0000 -1.50 ± 1.08 -0.43 ± 1.09 2.26 ± 1.39

deformation. A point source of this type can be constructed by 

superimposing three collocated orthogonal point cracks (for 

example, Okada, 1985, 1992), with two of the cracks having 

equal amplitudes related to the aspect ratio w of the minor 

radii to the major radius of the spheroid by

M 0 ( ) −= 1.03.0/1 w , (3)

and the third crack having an amplitude

M 0 ( ) w −= 2.04.0/1 . (4)

M0 is the moment, which in this case is equivalent to volume 

change (the amount of opening times the area affected) multiplied by the shear modulus. If we substitute w = 1, then M0

= 5 for the sum of all three cracks, and the source obtained is 

equivalent to a spherical point pressure source. Decreasing w

elongates the spheroid.

A point prolate spheroidal pressure source with its long 

axis vertical fits the data better than do the spherical pointsource models (table 10, fig. 19B). The best-fitting aspect ratio 

(0.66) and depth (7.9 km) is roughly consistent with other 

models of the Mount St. Helens magma chamber (Pallister and 

others, 1992). Only a slight improvement in the fit is obtained 

by adding a systematic horizontal velocity (0.2 mm/yr west, 

1.4 mm/yr north), which shifts the center of the source to the 

southwest. The estimated velocity shift is similar to the misfit 

of the tectonic model to the pre-event velocity of JRO1.
328 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

 122 ̊20' 122 ̊15' 122 ̊10' 122 ̊05' 122 ̊00'

46 ̊05'

46 ̊10'

46 ̊15'

46 ̊20'

0 5 10

KILOMETERS

AREA

OF MAP

WASHINGTON

+

+

WGOT

LVCY

NEBU

LOOW

BIVO

P696

P695

P697

P698

P690

P699

P693

TSTU

TGAU

TWIW

TWRI

P687

P702

JRO1

+

VELOCITY SCALE

MODEL

OBSERVED

10 MM/YEAR

0 5 10

KILOMETERS

TGAU

TWIW

JRO1

P702

P693

P695

P696 P697

P698

P699

P690

P687

TWRI

122 ̊20' 122 ̊15' 122 ̊10' 122 ̊05' 122 ̊00'

46 ̊05'

46 ̊10'

46 ̊15'

46 ̊20'

AREA

OF MAP

WASHINGTON

WGOT

LVCY

NEBU

LOOW

BIVO

VELOCITY SCALE MODEL OBSERVED -10 MM/YEAR

TSTU

Figure 17. Map showing corrected 

average GPS station horizontalvelocity vectors (black arrows tipped 

with 95-percent-confidence error 

ellipses) and displacements predicted 

by best-fitting, tilted point prolate 

spheroid (red arrows, scale at bottom 

of figure) for the Mount St. Helens 

area from October 15, 2004, to October 

15, 2005. See table 10 for model 

parameters. Surface projection of 

center of tilted point prolate spheroid 

marked by red cross, that of vertical 

prolate spheroid by blue cross, and 

that of spherical point pressure source 

by green cross.

Figure 18. Map showing corrected 

average GPS station vertical-velocity 

vectors (black arrows tipped with light 

blue 95-percent-confidence error bar) 

and vertical displacements predicted 

by best-fitting, near-vertical, pointprolate spheroid (red arrows, scale at 

bottom) for Mount St. Helens area from 

October 15, 2004, to October 15, 2005. 

See table 10 for model parameters.
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 329

60 40 20 0 20 40 60

DISTANCE, IN KILOMETERS

-15

-20

-10

- 5

0

5

60 40 20 0 20 40 60

DISTANCE, IN KILOMETERS

-15

-20

-10

-5

0

5

-15

-20

-10

-5

0

5

60 40 20 0 20 40 60

DISTANCE, IN KILOMETERS

A

B

SPHERE

VERTICAL

PROLATE

SPHEROID

PROLATE

SPHEROID

PLUNGE 83° 

N37°W

SOUTH NORTH

VELOCITY, IN MILLIMETERS PER YEAR

C

Figure 19. Observed radial (blue squares) and vertical (red 

diamonds) velocities for deformation of Mount St. Helens from 

October 15, 2004, to October 15, 2005, projected onto northsouth profile and compared with predicted radial (solid blue 

curve) and vertical (red curve) velocities from best-fitting point 

sources. Error bars are ±1 standard deviation. Details of models 

given in table 9. A, spherical point pressure source; B, vertical 

point prolate spheroid; C, tilted or dipping point prolate spheroid 

with a small velocity correction. 

Another way for the model to account for asymmetry in 

the deformation is to allow the source to plunge at an angle 

less than 90o and to allow the azimuth of this axis to vary. If 

we do not allow for a systematic rate of translation, the parameter search returns a vertical optimal plunge. If we allow for 

a systematic velocity error, the best-fitting, plunging prolate 

source is nearly vertical (83o) with azimuth N. 37o W. (fig. 

19C). The systematic velocity is again similar to the misfit of 

the tectonic model to the pre-event velocity of JRO1.

To calculate the volume change consistent with the 

source strength, we assume that the spherical source has a 

radius of 1 km and the prolate spheroid has a minor radius 

of 1 km. Once the model geometry is fixed to the values 

obtained by the parameter search, the radius change (equivalent to strength of the source) is inverted directly. The equivalent source cavity-volume decrease for the one-year period 

covered by the data is in the range 8–12×106 m3 in the various 

prolate spheroid models.

Only JRO1 recorded the relatively rapid deformation that 

occurred during the first two weeks of the eruption, and we use 

its motion as a proxy to estimate the total cavity-volume loss. 

The predicted 12.8 mm/yr of radially inward motion at JRO1 

from the dipping prolate spheroid model with a systematic 

velocity correction is about half of the total movement of JRO1 

from September 23, 2004, to October 31, 2006. A total cavityvolume loss of 16–24×106

 m3 is consistent with the JRO1 

displacements, and it is a fraction of the >83×106 m3 dome 

volume as of August 2006. Mastin and others (this volume, 

chap. 22) explain this discrepancy as a combination of exsolution of gas, which buffers pressure loss in the magma chamber, 

and recharge by new magma (recharge estimated to be only 

about one-tenth of the erupted volume). The recharge, if it is 

occurring, has thus far not increased the flux of volcanic gases 

(Gerlach and others, this volume, chap. 26).

Conclusions

Remarkably little far-field volcanic deformation has 

occurred around Mount St. Helens since shortly after the 

crater-forming collapse and eruption in 1980. Data collected in 

1982 and 1991 for surveys of a regional high-precision trilateration network provide the clearest evidence for recharge of the 

volcano’s magma system. During that interval, areal dilatation accumulated at an average rate of 144±39 nanostrain/

yr, an order of magnitude greater than and opposite in sign 

to the estimated background rate of −15±3 nanostrain/yr. No 

significant deformation occurred in the same network between 

1991 and 2000.

GPS surveys conducted since the 1990s reveal background deformation, which is used to define a tectonic 

model that includes a rigid-body clockwise rotation and a 

small amount of strain accumulation. The background strain 

accumulation within the network (in the absence of volcano 

deformation) is essentially a uniaxal contraction in the eastnortheast direction, a possible effect of a locked Cascadia 

subduction-zone thrust fault. After removing this secular 

deformation, we find no coherent pattern of volcanic deformation around Mount St. Helens between 2000 and 2003. 

Furthermore, daily position measurements made at a CGPS 
330 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Source 

description

Longitude

(degrees)

Latitude

(degrees)

Depth

(km)

∆Radius

(m/yr)

∆East

(m/yr)

∆North

(m/yr)

Aspect

ratio

Long axis

plunge

(degrees)

Long 

axis

(trend)

Model 

parameters

Reduced 

χ2

∆Volume

(106

 m3/

yr)

Spherical point 

pressure (Mogi) -122.1765 46.1961 12.93 -0.7653 -- --

1.00 

(fixed) -- -- 4 2.17 9.7

Spherical point 

pressure + 

translation rate

-122.1926 46.2119 10.39 -0.5028 -0.0014 0.0021 1.00 

(fixed) -- -- 6 1.95 6.3

Vertical prolate 

spheroid -122.1746 46.2133 7.99 -1.1485 -- -- 0.66 90 (fixed) n.a. 5 1.59 11.9

Vertical prolate 

spheroid + 

translation rate

-122.1758 46.2000 7.60 -0.9651 -0.0002 0.0014 0.673 -- n.a. 7 1.54 9.6

Tilted prolate 

spheroid -122.1748 46.2133 7.94 -1.1356 -- -- 0.660 90 (fixed) n.a. 7 1.60 11.9

Tilted prolate 

spheroid + 

translation rate

-122.1811 46.2133 7.14 -0.7927 -0.0004 0.0013 0.676 82.4 N36°W 9 1.49 7.9

Table 10. Elastic half-space deformation source models.

[ΔEast and ΔNorth data show estimates for rigid-body translation rate that account for possible error in tectonic model. --, elements not part of solution; n.a., not 

applicable to long-axis trend because model is of vertical prolate geometry.]

station 9 km north of Mount St. Helens show no anomalous 

deformation from 1997 until the start of seismic unrest on 

September 23, 2004.

The movement of magma to the surface at the start of 

unrest did not systematically displace GPS stations on the volcano’s flanks, but it did displace station DOM1 on the September 1984 lobe of the old dome by more than 0.2 m to the north. 

The lack of systematic deformation around the crater suggests 

that solidified magma plugging the top of the conduit was 

pushed out, rather than being bypassed. Much of the movement of the L1 CGPS station DOM1 occurred by September 

28, well before the first phreatic explosion, suggesting that the 

movement of magma began with the start of seismic unrest. 

At the same time, CGPS station JRO1 moved toward Mount 

St. Helens an average of about 0.5 mm/day from September 

23 until a few days after the phreatic explosion on October 5, 

2004, when its motion slowed and began to decay.

Analysis of eruption-related deformation data reveals a 

single space- and time-coherent mode of time-varying deformation, which has been decaying exponentially since the start 

of unrest. All GPS stations are moving toward Mount St. Helens and downward in a manner consistent with the predictions 

of an elastic half-space model that represents pressure loss in 

a vertically elongate magma chamber. The center of Mount St. 

Helens’ magma chamber is estimated to be at a depth of 7 to 8 

km, and the chamber’s long axis, which is vertical or near vertical, is estimated to be approximately three times longer than 

its diameter. The actual dimensions of the magma chamber 

are poorly constrained by the deformation data. The sourcecavity volume loss, which is proportional to the strength of 

the source, is about 16–24×106

 m3, about one-third of the 

erupted volume. This discrepancy can be accounted for mostly 

by buffering of pressure loss in the magma chamber through 

exsolution of gas and with a small amount of recharge (Mastin 

and others, this volume, chap. 22).

Acknowledgments

We thank the NSF-sponsored Plate Boundary Observatory for rapid installation of CGPS monitoring stations in 

October 2004 in response to renewed unrest at Mount St. 

Helens. In particular, we are grateful for the timely efforts of 

Katrin Hafner and other field engineers from UNAVCO, Inc., 

the contractor responsible for installing PBO stations.

We thank the CVO staff, temporary staff, and volunteers who helped with the field surveys and permanent CGPS 

station construction. We are grateful that Elliot Endo had 

the foresight to install a CGPS station at the Johnston Ridge 

Observatory. Richard LaHusen established telecommunications at the U.S. Forest Service’s Coldwater Ridge maintenance shed in order to download CVO GPS data, and he 

provided L1 CGPS data. We used computer code from John 

Langbein of the USGS Earthquake Hazards Team to invert 

for best-fitting source-model parameters. The original code 

was developed by James C. Savage; Mark H. Murray added 

random-cost extensions and the Okada dislocation routines, 

and John Langbein added code for the ellipsoidal pressure 

sources. John Langbein and Larry Mastin provided constructive reviews of this paper.
15. Analysis of GPS-Measured Deformation Associated with the 2004–2006 Dome-Building Eruption 331

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Chapter 16

Instrumentation in Remote and Dangerous Settings; 

Examples Using Data from GPS “Spider” Deployments During 

the 2004–2005 Eruption of Mount St. Helens, Washington

By Richard G. LaHusen1, Kelly J. Swinford1, Matthew Logan1, and Michael Lisowski1

Abstract

Self-contained, single-frequency GPS instruments fitted 

on lightweight stations suitable for helicopter-sling payloads 

became a critical part of volcano monitoring during the 

September 2004 unrest and subsequent eruption of Mount St. 

Helens. Known as “spiders” because of their spindly frames, 

the stations were slung into the crater 29 times from September 2004 to December 2005 when conditions at the volcano 

were too dangerous for crews to install conventional equipment. Data were transmitted in near-real time to the Cascades 

Volcano Observatory in Vancouver, Washington. Each fully 

equipped unit cost about $2,500 in materials and, if not 

destroyed by natural events, was retrieved and redeployed as 

needed. The GPS spiders have been used to track the growth 

and decay of extruding dacite lava (meters per day), thickening 

and accelerated flow of Crater Glacier (meters per month), and 

movement of the 1980–86 dome from pressure and relaxation 

of the newly extruding lava dome (centimeters per day).

Introduction

Typically, volcano monitoring and associated eruption 

forecasting relies on several disciplines of volcanology, principally seismology, gas geochemistry, and geodesy (Dzurisin, 

2006). No single tool or technique can adequately monitor or 

predict the range of volcanic behaviors—from aseismic deformation to relatively benign dome building to major explosive 

eruptions. Accordingly, volcanologists rely on an assortment 

of instruments and techniques to monitor volcanic unrest. 

Sensors and related instrumentation have been developed in 

attempts to accommodate the needs of each particular discipline. However, even when an instrument is available that is 

capable of making a desired measurement, use of the instrument may be limited by expense or by an inability to deploy it 

in dangerous or inaccessible sites close to volcanic vents. New 

techniques and instruments are notable to volcanologists when 

they are affordable and minimize exposure of personnel to 

hazards. This paper describes the rapid development and application of a self-contained instrument package that was used 

successfully to monitor deformation close to the vent during 

the renewed eruption of Mount St. Helens in 2004–2005.

Prior Near-Vent Geodesy at Mount St. Helens

Pioneering geodetic work done during the 1980–86 eruptions of Mount St. Helens demonstrated that some eruptions 

could be predicted by monitoring accelerating deformation on 

localized parts of the dome (fig. 1; Swanson and others, 1983; 

Dzurisin and others, 1983). The 1980s geodetic work demonstrated that substantial preeruptive dome deformation typically was limited to areas near the active vent. The geodeticdeformation measurements commonly required the repeated 

and prolonged presence of personnel working close to the vent 

to bury electronic tiltmeters or to measure distances between 

fixed monuments by using electronic distance meters. Fortunately, with the subsequent availability of commercial GPS 

instruments, it is now possible to make repeated high-precision 

geodetic measurements on the volcano without exposing personnel to prolonged periods of work in hazardous areas.

Prototype GPS Instrument

For 4 years prior to the September 2004 seismic unrest at 

Mount St. Helens, an automated L1 GPS system was intermittently operated as a prototype monitoring tool. The system was 
336 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

100

200

300

JULY AUG.

1981

SEPT.

CUMULATIVE CONTRACTION, IN CENTIMETERS

ERUPTION

Figure 1. Movement of monument on lava 

dome in 1981 as determined by sequential EDM 

measurements, Mount St. Helens, Washington. 

Arrow indicates issuance of eruption prediction 

for time window shown by rectangular bar. 

Vertical dashed line shows time that dome 

eruption began (from Swanson and others, 1983).

being developed by the U.S. Geological Survey’s Cascades 

Volcano Observatory (CVO) as an inexpensive, near-realtime, ground-deformation monitoring tool with design 

emphasis on low cost, power conservation, and telemetered 

data integrity (LaHusen and Reid, 2000). Station DOM1 was 

installed on the 1980–86 dome within the crater of Mount St. 

Helens, and stations SFT2 and POA3 were installed outside 

the crater about 4 km to the west and east of DOM1, respectively. Each station was powered by a 20-W solar panel and 

a rechargeable lead-acid battery. Within each was a USGS 

microcontroller controlled remotely from a desktop computer 

at CVO using a 100-km multi-hop 900-MHz radio link. The 

operational scheme conserves power at the remote stations 

by alternately powering the GPS receivers and radio-telemetry components. During data-acquisition cycles, a CMC 

Allstar model L1 GPS receiver coupled with a Micropulse 

model 1372 survey-grade GPS antenna is powered for 20 

minutes while raw GPS data at 10-s epochs are logged to 

the microcontroller. At the end of each data-logging session, 

the GPS receiver is switched off to conserve power, and the 

radio is switched on to relay data packets to CVO. Each datatransmission session lasts several minutes and includes error 

checking and retransmission to achieve error-free blocks of 

raw GPS data. The operational duty cycle is adjustable to 

balance power conservation, frequency of measurement, and 

station longevity.

Fixed, static, double-differential solutions between stations were calculated automatically using USGS control and 

scheduling software that applied a commercial software module, Waypoint Precise DLL. An independent position solution 

was calculated for every 20-minute data-acquisition period. 

Although GPS data acquisition within the crater posed substantial challenges owing to obscured views of the sky, mantling by 

rime ice, and noise from multipath reflections, this monitoring 

scheme typically transferred data reliably and repeatedly. The 

accuracy of each single solution was 1 cm plus 1 ppm of the 

differential baseline length (4 km), or 1.4 cm for the horizontal 

components. Vertical accuracy was found to be about double 

the horizontal value. This noise, inherent to single-frequency 

GPS solutions, was reduced greatly by applying a moving median filter that allowed discrimination of more subtle 

motions of less than a centimeter over longer time periods. 

Between 2002 and February 2004, the system measured dome 

subsidence that possibly reflected contractive cooling of the 

dome interior. The subsidence was at an annual rate of about 

8 cm/yr downward and 2 cm/yr eastward, toward the center of 

the dome (fig. 2). Batteries at DOM1 failed in February 2004 

and were not replaced until September.

When the GPS system was reactivated on September 27, 

2004, following initiation of seismicity at Mount St. Helens, 

station DOM1 was in a location significantly different from 

what would have been predicted based on the subsidence trend 

of the preceding years. It was about 20 cm north and 

12 cm higher than the predicted location (fig. 2). These location changes occurred between February 2004 and September 

27, 2004, but we cannot further constrain the rate and timing 

of this deformation. Automated measurements on September 

27 showed that DOM1 was moving northward at 2 cm per day. 

This station and a companion seismometer, the only stations 

on the old dome, were destroyed 4 days later, on October 1, by 

ballistic fragments ejected during the first phreatic explosion 

of the 2004 eruption (fig. 3).

GPS Spider Deployments

Loss of the only GPS station in the crater presented a 

dilemma because near-vent deformation could no longer be 

monitored. The time-consuming task of reinstalling a solarpowered GPS station onsite was considered unsafe because of 

the risk of additional, unpredictable phreatic explosions. Other 

challenges for deformation monitoring in the crater included 

the likelihood of additional ashfall from explosions and the 

impending onset of winter snowfall.

As an alternative to installing a permanent, monumented 

station, we assembled three GPS stations electronically similar 

in design to the original DOM1. However, unlike DOM1, which 
16. Instrumentation in Remote and Dangerous Settings; Examples Using Data from GPS “Spider” Deployments 337

7/1/02 1/1/03 7/1/03 

DATE

1/1/04 7/1/04 11/1/04

0

0.1

-0.1

-0.2

-0.3

-0.4

0.2

0.3

0.4

Northings

Eastings

Vertical

DISPLACEMENT, IN METERS

DOM1

A B

Figure 2. Time series of preeruptive deformation at GPS station DOM1, Mount St. Helens, 

Washington. Prior to February 2004, the 1980–86 dome was subsiding, possibly due to 

contraction of cooling interior. When DOM1 was reactivated on September 27, 2004, it was 20 

cm north and 12 cm above its predicted location based on its previous trend (dashed lines).

Figure 3. Photographs showing station DOM1 GPS before and 

after its destruction by ballistic fragments from explosion at Mount 

St. Helens on October 1, 2004. A, Before explosion. Antenna 

mount (foreground behind boulder) and instrument shelter (black 

box), with helicopter hovering just beyond. USGS photo by R.G. 

LaHusen. B, After explosion, from slightly closer viewpoint. White 

jagged sheet of plywood is the only side of instrument box still 

erect. USGS photo by S.C. Moran, November 5, 2004.
338 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

was pieced together on site, these new stations were constructed 

as self-contained portable units. Each unit consisted of an aluminum case that housed a 1,200-Ah supply of nonrechargeable airalkaline batteries and a weatherproof ABS plastic case enclosing 

the electronic components. A GPS antenna was attached to a 

1.5-m-long steel pipe mast on one end of the aluminum case, 

and a similar mast on the other end supported a radio antenna. 

A rope sling with a swivel eye was bolted to the outer metal 

case for attachment to a helicopter sling-cable remote-release 

hook. Each of these three stations weighed approximately 70 kg, 

and parts for each unit cost about $2,500, making it practical to 

deploy several stations and constituting an acceptable loss if a 

station were destroyed. With this design, stations could be set in 

the crater near the source of the recent explosion, and personnel 

would be exposed only briefly to potential hazards. Initial results 

from these portable installations were promising, demonstrating 

that GPS stations could be installed quickly to provide repeatable 

results at centimeter accuracy in near-real time. One unit toppled 

shortly after deployment, so we redesigned the frame with three 

widely spaced legs for better stability on rocky, uneven terrain.

Spider Frame Design

Field tests of several frame prototypes helped to determine design requirements and constraints. These included the 

need for a frame that was: (1) strong and rigid for its weight, 

(2) corrosion resistant, (3) relatively inexpensive and simple to 

build, (4) capable of accommodating various onboard electronic 

instruments and antennas, (5) capable of being slung safely 

beneath a helicopter, (6) capable of deployment by helicopter sling cable onto uneven and rocky terrain, (7) stable when 

placed on uneven and rocky terrain, and (8) capable of being 

retrieved by grappling hook from a helicopter for redeployment 

to another site or transport to a safe location to replace batteries.

Using these design considerations, we constructed 18 

frames at CVO (fig. 4). The frames were built of type 6061 

stainless steel square tubing with 1.6-mm-thick walls and a 

38×38-mm cross section. The tubing was cut with a horizontal 

band saw and welded using stainless steel welding wire.

The large leg span (1.4 m) of the three-legged frames 

and the low center of gravity of the welded aluminum battery 

and electronics box (0.34 m above ground surface) provided a 

stable platform for the onboard batteries and electronics. The 

GPS antenna was mounted to one of the three framing legs that 

extended higher than the other two, 1.7 m from the ground. 

A short leg served as a mount for a 1-m-long omnidirectional 

antenna for the data transceiver.

A pyramidal stainless steel tripod canopy of the same 

tubular material was welded to the legs, culminating 1.8 m 

above the ground where a lifting eye, cable, and swivel were 

attached. In addition to strengthening the frame, this canopy 

provided a centered attachment point for airborne transport 

and presented a large, open target for helicopter retrieval using 

a grappling hook. With the addition of the spindly-legged 

framework, these self-contained stations took on the appearance and nickname of “spiders.”

Canopy

Electronics

and batteries

in box

Leg

GPS

antenna

Lifting cable

Swivel and ring

Radio

antenna

Figure 4. Photograph of GPS spider with stainless steel leg and 

canopy framework (leg span, 1.4 m) and aluminum box housing 

the unit’s electronics and batteries. GPS antenna is located 

on extended leg in right foreground, and data radio antenna is 

attached to right rear leg. Stainless steel lifting cable with swivel 

and ring is attached to top of canopy. USGS photo by M. Logan.

The spider is stable when in flight beneath a helicopter, 

and the low center of gravity and wide leg span make the station 

resistant to tipping during deployment and retrieval, as well 

as resistant to toppling by high winds after deployment. Over 

the first 18 months of use, the stainless steel frames did not 

deteriorate in the harsh volcanic environment except for damage 

from direct ballistic impact. During 2004–5, we experimented 

with additional instruments on board some spiders including 

tiltmeters, cameras, gas sensors, and seismic accelerometers 

(McChesney and others, this volume, chap. 7).

Examples of GPS Spiders at Mount St. 

Helens, 2004–2005

GPS Spiders on the 1980–86 Dome

Remotely placed GPS spiders have occupied five sites on 

the 1980–86 dome. These sites were about 400–500 m from 
16. Instrumentation in Remote and Dangerous Settings; Examples Using Data from GPS “Spider” Deployments 339

the center of September 2004 deformation, ensuing phreatic 

explosions, and lava-spine extrusion. The history of emplacement and subsequent life is summarized in table 1; distribution 

and mapped progression is shown in figure 5.

The first three portable GPS stations were set on the 

north, east, and approximate center of the 1980–86 dome on 

October 3 and 4, 2004; station designations were NDM5, 

NRM6, and TOP7, respectively. Results from these installations showed that (1) motion of the three stations was similar, 

indicating the relative stability of the remotely deployed packages, and (2) the northward movement of the 1980–86 dome 

was continuing and was not restricted to the area of DOM1 

on the west side of the dome. The 1980–86 dome lay north of 

the vent, so these spiders typically moved slightly northward. 

It appeared that the entire 1980–86 dome was moving to the 

Station Start date End date Displacement (m) Fate

1980–86 dome

DOM1.0 1 9/27/04 10/1/04 -- Explosion October 1, 2004

NDM5 10/3/04 1/2/05 0.15 Toppled, recovered

NRM6 10/4/04 11/20/04 -- Toppled, recovered

TOP7 10/4/04 3/8/05 -- Explosion March 8, 2005

DOM1.1 11/6/04 3/8/05 0.4 Explosion March 8, 2005

NEDB 11/20/04 3/8/05 0.16 Explosion March 8, 2005

DOM1.2 3/10/05 8/5/05 0.1 Moved

NEDB.1 4/6/05 3/2/06 0.1 Batteries expired

DOM1.3 8/5/05 12/25/05 0.05 Batteries expired

HIE5.2 9/21/05 12/31/05+ 0.05 Continued operating into 2006

Middle zone

MID9.0 10/27/04 2/3/05 1.4 Explosion damage Jan. 16, 2005

MIDE.0 2/11/05 3/8/05 0.22 Explosion March 8, 2005

MID9.1 4/6/05 6/29/05 2.3 Removed

MID9.2 11/17/05 12/16/05 0.17 Buried by talus

New dome

CLF4 10/27/04 1/21/05 48 Rockfall

ELEA.0 11/20/04 11/27/04 67 Rockfall

HNY0 1/3/05 1/29/05 78 Rockfall

CDAN 1/15/05 1/16/05 8 Explosion Jan. 16, 2005

AHAD 2/8/05 2/16/05 41 Removed

ELE4.0 4/19/05 4/21/05 2 Moved

SEV7 5/24/05 3/24/06 10 Batteries expired

East Crater Glacier

ICY4 2/16/05 4/8/05 22 Lost in crevasse

ICY5.0 2/16/05 3/8/05 28 Explosion March 8, 2005

ELE4.1 4/21/05 6/30/05 26 Moved

ELE4.2 6/30/05 7/28/05 6.7 Moved

HIE5.0 7/18/05 8/19/05 8.5 Moved

ELE4.3 7/28/05 8/19/05 7.8 Moved

West Crater Glacier

WES6 7/14/05 9/14/05 70 Removed

ELE4.4 8/19/05 11/9/05 122 Batteries expired

HIE5.1 8/19/05 9/14/05 24 Moved

1 Permanent monumented station; restarted Sept. 27, 2004.

Table 1. History of GPS spider deployments in the crater of Mount St. Helens, Washington, 2004–2005.

north, away from the vent area, as if it were being shouldered 

aside in response to the growing mass beneath its south flank. 

Consequently, the northing component of the GPS solution 

best depicts the predominant motion. Magnitude of total 

motion during this period was less than 1 m (fig. 6). Rates 

were as much as 2 cm per day northward. Vertical displacement was typically within the noise of the analysis.

The 1980–86 dome appeared to behave as a semicoherent block so that, with few exceptions, spiders at different 

sites responded similarly (fig. 6). The rate of displacement 

varied with time. The limited evidence suggests that the rate of 

old-dome movement varied directly with the linear extrusion 

rate of the new dome, as shown by the coincidence of highest 

1980–86 dome displacement rates with the greatest velocity for spiders riding on the extruding spine (fig. 6). We were 
340 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

11/1/04 12/1/04 1/1/05 

DATE

2/1/05 3/1/05 

HORIZONTAL DISPLACEMENT, IN METERS

0

-0.1

0.1

0.2

0.3

0.4

0.5

0.6

0.7

TOP7

DOM1

10 m/d

ELEA

6 m/d

AHAD

8 m/d

CDAN

Figure 5. Shaded-relief digital elevation model (2003) showing GPS spider deployments (red triangles) in 

crater of Mount St. Helens, Washington. Tracks of triangles indicate total movement of stations (table 1).

Figure 6. Graph showing movement of two GPS spiders on 

1980–86 dome (DOM1, TOP7) and three GPS spiders on 2004–5 

dome (ELEA, CDAN, and AHAD) during first 5 months of eruption. 

Location of stations shown in figure 5. The green boxes show 

duration of new-dome spiders, each labeled with elapsed 

horizontal velocity, which was nearly halved from November 

2004 to March 2005. These three spiders had substantial 

displacements (as much as 67 m over a 7-day period for ELEA), 

which makes it difficult to portray their data in a manner 

comparable to the old-dome spiders TOP7 and DOM1.

0 500 METERS

WEST

CRATER

GLACIER

2004–2005

DOME

1980–1986

DOME EAST

CRATER

GLACIER

46°12'

46°11'30"

122°12' 122°11'30" 122°11' 122°10'30"

HIE5.1

CLF4

CDAN

NDM

NEDB

WES6

SEV7

TOP7

DOM1.X

HNY0

AHAD

ELE4.1

ELE4.4

ELEA.0

ELE4.3

ELE4.2

ICY5.0

ICY4

HIE5.0

HIE5.2

ELE4.0

MIDE.0

MID9.X
16. Instrumentation in Remote and Dangerous Settings; Examples Using Data from GPS “Spider” Deployments 341

7 8 9 10 11 12 13 14 15

JANUARY 2005

16 17 18 19 20

NORTHINGS,

IN METERS

1-MINUTE RSAM,

IN COUNTS

-0.02

0

0.02

0.04

0.06

0.08

0

20

40

60

80

100

120

140

SEP

TOP7

unable to maintain a spider on the extruding spine long enough 

to demonstrate this relation more convincingly.

Intermittently, velocity of GPS spiders on the old dome 

increased as the rate of seismicity increased, as measured by 

real-time seismic amplitude measurements, or RSAM (Endo 

and Murray, 1991; Moran and others, this volume, chap. 2). 

One example of this correlation occurred on January 15, 2005, 

when TOP7 and other near-vent stations accelerated away from 

the vent as RSAM values started to increase (fig. 7). An explosion the following day destroyed several spiders (table 1).

GPS Spiders on the New Dome

Measurements of extrusive rates are the most effective 

way to monitor the progress of effusive eruptions. The best 

day-to-day proxy for extrusive rate at Mount St. Helens was 

the linear rate at which lava emerged from the vent; thus, we 

first targeted sites on and immediately adjacent to the actively 

extruding dome. Prior to new lava appearing at the surface 

on October 11, 2004, intense surface deformation created an 

uplift of part of the 1980–86 dome, deformed glacier ice, and 

crater-floor debris, or the “welt,” through which the first lava 

spines emerged (Vallance and others, this volume, chap. 9). 

Sites on the welt had the advantage of longevity compared to 

placement directly on the active lava spines and were used first 

in October 2004. Although not directly on the extruding lava, 

these spiders were useful as proxies of eruption vigor because 

they were being actively pushed away from the vent. By late 

November 2004, we began placing spiders directly astride the 

actively extruding lava spine.

Figure 7. Graphs showing relation 

between increases in real-time 

seismic amplitude (RSAM) and 

northward motion of station TOP7 in 

mid-January 2005. Seismic station 

SEP is located 100 m northwest of 

TOP7, on the September lobe of the 

1980–86 dome.

The first spider was set on the welt adjacent to the northeast side of the new dome on October 12, 2004, but its GPS 

antenna was damaged on deployment, rendering it geodetically useless. The spider also had seismic instrumentation on 

board that continued to function (McChesney and others, this 

volume, chap. 7). The next GPS spider, station CLF4, was 

installed on October 27, adjacent to the south flank of newly 

emergent spine 3, on uplifted crater-floor debris of the welt 

(fig. 5; table 1). Station CLF4 moved downward and to the 

south, traveling 10 m in the first three days as spine 3 plowed 

across the crater floor. Station CLF4 was pushed away from 

the growing spine as if riding the bow wave of a ship. It operated for 85 days, outliving the growth period of spine 3 and 

persisting into the first third of spine 4’s life before succumbing to rockfall.

Some of the most interesting correlations between eruptive phenomena and movement of GPS spiders came from 

site MID9, which was on the saddle midway between the new 

and old domes (figs. 5, 8). On November 12, November 22, 

and December 8, 2004, MID9.0, the first of the MID9 spiders 

(table 1), accelerated away from the vent (fig. 9) coincident with 

increased RSAM counts. These velocity changes appeared as 

surge-pause-surge phases that lasted several days. The correlation with RSAM was masked for some other events that 

occurred in stormy weather, owing to heightened seismic noise 

that accompanied high wind.

On December 21, MID9.0 stalled and slowly reversed 

direction, heading back toward the vent and downward (figs. 

9, 10). This change may have been in response to depressurization within the conduit and relaxation at the surface. 

Other crater GPS stations also slowed or stopped moving away 

from the vent. This event coincided with a change in seismicity, during which several large 

earthquakes had downward first 

motions as opposed to the typical 

pattern of upward first motions (S. 

Malone, written commun., 2004). 

But station MID9.0 continued its 

southward and down motion into 

2005 while extrusion continued, 

so a more likely explanation for 

the ventward motion is that ventadjacent bedrock was shifting in 

response to some other factor. 

Because this area between the 
342 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

11/1/04 11/15/04 11/29/04 12/13/04

DATE

12/27/04 1/10/05 1/24/05

DISPLACEMENT, IN METERS

0.0

-0.2

-0.4

-0.6

0.2

0.4

0.6

0.8

1.0

1.2

Eastings

Northings

Vertical

domes was so responsive to changes in eruption dynamics, 

we tried to keep a functioning GPS spider near the original 

MID9.0 location. Accordingly, after MID9.0 suffered antenna 

damage in an explosion on January 16, it was replaced by 

MIDE.0 followed by MID9.1 and MID9.2 (table 1).

On November 20, spider ELEA.0 was placed on the 

highest point of spine 3 (fig. 5). Its initial motion was an 

astonishing 10 m per day. Before it was destroyed by a rockfall 6 days later, ELEA.0 had moved 67 m south-southeastward and 8 m up.

The remarkable record of ELEA.0 reinforced our decision to keep a GPS spider on the actively growing spine of the 

new dome. A GPS spider, especially one carrying an accelerometer (seismometer), could add substantially to the amount 

of data available for evaluating steady-state extrusion models 

and for locating seismic sources more precisely. The task of 

slinging spiders onto hot lava spines proved more difficult 

than expected, owing to turbulence and decreased lift in heated 

air. Thermal surveys showed surface temperatures of spines 

away from the vent typically were less than 50°C, but cracks 

exposed interior lava as hot as 700°C (Vallance and others, 

chap. 9; Schneider and others, this volume, chap. 17) and created strong hot updrafts. A successful installation on January 

3, 2005, deployed spider HNY0 (Happy New Year) on a large 

block near the south end of spine 4 (fig. 5). For more than 3 

weeks, spider HNY0 moved 2–2.5 m/day east-southeastward, 

consistent with eastward spreading as the south end of the 

spine broke apart.

Because spine 4’s surface rose steeply from the vent, we 

built a spider with legs of differing length, cut appropriately 

to match the spine’s 32° slope. This spider, CDAN, held fast 

in the soft fault gouge that mantled the spine (fig. 11). However, station CDAN was toppled and buried in talus during an 

explosion on the day after its installation, but not before it had 

moved about 8 m southward and upward (table 1).

Spiders were sited on the active spine sporadically into 

2006 (table 1). Spines 5–7 grew more vertically than their 

recumbent predecessors and became increasingly mantled 

by talus, preventing us from finding sites suitable for setting 

a spider. Sites in talus were notably perilous, owing not only 

to instability of blocks on which spiders could be placed, but 

also to destruction by rockfalls from upslope debris. Summits 

of spines were free of risk from rockfall, but unlike the earlier 

recumbent spines that extended hundreds of meters from the 

vent—the site of greatest heat discharge—spines 5–7 grew 

more steeply, formed smooth surfaces for relatively short 

distances, and produced mostly rubble and large flanking talus 

aprons. Consequently, the summits of spines 5–7 were too hot 

and the air there too turbulent to safely deploy the instruments.

Spiders on Crater Glacier

Throughout the 2004–5 eruption, 

lava-dome growth through formerly 

horseshoe-shaped Crater Glacier has 

caused dramatic disruption and deformation (Walder and others, this volume, 

chap. 13). This remarkable process was 

documented with intermittent aerial 

photographs and creation of digital 

elevation models (Schilling and others, 

this volume, chap. 8), albeit with some 

Figure 8. January 2005 photograph of station MID9 site (red box) 

between 1980–86 dome on the right and snow-free spine 4 on the 

left. Vent is marked by smooth emergent lava spine. Ground on 

which MID spiders were deployed was warm and was chosen to 

reduce burial by snow. USGS photo by D. Dzurisin, January 3, 2005.

Figure 9. Graph of displacement of MID9.0 

GPS spider, indicating accelerations, 

decelerations, and reversals as this nearvent station responded to subtle changes in 

eruption dynamics.
16. Instrumentation in Remote and Dangerous Settings; Examples Using Data from GPS “Spider” Deployments 343

EASTINGS, IN METERS

-0.5 0.0 0.5 1.0 1.5 2.0 2.5

NORTHINGS, IN METERS

1.0

0.5

0

-0.5

1.5

2.0

2.5

10/28/04

12/19/04

3/1–3/8/05

5/11/05

difficulty owing to the lack of persistently identifiable features 

on the glacier surface. Those GPS spiders placed on the 

glacier provided a more continuous record of glacial deformation that allowed detailed examination of glacier compression, thickening, and increased rate of flow. Initially, growth 

of the welt and lava extrusions affected the southeastern part 

of the crater, cleaving Crater Glacier into east and west arms 

and compressing the east arm against the crater wall. Visual 

observations of this process indicated that the east arm was 

being greatly thickened. In response, the flow rate appeared to 

accelerate as a bulging lobe advanced northward toward the 

glacier’s terminus east of the old lava dome. This deformation 

was accompanied by pervasive fracturing of the glacier with 

the formation of deep crevasses and, ultimately, extensive 

fields of seracs. In order to quantify these phenomena, several 

spiders were deployed temporarily on the glacier in nine locations between February and August 2005 (table 1).

Spiders ICY4 and ICY5.0 were installed on the east arm 

of Crater Glacier on February 16, 2005. Spider ICY5.0 was 

placed on the thickened and bulging part of the glacier, and 

ICY4 was placed on the relatively undisrupted part downslope 

of the bulge. Resultant velocity measurements from ICY4 

showed the upper part of the glacier was moving 1.4 m/d 

and the lower undisturbed part was moving only 0.4 m/d. Of 

particular interest was the lack of diurnal velocity changes 

that typify temperate glaciers, indicating that Crater Glacier 

has a permeable bed that precludes basal slip resulting from 

meltwater accumulation (Walder and others, this volume, 

chap. 13). Additional spiders placed on the east and west arms 

confirmed these observations (table 1).

When dome growth shifted westward in midsummer 

2005, the west arm of Crater Glacier started to show signs of 

bulging and crevassing, so in July and August 2005 a series of 

spider deployments was begun in order to collect data necessary to track changes. Stations WES6, ELE4.4 and HIE5.1 

showed velocities of 1 m/d or more (fig. 12). Of particular 

interest was the stations’ utility as an indirect confirmation of 

continuing lava extrusion because their instruments indicated 

continued upward motion as the advancing lava compressed 

the glacier against the crater wall. In times of limited visibility 

when remote cameras were ineffective, these spiders were 

our only means of confirming continuing dome growth. Their 

utility was limited by the occasional need to move them away 

from widening crevasses in the summer and by their inability 

to function under accumulation of several meters of snow during winter months.

Conclusions

During the 2004–5 eruption of Mount St. Helens, 

portable GPS stations, nicknamed spiders, installed by 

helicopter-sling operations proved to be an invaluable volcanomonitoring tool at sites in hazardous settings or where landing 

a helicopter was not possible. With real-time telemetry of 

data, spiders transmitted day and night. Cloudy weather that 

obscured camera images had no effect on the monitoring 

capability of GPS spiders. Although seismicity has become 

the most widely used real-time tool to detect explosive eruptions, real-time deformation monitoring may be equally or 

Figure 10. Schematic map showing horizontal path of three 

sequentially placed GPS spiders as they moved from the initial 

MID site. Site was alternately pushed away or relaxed toward the 

vent area during the course of deployment. Vent lay about 200 m 

southwest of MID. Change of symbol colors is solely to clarify the 

trace of displacement path.

Figure 11. Mid-January 2005 photograph of station CDAN GPS 

spider on extruding spine 4. Legs of differing lengths permitted 

level deployment on 32° slope. Legs are embedded several 

centimeters into soft gouge that mantles spine. USGS photo by 

J.S. Pallister, January 14, 2005.
344 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

8/19 8/23 8/27 8/31 9/4 9/8 9/12 9/16

Northings

Eastings

Vertical

Northings

Eastings

Vertical

8/19 8/23 8/27 8/31 9/4 9/8 9/12 9/16

DISPLACEMENT, IN METERS

20

0

-20

-40

40

DISPLACEMENT, IN METERS

20

0

-20

-40

40

DISPLACEMENT, IN METERS

20

0

-20

-40

40

8/19 8/23 8/27 8/31

2005

9/4 9/8 9/12 9/16 

Northings

Eastings

Vertical

HIE 5.1

ELE 4.4

WES6

Figure 12. Plot of three GPS spiders on west arm of Crater 

Glacier during summer 2005, showing similar uplift and motion 

away from the advancing lava dome as the glacier was 

compressed against crater wall. Site locations shown in figure 5.

more valuable to track dome-building eruptions. At present, 

nothing demonstrates ongoing extrusion of lava as directly as 

a GPS receiver riding on an active spine. Costing about $2,500 

each in materials, spiders can be built relatively quickly and 

deployed in numbers needed to obtain data that provide a 

detailed record of near-vent deformation, lava extrusion, and 

effects on adjacent glaciers.

Acknowledgments

Our success with spiders stems in part from highly 

skilled helicopter pilots. In particular Jeff Linscott of JL Aviation and Morgan Kozloski of Hillsboro Aviation have carried 

greatest responsibility for installing and retrieving equipment. 

We gratefully acknowledge manuscript reviews by Tom Murray and Mark Reid.

References Cited

Dzurisin, D., ed., 2006, Volcano deformation—geodetic monitoring techniques: Berlin, Springer, Springer-Praxis Books 

in Geophysical Sciences, 441 p.

Dzurisin, D., Westphal, J.A., and Johnson, D.J., 1983, Eruption prediction aided by electronic tiltmeter data at Mount 

St. Helens: Science, v. 221, no. 4618, p. 1381–1383.

Endo, E.T., and Murray, T., 1991, Real-time seismic amplitude 

measurement (RSAM); a volcano monitoring tool: Bulletin 

of Volcanology, v. 53, no. 7, p. 533–545.

LaHusen, R.G., and Reid, M.E., 2000, A versatile GPS system 

for monitoring deformation of active landslides and volcanoes [abs.]: Eos (American Geophysical Union Transactions), v. 81, no. 48, p. F320.

McChesney, P.J., Couchman, M.R., Moran, S.C., Lockhart, 

A.B., Swinford, K.J., and LaHusen, R.G., 2008, Seismicmonitoring changes and the remote deployment of seismic 

stations (seismic spider) at Mount St. Helens, 2004–2005, 

chap. 7 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., 

eds., A volcano rekindled; the renewed eruption of Mount 

St. Helens, 2004–2006: U.S. Geological Survey Professional Paper 1750 (this volume).

Moran, S.C., Malone, S.D., Qamar, A.I., Thelen, W.A., 

Wright, A.K., and Caplan-Auerbach, J., 2008, Seismicity associated with renewed dome building at Mount St. 

Helens, 2004–2005, chap. 2 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Schilling, S.P., Thompson, R.A., Messerich, J.A., and Iwatsubo, E.Y., 2008, Use of digital aerophotogrammetry to 

determine rates of lava dome growth, Mount St. Helens, 

Washington, 2004–2005, chap. 8 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Schneider, D.J., Vallance, J.W., Wessels, R.L., Logan, M., 

and Ramsey, M.S., 2008, Use of thermal infrared imaging 

for monitoring renewed dome growth at Mount St. Helens, 

2004, chap. 17 of Sherrod, D.R., Scott, W.E., and Stauffer, 

P.H., eds., A volcano rekindled; the renewed eruption of 

Mount St. Helens, 2004–2006: U.S. Geological Survey 

Professional Paper 1750 (this volume).

Swanson, D.A., Casadevall, T.J., Dzurisin, D., Malone, S.D., 
16. Instrumentation in Remote and Dangerous Settings; Examples Using Data from GPS “Spider” Deployments 345

Newhall, C.G., and Weaver, C.S., 1983, Predicting eruptions 

at Mount St. Helens, June 1980 through December 1982: 

Science, v. 221, no. 4618, p. 1369–1376.

Vallance, J.W., Schneider, D.J., and Schilling, S.P., 2008, 

Growth of the 2004–2006 lava-dome complex at Mount St. 

Helens, Washington, chap. 9 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Walder, J.S., Schilling, S.P., Vallance, J.W., and LaHusen, 

R.G., 2008, Effects of lava-dome growth on the Crater 

Glacier of Mount St. Helens, Washington, chap. 13 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano 

rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).
346 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

This page intentionally left blank
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 17

1

 U.S. Geological Survey, 4200 University Drive, Anchorage, AK 99508

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

3

 Department of Geology and Planetary Science, University of Pittsburgh, 

200 SRCC Building, Pittsburgh, PA 15260

Use of Thermal Infrared Imaging for Monitoring Renewed 

Dome Growth at Mount St. Helens, 2004

By David J. Schneider1, James W. Vallance2, Rick L. Wessels1, Matthew Logan2, and Michael S. Ramsey3

Abstract

A helicopter-mounted thermal imaging radiometer documented the explosive vent-clearing and effusive phases of the 

eruption of Mount St. Helens in 2004. A gyrostabilized gimbal controlled by a crew member housed the radiometer and 

an optical video camera attached to the nose of the helicopter. Since October 1, 2004, the system has provided thermal 

and video observations of dome growth. Flights conducted 

as frequently as twice daily during the initial month of the 

eruption monitored rapid changes in the crater and 1980–86 

lava dome. Thermal monitoring decreased to several times 

per week once dome extrusion began. The thermal imaging 

system provided unique observations, including timely recognition that the early explosive phase was phreatic, location 

of structures controlling thermal emissions and active faults, 

detection of increased heat flow prior to the extrusion of lava, 

and recognition of new lava extrusion. The first spines, 1 and 

2, were hotter when they emerged (maximum temperature 

700–730°C) than subsequent spines insulated by as much as 

several meters of fault gouge. Temperature of gouge-covered 

spines was about 200°C where they emerged from the vent, 

and it decreased rapidly with distance from the vent. The 

hottest parts of these spines were as high as 500–730°C in 

fractured and broken-up regions. Such temperature variation 

needs to be accounted for in the retrieval of eruption parameters using satellite-based techniques, as such features are 

smaller than pixels in satellite images.

Introduction

Detection and measurement of elevated temperature 

can play an important role in understanding the processes 

involved with active volcanic processes. As magma moves into 

a volcano, an increase in surface heat flow may occur that is 

manifested by various phenomena, such as increases in fumarole temperatures, opening of new fumaroles, development of 

hot faults and fractures, increases in hot spring and crater lake 

temperatures, and melting of snow and ice. Once an eruption 

begins there are numerous potential heat sources such as eruption columns, lava lakes, lava domes, lava flows, and pyroclastic-flow deposits. Analysis of thermal data can provide a 

means to detect unrest, observe phenomena not discernible at 

visible wavelengths, see through thin cloud cover and steam, 

and track changes in eruption style and eruption volume.

Thermal data are acquired at a variety of spatial and 

temporal resolutions: (1) low spatial resolution (kilometer 

scale) from high temporal resolution (multiple images per day) 

satellite sensors such as GOES, AVHRR, and MODIS (Dehn 

and others, 2002; Harris and others, 1997; Wright and others, 

2002); (2) moderate spatial resolution (tens of meters) from 

low temporal resolution (weekly or greater) satellite sensors 

such as ASTER and Landsat TM/ETM+ (Flynn and others, 

2001; Ramsey and Dehn, 2004; Vaughan and others, 2005); 

and (3) high spatial resolution (meter scale) from airborne 

or ground-based thermal imaging radiometers deployed on 

a tactical basis. This study focuses on an application of the 

latter—high-resolution thermal monitoring of eruptive activity 

at Mount St. Helens during October and November 2004.

The use of thermal imaging radiometers for monitoring 

volcanic activity has increased greatly over the past 5 years. A 

new generation of instruments that has entered the commercial 

marketplace utilizes detector arrays without external cooling 

(such as liquid nitrogen), thereby simplifying field logistics 

and increasing portability. These instruments provide calibrated temperature images and are useful in such applications 
348 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

as detecting heated fractures prior to lava eruption (Andronico 

and others, 2005) or flank collapse (Calvari and others, 2005), 

estimation of lava effusion rate (Harris and others, 2005a), and 

study of eruption dynamics (Harris and others, 2005b; Patrick 

and others, 2007).

The reawakening of Mount St. Helens in September 2004 

began with a swarm of shallow earthquakes on September 

23 (Moran and others, this volume, chap. 2) and progressed 

to deformation and cracking of glacier ice on the crater floor 

south of the 1980–86 dome to form the so-called welt, which 

was first observed on September 26 (Dzurisin and others, this 

volume, chap. 14). In response to this unrest, thermal-image 

monitoring began on October 1, 2004, to establish baseline 

thermal imagery and to examine the 1980–86 dome for signs of 

deformation. This paper reports results of the thermal surveys 

conducted during the first 2 months of activity during vent 

clearing and early stages of dome extrusion from October to 

December 2004. Thermal monitoring has continued throughout 

the eruption and has aided in the analysis of subsequent dome 

growth (Vallance and others, this volume, chap. 9).

Instrumentation and Methodology 

The imaging system used in this study comprises a 

FLIR Systems ThermaCAMTM PM595 infrared camera and 

a Sony EVI-370 (visual-wavelength) video camera housed in 

a four-axis gyrostabilized gimbal that is mounted to the nose 

of a Bell-206 helicopter (fig. 1A). Although these types of 

thermal imaging cameras are generally referred to as FLIR, 

for Forward Looking Infrared, the gimbal can rotate in any 

direction. A crew member controls gimbal orientation and 

camera operation with a remote control (fig. 1B). The infrared 

camera utilizes a 320×240 microbolometer array that detects 

long-wave thermal radiation in the 7.5–13-μm atmospheric 

window and converts these values to brightness temperature. 

A 12°-wide telephoto lens provides a horizontal field of view 

of 210 m and a pixel resolution of 1.5 m at a distance of 1,000 

m. Temperature is measured in three ranges: –40–200oC, 

30–800oC, and 350–1,500oC. An internal flash memory card 

stores data from the infrared camera at 14-bit resolution and 

maximum frequency of 1 Hz. The visual video camera has a 

resolution of 720×480 pixels and a 12x auto-focus zoom lens. 

Analog video signals from the infrared and visual cameras are 

recorded simultaneously. These data are used to provide context and geographical reference for the individually calibrated 

thermal images. 

The total radiant emittance measured by the infrared 

camera is

Wtot = esWobj + (1 – e) sWrefl + (1 – s)Watm , (1)

where e is the emissivity of the object (a measure of how 

much radiation is emitted from an object compared to a perfect blackbody emitter, see discussion), s is the atmospheric 

transmission, Wobj is the emission from the object, Wrefl is the 

reflected emission from ambient sources, and Watm is the emission from the atmosphere. The camera system converts radiant 

emittance (as measured by camera detector voltage) to brightness temperature internally using factory-defined calibration 

coefficients. We use an average emissivity value of 0.96 based 

on laboratory measurements of a dacite reference (Vaughan 

and others, 2005). We use ThermaCAMTM Researcher software to calculate atmospheric transmissivity using measured 

values of the ambient air temperature, relative humidity, and 

atmospheric path length (from the object to the camera).

The telephoto capabilities of the infrared camera allowed 

for safe helicopter operations from outside the 1980 crater at 

a typical altitude of 3,000 m above sea level (600 m above 

the crater rim) and a horizontal distance of 1,000 m from the 

B

A

Figure 1. Forward Looking Infrared (FLIR) thermal and visual 

video imaging system used during airborne observations of Mount 

St. Helens, 2004–2007. A, FLIR instrument mounted on nose of 

Bell 206-L helicopter. Inset shows infrared (top) and visual video 

(bottom) lenses. B, Remote control panel used to display infrared 

and visual video streams, to control orientation of sensors using a 

joystick, and to capture data.
17. Use of Thermal Infrared Imaging for Monitoring Renewed Dome Growth at Mount St. Helens 349

welt on the south crater floor. This was especially important 

for flights conducted during the first week of October 2004 

when there was a high level of uncertainty regarding eruption 

prognosis and associated hazard. Infrared images, digital still 

photographs, and infrared and visual video were collected 

primarily at an oblique perspective, in many cases allowing for 

observation under steam and gas plumes that would have been 

impossible from a vertical viewing perspective. The thermal 

infrared and visual video collected using this system also provided striking graphical images for distribution to the media 

(Driedger and others, this volume, chap. 24).

In this study, thermal images are displayed in pseudocolor. Minimum and maximum temperature values (shown by 

a scale bar) are selected to provide optimal image enhancement. In some cases, the maximum temperature value 

reported in the text is higher than the maximum value shown 

by the scale bar.

Observations

Vent Clearing

We began thermal infrared monitoring on October 1, 

2004, to develop baseline thermal images of the 1980–86 

dome, to investigate the uplifted welt on the south crater floor, 

and to determine whether there was any thermal evidence 

for fracturing or faulting in the 1980–86 dome in response 

to ongoing uplift. Imaging began at 1156 PDT and showed 

several warm fumarolic regions with temperatures of 40oC to 

65oC, primarily on the north flank of the dome, as well as the 

influence of solar heating on the south flank. The region of the 

dome and the south crater floor adjacent to the welt (an area of 

visible deformation and faulting) was near ambient temperature (0–15oC) except for a t-shaped thermal lineament (figs. 

2A, 2B) at the west end of the graben that formed during dome 

extrusion in 1985 near the area of the dome informally known 

as Opus. This lineament had temperatures of 30oC to 38oC.

At approximately 1202 PDT an ash-rich explosion ruptured ice at the west end of the welt. This event began without 

recognizable precursory seismicity (Moran and others, this 

volume, chap. 6) and without any observable thermal change 

in images collected 15 seconds prior to onset. The eruption column comprised ballistic projectiles, jetting, ash-rich 

projections inclined northward, and a convecting steam-rich 

cloud (figs. 2C, 2D). The maximum temperature, 160oC, 

recorded from the base the column, indicated a phreatic rather 

than a magmatic mechanism. Infrared video shows that ballistic blocks were ejected primarily northward, where they 

destroyed seismic station SEP and GPS station DOM1 on the 

1980–86 dome, located about 400 m north of the vent. Within 

5 minutes of the eruption onset, a horizontally convecting ashrich cloud developed and moved westward from the vent until 

it encountered the crater wall (figs. 2E, 2F). The steam-andash cloud ultimately reached an altitude of ~4,500 m above sea 

level. Ash emissions ended abruptly at 1220 PDT, and images 

collected several minutes later once the residual steam dispersed revealed a vent with a diameter of ~20 m in ice of the 

welt and a broad area of ash deposits on the west crater floor 

(fig. 2G). Vent temperatures were only slightly above freezing 

(2 –6oC). Deposits adjacent to the vent had temperatures that 

ranged from 45oC to 50oC, and ballistic blocks with similar 

temperatures lay in small areas on the south and west flank of 

the 1980–86 dome (fig. 2H). The relatively low temperatures 

of the vent and deposits provide further evidence for a phreatic 

rather than magmatic eruption mechanism. A minor amount of 

snow and ice was melted, generating a small channelized flow 

that extended several hundred meters west from the vent. 

Rapid deformation and uplift of the welt continued following the explosion, and thermal images tracked increasing 

temperatures of certain features (figs. 3, 4). Uplift of Opus 

(south of the 1985 graben) began to expose warmer interior 

rocks primarily on its eastern margin. Images of the Opus 

area collected on October 1, 2004 (figs. 3A, 3B), show several 

areas of elevated surface temperatures that ranged from 40oC 

to 50oC. Thermal images taken the following day showed 

that this area had been uplifted further and was the source of 

nearly constant warm rockfalls. By October 3, a broad region 

of oxidized rock from the interior of the dome was exposed 

(fig. 3C), and warm rockfalls and streaks of elevated temperature were more pronounced (fig. 3D). An elongate region 

of elevated temperature, detected at the north margin of the 

welt, trended east to west along the axis of the 1985 graben 

and intersected the October 1 vent (figs. 4A, 4B). Areas of 

enhanced heat flow along normal faults appeared in images as 

shorter lineaments perpendicular to the 1985 graben (fig. 4B).

Three additional phreatic explosions on October 3, 4, 

and 5 (Moran and others this volume, chap. 6) indicated the 

presence of a shallow heat source. These events reamed out 

a small basin on the west end of the welt, wherein water 

and ice collected (fig. 5A). On October 4, small ash-rich 

geysers (less than 50 m high) erupted through the lake, after 

which hot upwelling gas bubbles with temperatures as high 

as 200oC appeared (fig. 5B). The last phreatic explosion on 

October 5 produced a new vent north of the earlier one. Over 

the next several days, new fumaroles developed near these 

vent regions, with temperatures as high as 55oC observed on 

October 7.

On October 10, after 2 days of poor weather, thermal 

imagery detected a broad steaming area of ice- and snow-free 

rock at the west end of the welt (fig. 6A). Elevated temperatures that ranged from 30oC to 50oC characterized this broad 

area, and numerous hot point sources within it ranged from 

100oC to more than 270oC (fig. 6B). This marked increase 

in temperature was the greatest yet observed and presented 

strong evidence that magma was nearing the surface.

Lava Extrusion Begins

Lava, first observed on October 11, extruded from the 

northern part of the zone of elevated temperatures shown in 
350 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

G H

E F

C D

A B

1980–86 Dome

Opus Opus

South Crater Floor

Welt Welt

1985 Graben

Ash deposit

Vent

Opus

1980– 86 Dome

60

40

20

0

60

40

20

0

60

80

40

20

0

60

80

40

20

0

TEMP (°C)

TEMP (°C)

TEMP (°C)

TEMP (°C)

Figure 2. Comparisons between visual and infrared images of Mount St. Helens collected on October 1, 2004. 

Red rectangles in visual images show approximate boundaries of contemporaneous infrared images. A, Visual 

view from southwest of south crater floor and deforming welt collected at 11:56:30 PDT. B, Infrared image of area 

shown in A. C, Visual image from east showing phreatic eruption column collected at 12:05:50 PDT. D, Infrared 

image of area in C. E, Visual image from south showing phreatic eruption column collected at 12:07:07 PDT. 

F, Infrared image of area in E. G, Visual image from southeast showing vent region of phreatic explosion and 

associated ashfall deposits. H, Infrared image of area in G.
17. Use of Thermal Infrared Imaging for Monitoring Renewed Dome Growth at Mount St. Helens 351

A B

C D

1980–86 dome

1980–86 dome

1985 graben

1985 graben

Phreatic

vent

Opus

Welt

Opus

Welt

40

20

0

TEMP (°C)

40

20

0

TEMP (°C)

A B

Opus

Opus

Welt

Phreatic

vent

1985 graben

50

40

20

0

10

30

TEMP

(°C)

Figure 3. Comparisons between visual and infrared images of southeast part of 1980–86 dome and welt at Mount St. 

Helens. View from southeast. Red rectangles in visual images show approximate boundaries of contemporaneous infrared 

images. A, Visual image collected on October 1, 2004. B, Infrared image of area in A. C, Visual image collected on October 3, 

2004. D, Infrared image of area in C.

Figure 4. Comparison between visual and infrared images collected on October 3, 2004, of southern part of 1980–86 

dome and welt at Mount St. Helens, showing thermal lineaments related to deformation. Red arrows indicate similar 

features between images. A, Visual image, view from the west. B, Contemporaneous infrared image.
352 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A B

Opus

Ephemeral

lake Venting

gas

30

10

40

20

0

TEMP (°C)

A B

Welt

South crater floor

80

40

60

20

0

TEMP (°C)

Figure 5. Comparison between visual and infrared images collected on October 4, 2004, of southern part of 1980–86 

dome and phreatic vent on welt at Mount St. Helens, showing hot gas venting through ephemeral lake. A, Visual 

image, view from west. B, Contemporaneous infrared image.

Figure 6. Comparison between visual and infrared images collected on October 10, 2004, of welt at Mount St. Helens, 

showing increased heat flow due to magma nearing surface. View from west. Red rectangle in visual image shows 

approximate boundary of contemporaneous infrared image. A, Visual image. B, Infrared image.

figure 6. Following the usage of Cascades Volcano Observatory scientists, this and all subsequent lava extrusions are 

referred to as “spines,” with a number assigned based on 

order of appearance. Spine 1 emerged as a fin-shaped slab 

approximately 80 m wide, 40 m high, and 10 m thick (fig. 

7). Although the spine was partially obscured by steam during thermal observation (figs. 7A, 7C), temperatures as high 

as 580oC were measured at its base (figs. 7B, 7D). A small 

warm debris fan extended from spine 1, but there was little 

melting evident on adjacent snow and ice. Extrusion of spine 

1 continued through at least October 14, with the hottest temperatures of the entire eruption sequence, 730oC, measured in 

cracks and fractures on that date.

Poor weather prevented thermal imaging flights from 

October 15 to 20, at which time extrusion of spine 2 had 

extended southward from spine 1 across the south crater 

floor. By October 24, spine 2 encountered the south crater 

wall (fig. 8). Spine 1 had cooled considerably since observation on October 14, with temperatures that ranged from 40oC 

to 60oC where it emerged from the ground, whereas the temperature of spine 2 ranged as high as 430oC within fractured 

and fragmented rock near the likely zone of extrusion (fig. 

8B). Intense steaming, caused by high relative humidity on 

this day (~80 percent), attenuated measured temperatures. 

The surfaces of spines 1 and 2 were rough, fractured, and 

blocky. Hot rockfalls on their west sides caused minor melting of snow and ice. An overflight on October 27 detected 

temperatures as hot as 700oC where spine 2 emerged, as 

well as a broad warm area in the debris cover southeast of 

spine 1. Observations were hampered by thick steam, but 

retrospective analysis suggests that extrusion of spine 3 had 

begun by this time.
17. Use of Thermal Infrared Imaging for Monitoring Renewed Dome Growth at Mount St. Helens 353

A B

C D

100

200

0

TEMP (°C)

100

200

0

TEMP (°C)

By November 4, spine 3 had grown considerably. Its surface was smooth with a few large cracks (fig. 9A), but its west 

and south sides were blocky and fractured. As much as several 

meters of fault gouge covering spine 3 formed as largely crystallized magma ascended the upper part of the conduit (Cashman and others, this volume, chap. 19). Its smooth surface and 

recumbent attitude gave spine 3 a whaleback-like morphology. 

The fault gouge served as an effective insulator so that surface 

temperatures where it emerged from the vent on its north end 

ranged from 150oC to 200oC and decreased exponentially to 

typical values of 25oC to 50oC within 50 to 75 m (fig. 9B). The 

hottest temperature observed in small regions of collapse and 

fractures was 625oC.

Spine 3 moved toward the south as it was extruded and 

encountered the south crater wall by November 12 (Vallance 

and others, this volume, chap. 9). By late November, continued southward compression against the crater wall resulted 

in crumbling of spine 3 at its south end (fig 10A). Large open 

fractures in the fault-gouge carapace observed in thermal 

images (fig. 10B) suggested tensional cracking caused by bending of the solid spine as the extrusion continued southward. A 

typical temperature profile during this time period is shown in 

Figure 7. Comparison between visual and infrared images of Mount St. Helens collected on October 11, 2004, showing the 

first views of lava at the surface. Red rectangle in visual image shows approximate boundary of contemporaneous infrared 

image. A, Visual image, view from southwest. B, Infrared image, view from southwest. C, Visual image, view from northwest. 

D, Infrared image, view from northwest.

figure 11 along a line parallel to direction of extrusion. Temperature at the vent was typically 200oC and decreased abruptly 

southward, but was as high as 700oC in fractures.

Thermal infrared images continue to be used as a monitoring tool to observe dome growth and have documented the 

thermal characteristics of recumbent (whaleback) extrusion 

from November 2004 to April 2005, vertical spine extrusion 

from April to July 2005, and composite spine growth (composed of massive rock and disintegrating rock debris) from 

July 2005 to present (April 2007) (Vallance and others, this 

volume, chap. 9).

Discussion

Sources of Temperature Measurement Error

Accurate measurement of temperature using a thermal 

radiometer depends on several factors including viewing 

angle, characteristics of the atmospheric path, absorption 

by volcanic gases (water and SO2), and the emissivity of the 
354 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

1

2

B

2

1 200

250

150

100

50

0

TEMP (°C)

A B

1

3

3 1

200

100

0

TEMP

(°C)

Figure 9. Comparison between visual and infrared images of Mount St. Helens collected on November 4, 2004, showing the 

development of spines 1 and 3 (as labeled). A, Visual image, view from the northeast. B, Composite infrared image.

Figure 8. Comparison between visual and infrared images of 

Mount St. Helens collected on October 24, 2004, showing the 

development of spines 1 and 2 (as labeled). A, Visual image, view 

from the west. B, Composite infrared image.

target (Ball and Pinkerton, 2006). The maximum temperature 

values reported herein are most strongly affected by viewing 

angle, because the source of heat typically was a fracture or 

fissure. In these cases, physical blockage of the source(s) of 

thermal emission occurred, and imaging of the target from 

multiple viewing angles was required to detect the maximum values. In other cases, opaque steam clouds completely 

obscured the source, making it impossible to measure the 

temperature accurately.

Corrections for atmospheric-path absorption were made 

using ThermaCAMTM Researcher software and estimates of 

atmospheric temperature, relative humidity, and path length. 

For each mission, temperature and humidity values for Mount 

St. Helens were estimated using archived EDAS40 model 

sounding data provided by the NOAA Air Resources Laboratory (http//www.arl.noaa.gov/ready/amet.html). These values 

were used to compute the atmospheric transmission; the relation between these parameters is shown in figure 12 for a typical range of values. Uncertainty in estimating a representative 

relative humidity will contribute to a larger error in calculating 

transmission as the temperature of the atmosphere increases; 

the error will be compounded as the path length increases. 

Many of the thermal imaging missions were conducted at 

atmospheric temperatures near 0oC, the relative humidity 

values varied by a maximum of ±15 percent, and variation in 

path length was ±250 m. This results in a variation in atmospheric transmission of about ±2 percent. For an object with a 

temperature of 500oC, this results in a variation of ±10oC.

Although water is the dominant volcanic gas that absorbs 

infrared radiation, SO2

 has two strong absorption features 

(centered at wavelengths of 7.3 and 8.5 μm) within the spectral 

range of the thermal radiometer. Although techniques to 

quantify the systematic error in ground-based thermal infrared measurements caused by gas absorption are still being 

developed, it may be a significant issue in situations where gas 

concentrations are high (Sawyer and others, 2005). The SO2

emission rates during the time period of this study were low, 

ranging from values at the lower detection limit of 1 metric ton 

per day (t/d) through October 4 and increasing to an average 

value of approximately 100 t/d thereafter (Gerlach and others, 

this volume, chap. 26). Thus, absorption of infrared emissions 

by SO2 is likely to be a relatively small source of error.
17. Use of Thermal Infrared Imaging for Monitoring Renewed Dome Growth at Mount St. Helens 355

A B

3

3

2 2 1

1

350

-15

TEMP

(°C)

0 50 100 150 200

A

B

A A'

A

A'

TEMPERATURE, IN

DEGREES CELSIUS

PIXEL DISTANCE

400

300

200

100

0

350

0

TEMP (°C)

Figure 10. Comparison between visual and infrared images of Mount St. Helens collected on November 29, 2004, showing 

the development of spines 1 to 3. Spine 3 (as labeled) is actively being extruded and is crumbling as it encounters the south 

crater wall. A, Visual image, view from the west. B, Composite infrared image.

Figure 11. Temperature profile 

along north end of spine 3 at 

Mount St. Helens on November 

29, 2004, showing typical 

temperatures observed during 

extrusion of whaleback spines. 

A, Temperature profile A–A’; 

profile length approximately 125 

m, north toward left. B, Infrared 

image of spine 3 viewed from the 

north showing the location of the 

north-south profile, A–A’.

Emissivity is a measure of how much energy is emitted from an object compared to a theoretical perfect emitter 

known as a blackbody. In nature, all materials emit less energy 

than a blackbody, because energy is absorbed by the vibration 

of molecular bonds. They are referred to as selective radiators wherein emissivity varies as a function of wavelength. 

Although the emissivity of silicate rocks has a characteristic 

decrease in emissivity at a wavelength range of 8 to 10 μm, an 

average value over the broader range of 7.5 to 13 μm is used 

where measuring temperatures with a broadband radiometer. 

Variations in emissivity can also occur for a given rock type as 

a result of vesicularity and particle size. In general, emissivity 

decreases as vesicularity decreases (Ramsey and Fink, 1999). 

The 2004 lava was largely degassed and dense (Pallister and 

others, this volume, chap. 30), suggesting that the actual 

emissivity may be lower than a laboratory reference. Furthermore, very fine grained silicates (<60 μm) exhibit a decrease in 

emissivity of the primary absorption band between 8 and 10 μm 

owing to the scattering of radiant energy (Ramsey and Christensen, 1998). It is possible that ash fallout and (or) the cover of 

the chalky fault-gouge carapace served to reduce the effective 

emissivity of the extruded lava (compared to the dacite reference value). The effect of an overestimated emissivity would be 

an increase in the temperatures above those reported herein. The 

result of a variation in emissivity was tested by changing the 

value by a factor of 0.1 within the ThermaCAMTM Researcher 

software, which resulted in a change in temperature on the order 

of 20oC for an area with a temperature of 450oC. Thus, variations in emissivity may account for changes on the order of 5 

percent in the absolute temperature recorded.
356 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

15 25 35 45 55 65 75

RELATIVE HUMIDITY, IN PERCENT

ATMOSPHERIC TRANSMISSION

1,500 m; 0°C

500 m; 15°C

1,000 m; 15°C

500 m; 0°C

EXPLANATION

1,000 m; 0°C

1,500 m; 15°C

1.0

0.9

0.8

0.7

Figure 12. Atmospheric transmission as a function of relative humidity, atmospheric temperature, and path length.

Thermal Imaging as a Response Tool

The rapid pace of unrest associated with the reawakening 

of Mount St. Helens in September 2004 resulted in an intense 

monitoring response by the Cascades Volcano Observatory 

(CVO) and the broader U.S. Geological Survey Volcano 

Hazards Program (VHP). Although handheld thermal imaging cameras have been used successfully in eruption response 

by other volcano observatories worldwide (Kauahikaua and 

others, 2003; Andronico and others, 2005; Calvari and others, 

2005; Harris and others, 2005a), the gimbal-mounted system used during this response was a relatively new tool for 

VHP scientists. It had been flown on several missions at the 

Alaska Volcano Observatory (AVO) to collect baseline thermal 

data and on one flight in response to unrest at Mount Spurr 

(Power, 2004). Thus, the gimbal-mounted instrument was 

largely untested for eruption response by the USGS when first 

flown at Mount St. Helens on October 1, 2004, arriving only 

6 minutes prior to the initial phreatic explosion. The utility of 

the instrument was quickly recognized as it was determined 

rapidly that the initial explosion was of phreatic origin. This 

information was quickly radioed to the operations center at 

CVO and used to apprise land management and emergency 

response personnel. In addition, the thermal and visual video 

imagery was made available to the media and was broadcast 

widely. The ability to control the gimbal and shoot video in a 

wide range of orientations was extremely useful as the helicopter maneuvered in response to the eruption cloud, topographic barriers, and wind. The telephoto capabilities were 

also vital, allowing for observation at a safe distance beyond 

the crater rim.

In addition to documenting thermal features on the crater 

floor and dome, use of the instrument during the first 2 weeks of 

October provided opportunities to test thermal-observation strategies and to determine what kinds of useful information could 

be obtained. The stability of the 1980–86 dome in response to 

the deformation on the south crater floor was initially of great 

concern, as potential eruption scenarios were evaluated involving (1) large-scale collapse of the 1980s dome toward the northern open end of the 1980 crater and (2) potential hazards to the 

public who had come to view the volcano. Thus, an important 

early use of the system was to look for the thermal expression of 

faults or fractures on the northern flank of the dome. Observations made throughout the eruption have shown no significant 

change in thermal response of the 1980–86 dome and thus have 

eased concerns about its stability.

Structural Control of Heat Emissions

The south flank of the 1980–86 dome, which formed the 

northern part of the welt, showed several examples of structural control of heat emissions during a time of rapid uplift, 

faulting, and deformation (figs. 3, 4) along a graben formed 

during an episode of dome growth in 1985. Between flights 

on October 1 and 2, rapid uplift of tens of meters on the east 

end of the graben caused warm rockfalls from the newly 

exposed interior of the dome. Over the next several days, the 

thermal imaging system was able to record continued uplift in 

this region as expressed by displacement of thermal features 

and nearly continuous warm rockfalls. On the west end of 

the graben (fig. 4), the thermal expression of normal faulting was observed, suggesting that this area was not moving 

upward as rapidly as the east end. Although these features 

can be seen in visible images, the advantage of the thermal 

system was to highlight areas of interest during a time period 

of very rapid changes in surface topography. Similar phenomena (on a larger scale) were observed at Mount St. Helens in 

April and May 1980 during development of the bulge. For 

example, Kieffer and others (1981) documented increased 

heat flow from several faults, fractures, and brecciated zones 

using a variety of thermal imaging devices, and they reported 
17. Use of Thermal Infrared Imaging for Monitoring Renewed Dome Growth at Mount St. Helens 357

temperature values in these regions of 15oC to >25oC above 

background, similar to those observed in October 2004. 

Structural control of heat emissions have also been noted at 

Mt. Etna in 2002 where hot fractures were observed several 

weeks prior to the eruption of a lava flow (Andronico and 

others, 2005), at Stromboli in 2002 where fractures were 

observed several hours prior to a major slope failure (Calvari 

and others, 2005), and at Augustine Volcano in 2006 where 

hot fractures and fissures were observed in the weeks preceding vulcanian explosions (D. Schneider, unpub. data). These 

studies suggest the broad utility of thermal imaging systems 

for detecting and monitoring deformation in the weeks and 

months prior to eruption.

The thermal measurements at Mount St. Helens documented an increase in heat flow on October 10 as magma 

neared the surface. On October 11, they provided unambiguous confirmation that lava (rather than uplifted crater floor) 

was being extruded. The imaging capabilities provided better 

viewing conditions at times when steam otherwise obscured 

observation (fig. 7) and allowed for an estimate of the volume 

of spine 1. The maximum temperature observed in fractures 

and regions of collapse during the eruption sequence was 

730oC, well below the solidus temperature of dacite. Although 

the thermal measurements may underestimate the maximum 

temperature as a result of thermal stratification due to radiative 

cooling, the observations are consistent with the petrologic 

(Pallister and others, this volume, chap. 30; Cashman and others, this volume, chap. 19) and seismic observations (Moran 

and others, this volume, chap. 2) that suggest a viscous to solid 

transition at a depth of approximately 1 km.

Comparisons to Satellite Methods

The detection limit of satellite-based measurements 

of thermal emissions is largely a matter of spatial resolution. With the exception of large thermal features (such 

as lava lakes, crater lakes, and pyroclastic-flow deposits), 

most sources of volcanogenic heat are subpixel in size when 

viewed with moderate-resolution sensors such as ASTER 

(90-m pixels) or coarse-resolution sensors such as AVHRR 

or MODIS (1-km pixels). Thus, the signal detected by the 

satellite is a pixel-integrated value of (at least) a hot and a cold 

(background) component. In order for a thermal anomaly to 

be detected, the proportion of hot material within a pixel, or 

the temperature of the hot component, needs to increase to 

the point where the integrated temperature exceeds a threshold (compared to neighboring pixels). So it follows that the 

smaller the spatial resolution of a sensor, the more sensitive it 

will be to detecting anomalous pixels.

Vaughan and Hook (2006) examined nighttime (to eliminate the solar reflected radiance) ASTER thermal infrared 

data for Mount St. Helens and reported that a single-pixel, 

weak (10°C) thermal feature was observed on the evening 

of October 1 approximately 11 hours after the first phreatic 

explosion. They noted that this value was within the range 

of typical ASTER-derived temperatures observed over the 

previous 5 years and would likely have not been viewed as 

significant without knowledge of the explosion. The next 

clear nighttime ASTER image was collected in February 

2005 and showed an unambiguous maximum temperature of 

about 50°C. Although this instrument has been used successfully to monitor eruptive activity at many other volcanoes, the 

long period of time between successful, cloud-free nighttime 

images illustrates that significant eruptive activity can occur 

without being observed.

Manual analyses conducted by AVO of mid-infrared data 

from the AVHRR and MODIS sensors (1-km spatial resolution with multiple image collections per day) were not able 

to detect any unusual thermal (or volcanic ash) emissions 

during the precursory and vent-clearing phases of the eruption. 

The first weak thermal anomaly of 10°C (above background) 

was observed in data collected on the evening of October 11, 

shortly after the first observations of lava extrusion using the 

thermal imaging system (Alaska Volcano Observatory, unpub. 

data). It is unclear whether this level of activity would have 

been observed or reported without the knowledge of lava 

being extruded, as thermal anomalies of this magnitude can 

be masked by noise or other environmental factors. Thermal 

anomalies of 10°C to 20°C above background were reported 

by AVO throughout most of October during times of clear 

weather and increased to a one-pixel anomaly of 32°C above 

background in nighttime data on November 8 (Alaska Volcano Observatory, unpub. data). This observation corresponds 

most closely with the thermal image shown in figure 9 and 

occurred about 2 weeks after the start of the extrusion of 

spine 3. By comparison, MODVOLC, an automated global 

system designed to detect volcanic activity (with a rather high 

threshold to reduce false alarms) (Wright and others, 2004), 

did not trigger until December 18. By this time, spine 3 had 

encountered the south crater wall and started to disintegrate. 

Undoubtedly the detection of thermal activity by satellite 

techniques was greatly reduced by the insulating properties of 

the relatively cool fault-gouge carapace. Once the spines grew 

large and began to break apart, detection was more apparent. 

Still, lava extrusion was ongoing for a least a month before 

satellite methods utilizing 1-km-resolution data were able to 

unambiguously detect it.

Conclusions

The use of airborne thermal imaging to document eruptive activity at Mount St. Helens began in October 2004 to 

develop baseline thermal images of the 1980–86 dome and 

to investigate the rising welt on the south crater floor. Since 

that time, it has become a routine monitoring tool for tracking 

evolution of the lava dome (Vallance and others, this volume, 

chap. 9). Some of the unique information that has been derived 

from these data to date include (1) rapid identification of the 

phreatic nature of the early vent-clearing explosive phase; (2) 
358 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

observation of the structural control of heat flow during times 

of large-scale deformation of the south flank of the 1980–86 

dome, which was part of the welt; (3) observations of the venting of hot gas through an ephemeral crater lake and increased 

heat flow through the welt, indicative of a shallow magma 

source; (4) confirmation of new lava reaching the surface; (5) 

characterization of the change in surface temperature of the 

lava spines, from the initial, hotter spines 1 and 2 that were 

free of fault gouge, to the cooler gouge-covered whaleback 

spines (beginning with spine 3); (6) documentation that the 

relatively low maximum temperature of 730°C agreed with 

petrologic and seismic data suggesting a viscous-to-solid 

transition at shallow depth; and (7) identification of the source 

of active lava extrusion, dome collapse, and large variation 

in surface temperatures. Such temperature variation needs to 

be accounted for in the retrieval of eruption parameters using 

satellite-based techniques, because the thermal features are of 

subpixel size in satellite images.

Acknowledgments

We would like to thank Matthew Patrick and Cynthia 

Werner, who provided thoughtful and helpful reviews, and 

helicopter pilot Jeff Linscott of JL Aviation for his skillful, 

safe, and enthusiastic flying.

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360 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 18

1

 U.S. Geological Survey, PO Box 51, Hawaii National Park, HI 96718

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Radar Interferometry Observations of Surface 

Displacements During Pre- and Coeruptive Periods at 

Mount St. Helens, Washington, 1992–2005

By Michael P. Poland1 and Zhong Lu2

Abstract

We analyzed hundreds of interferograms of Mount St. 

Helens produced from radar images acquired by the ERS-1/2, 

ENVISAT, and RADARSAT satellites during the 1992–2004 

preeruptive and 2004–2005 coeruptive periods for signs of 

deformation associated with magmatic activity at depth. Individual interferograms were often contaminated by atmospheric 

delay anomalies; therefore, we employed stacking to amplify 

any deformation patterns that might exist while minimizing 

random noise. Preeruptive interferograms show no signs of 

volcanowide deformation between 1992 and the onset of eruptive activity in 2004. Several patches of subsidence in the 1980 

debris-avalanche deposit were identified, however, and are 

thought to be caused by viscoelastic relaxation of loosely consolidated substrate, consolidation of water-saturated sediment, 

or melting of buried ice. Coeruptive interferometric stacks are 

dominated by atmospheric noise, probably because individual 

interferograms span only short time intervals in 2004 and 

2005. Nevertheless, we are confident that at least one of the 

seven coeruptive stacks we constructed is reliable at about the 

1-cm level. This stack suggests deflation of Mount St. Helens 

driven by contraction of a source beneath the volcano.

Introduction

Continuous Global Positioning System (GPS) measurements at Mount St. Helens have provided unequivocal 

evidence of coeruptive deflation that started in 2004, coincident with the onset of seismicity in late September of that 

year (Lisowski and others, this volume, chap. 15). These data 

have been used to model an ellipsoidal source of volume loss 

between 6 and 12 km depth beneath the volcano (Lisowski and 

others, this volume, chap. 15; Mastin and others, this volume, 

chap. 22). In addition, GPS results from in and around the 

crater, including single-frequency instruments deployed on 

the growing lava dome, have recorded surface displacements 

associated with dome extrusion (LaHusen and others, this 

volume, chap. 16). Interferometric synthetic aperture radar 

(InSAR) data provide an excellent complement to these GPS 

data. Though lacking the high temporal resolution of continuous GPS or tiltmeter data, InSAR has the potential to provide 

much greater spatial resolution of deformation, which is 

useful both for identifying localized deformation sources and 

modeling complex source geometries. Further, radar images of 

Mount St. Helens are available from 1992 onward, making it 

possible to quantify surface displacements during the 1992–

2004 preeruption period, when few other geodetic measurements were collected.

Although InSAR characterizations of surface deformation at shield volcanoes (for example, Amelung and others, 

2000a) and calderas (for example, Hooper and others, 2004; 

Wicks and others, 1998; Wicks and others, 2006) have been 

quite successful, studies of stratovolcanoes have proven more 

challenging. This is largely due to the fact that most stratovolcanoes are steep sided and covered to varying degrees 

by vegetation, snow, and ice, leading to a loss of coherent 

interferometric signal on and around the volcano, especially 

for C-band radar wavelengths (Zebker and others, 2000; Lu 

and others, 2005b; Moran and others, 2006). The great height 

of many stratovolcanoes, relative to the surrounding terrain, 

can also create atmospheric conditions that introduce significant artifacts into InSAR results (Beaudecel and others, 

2000; Wadge and others, 2006). The combined effects of these 

conditions prevented the use of coeruptive InSAR data for the 

analysis of deformation at numerous stratovolcanoes, includ-
362 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ing Galeras, Colombia; Rincon de la Vieja, Costa Rica; Unzen, 

Japan; and Merapi, Indonesia (Zebker and others, 2000). 

Even when coeruptive data for a stratovolcano are coherent, 

there is no guarantee that deformation will be observed, as 

demonstrated by results from Fuego and Pacaya, Guatemala; 

Popocatepetl, Mexico; Sakurajima, Japan (Zebker and others, 

2000); Irruputuncu ( Zebker and others, 2000; Pritchard and 

Simons, 2004b) and Lascar (Pritchard and Simons, 2002; 

Pritchard and Simons, 2004a), Chile; and Shishaldin (Moran 

and others, 2006), Pavlof, Cleveland, and Korovin, Alaska 

(Lu and others, 2003b). Lu and others (2003b) and Moran and 

others (2006) suggested that, at least in the case of Shishaldin, 

the lack of observed line-of-sight displacements in coeruptive 

InSAR data may be caused by (1) posteruption motion that 

balanced preeruption displacements, resulting in no net deformation across the time spanned by the interferograms, (2) the 

occurrence of deformation in “blind zones” where InSAR data 

are not coherent, or (3) the lack of any significant deformation 

associated with the eruption. Pritchard and Simons (2004a; 

2004b) suggest similar reasons for the lack of observed deformation in InSAR data from Cerro Irruputuncu and Volcán 

Lascar, two volcanoes in Chile, in addition to the possibility 

that the magma reservoir may be too deep to produce surface 

displacements detectable by InSAR.

Poor coherence at a volcano does not necessarily imply 

that surface displacements cannot be observed by InSAR. 

Cerro Hudson is a 10-km-diameter, ice-filled caldera in 

southern Chile. Despite a complete lack of coherence within 

and limited coherence outside the caldera, displacements 

were of sufficient magnitude (15-cm line-of-sight inflation over 3 years) to be detected and distinguished from 

atmospheric noise (Pritchard and Simons, 2004b). Although 

large displacements were not observed in individual coeruptive interferograms of Soufrière Hills volcano1

, Montserrat, 

Wadge and others (2006) identified deformation caused by 

a variety of processes once they stacked (averaged) multiple 

interferograms. At stratovolcanoes where coherence is excellent, InSAR studies have resulted in important discoveries of 

aseismic magma accumulation, including Peulik, Alaska (Lu 

and others, 2002b) and South Sister, Oregon (Dzurisin and 

others, 2006; Wicks and others, 2002). In addition, InSAR 

has been critical for characterizing deformation at many 

composite volcanoes that have experienced unrest or eruptions, including Westdahl (Lu and others, 2003a), Makushin 

(Lu and others, 2002a), Akutan ( Lu and others, 2000; Lu 

and others, 2005b), Augustine (Masterlark and others, 2006), 

and Seguam (Masterlark and Lu, 2004), Alaska; Gada ‘Ale, 

Ethiopia (Amelung and others, 2000b); and several volcanoes 

in the central and southern Andes (Pritchard and Simons, 

2002; 2004a; 2004b). Clearly, the potential for advancing the 

understanding of pre- and coeruptive activity at Mount St. 

Helens through InSAR studies is substantial.

We completed an exhaustive analysis of InSAR data collected from Mount St. Helens between 1992 and 2005 from 

a variety of satellites, ground tracks, time spans, and satellite 

viewing geometries. Our goals were to (1) obtain high-spatialresolution coeruptive surface displacement data of the known 

volcanowide deflation (Lisowski and others, this volume, chap. 

15) for input into deformation source models, (2) identify any 

localized displacements that might not be detectable by GPS or 

other terrestrial measurements, and (3) identify any displacements precursory to the onset of eruptive activity in 2004. 

Coherence was poor and atmospheric artifacts were significant 

in most interferograms; nevertheless, we recognized localized displacements unrelated to volcanic activity and modeled 

coeruptive subsidence (although poor coherence prevented the 

use of complex model geometries). Preeruptive interferograms 

showed no evidence of precursory deformation; however, 

displacements localized within the crater cannot be ruled out, 

owing to poor coherence and atmospheric distortions.

Methodology

Traditional InSAR studies combine two radar images 

of the same area on the ground acquired at different times 

from nearly the same point in space to determine the deformation over the time spanned along the radar’s line-of-sight 

(LOS). This procedure, unfortunately, does not produce 

satisfactory results at Mount St. Helens for three reasons: 

(1) atmospheric path delays, which often correlate with 

topography, can introduce artifacts that amount to several 

centimeters of apparent LOS displacement (fig. 1), (2) interferograms that include scenes acquired during nonsummer 

months are generally incoherent due to seasonal snow cover, 

and (3) coherence breaks down quickly over time, causing 

interferograms that span more than one or two years to be 

mostly incoherent. In fact, in interferograms spanning several 

months or more, coherence was generally maintained only on 

deposits that were emplaced during the explosive eruptions 

of 1980 (for example, May 18, May 25, June 12, July 22), 

especially north of the volcano and in the upper North Fork 

Toutle River. As a result, we relied on interferometric stacks 

to study deformation of Mount St. Helens.

Stacking is a procedure that adds the LOS displacements from multiple interferograms (which may or may not 

be overlapping in time) acquired along the same look angle. 

Dividing the summed displacements by the cumulative time 

spanned results in an average displacement rate (for example, 

Fialko and Simons, 2001; Peltzer and others, 2001). The 

technique decreases the magnitude of random noise (primarily atmospheric artifacts) by N/u0.5, where N is the error in a 

single interferogram and u is the number of interferograms 

in the stack (Bevington and Robinson, 1992, p. 39–40). For 

1

 Capitalization of “Volcano” indicates adoption of the word as part of 

the formal geographic name by the host country, as listed in the Geographic 

Names Information System, a database maintained by the U.S. Board on Geographic Names. Noncapitalized “volcano” is applied informally—eds.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 363

122 ̊18' 122 ̊14' 122 ̊10' 122 ̊06'

46 ̊08'

46 ̊12'

46 ̊16'

5 km

Spirit Lake

-4 4 0 8

LOS displacement (mm)

AREA

OF MAP

WASH.

Figure 1. Interferogram of Mount St. Helens, Washington, formed 

from ERS-1 orbit 21415 (August 19, 1995) and ERS-2 orbit 2243 

(September 24, 1995). Interferogram spans only 36 days during summer 

months, hence high level of coherence. Phase correlates strongly with 

topography, so Mount St. Helens shows apparent line-of-sight (LOS) 

lengthening, while lower valleys to east show LOS shortening. Such 

behavior is characteristic of strong atmospheric influence. Topographic 

irregularities caused by areas of no data in SRTM DEM.

example, if atmospheric variations typically result in an error 

of 10 mm/yr, a stack of 25 interferograms would reduce the 

noise level to about 2 mm/yr (assuming atmospheric path 

delay errors have similar magnitudes in all interferograms). 

In contrast, signals that persist over time—including steady 

deformation due to volcanic or tectonic activity—are emphasized because the signal is manifested in all interferograms in 

the stack. InSAR stacks have been able to distinguish displacement rates as low as a few millimeters per year (Wright 

and others, 2001, 2004) but are useful only when the strength 

and geometry of the deformation source do not vary significantly over time, meaning that nonlinear deformation rates or 

changes in source location or geometry are not easily recovered in interferometric stacks.

We do not expect changes in source geometry over time 

to bias the interferometric stacks. The geometry of the magmatic system of Mount St. Helens has been investigated using 

both petrologic (for example, Rutherford and others, 1985; 

Cashman, 1988; Cashman, 1992; Pallister and others, 1992) 

and seismic (Scandone and Malone, 1985; Barker and Malone, 

1991; Lees, 1992; Moran, 1994; Musumeci and others, 2002) 

techniques and includes a reservoir at around 6–10 km depth 

connected to the surface by a near-vertical conduit. The conduit appears to have been sealed at a depth of about 2 km following the end of the 1980–86 eruptive period (Moran, 1994). 

Any pre- or coeruptive deformation associated with unrest 

beginning in 2004 will likely have a source within this system 

(especially at the levels around the 6–10-km-depth reservoir or 

2-km-depth seal).

Similarly, deformation rates during both the pre- and 

coeruptive periods appear to be approximately linear, so 

time-variable displacements will not bias deformation rates 

derived by stacking. Electronic Distance Measurement 

(EDM) and campaign GPS measurements in the inter-eruptive period, from late 1986 through the first half of 2004, 

indicate that little deformation occurred, especially after 

1991 (Lisowski and others, this volume, chap. 15). In addition, individual interferograms that span time periods before 
364 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

the start of the unrest in September 2004 show no unambiguous signs of volcanowide deformation. If preeruptive surface 

displacements occurred, the motion must therefore have been 

small in magnitude but may have accumulated over time. 

Coeruptive displacements, measured by continuous GPS 

starting on September 23, 2004, are generally linear after 

the first few weeks of the activity (Lisowski and others, this 

volume, chap. 15) and suggest only minor temporal or spatial 

variations in the contraction of a source at depth beneath 

Mount St. Helens.

An unusually large InSAR data set, summarized in 

table 1, is available for Mount St. Helens, including multiple 

satellites with a variety of imaging geometries. For preeruptive time periods, we analyzed data from ERS-1, ERS-2, 

and ENVISAT. Usable data from the first two satellites span 

1992–2001 and are available for four orbital tracks, whereas 

a sufficiently large archive of ENVISAT data is available 

from one track to allow an assessment of the surface displacement field in 2003–4. Our coeruptive InSAR data set 

includes five ENVISAT and two RADARSAT tracks. All 

possible interferograms from each independent track with 

perpendicular baselines of less than 300 m were created.

Topographic corrections to interferograms utilized the 

30-m resolution Shuttle Radar Topography Mission (SRTM) 

digital elevation model (DEM), which was generated in 2000 

(Farr and Kobrick, 2000). The older National Elevation Dataset (NED) 30-m DEM, created before 1986, contains fewer 

holes than the SRTM data, but its use resulted in deformation 

artifacts in interferograms in the area of Mount St. Helens 

crater. Significant topographic changes have occurred in the 

crater since the end of the 1980–86 eruptive period, most 

notably the growth of a glacier that had reached a maximum 

thickness of 200 m by the time eruptive activity resumed in 

2004 (Schilling and others, 2004; Walder and others, this 

volume, chap. 13). The more recent SRTM DEM accounts 

for growth of the glacier and other changes since 1986 and 

does not introduce significant artifacts related to topographic 

change. Interferograms were smoothed using the filtering 

strategy of Goldstein and Werner (1998) and unwrapped following Chen and Zebker (2001).

From this collection of interferograms, we omitted all 

images that spanned less than 100 days (except for coeruptive ENVISAT mode 2, track 385, where the only interferograms with perpendicular baselines less than 300 m spanned 

70 days or less), because images that span such short times 

have a low signal-to-noise ratio given the low displacement 

rates known from GPS (Lisowski and others, this volume, 

chap. 15). There is no upper limit on the time spanned, 

although interferograms that showed no or very limited 

coherence, including nearly all images that span more than 

two years, were omitted from the analysis. Stacks of each 

independent track were computed from this modified data 

set, and only pixels that are coherent in 60 percent of the 

input images are included in the final stack. Five preeruption 

and seven coeruption stacks were created, with 3 to 38 input 

interferograms each (table 1).

Preeruptive Stacks

Results

In preeruptive stacks (figs. 2–6), the range of displacements for each track is generally small, suggesting that the 

random atmospheric signal has been mostly removed. The lack 

of significant topography-correlated phase at Mount Adams 

(which we assume to be undeforming, based on the absence of 

other signs of unrest) in tracks that both cover and are coherent around both volcanoes (figs. 4, 5) further supports our 

assertion that atmospheric noise has been mostly suppressed 

by stacking.

None of the preeruption stacks show any signs of volcanowide deformation, suggesting that no significant surface 

displacements accumulated over the times spanned by the 

stacks. Unfortunately there are no InSAR results available 

between late 2001 and early 2003, so we cannot rule out 

deformation during this time period. No volcanowide displacements are apparent in the one stack that covers the early 

2003 to September 2004 epoch (up to the start of the eruption), suggesting that no centimeter-level displacements on 

the flanks of the volcano occurred in the months immediately 

preceding the start of the eruption.

Despite this lack of signal, the five preeruption stacks 

do reveal at least three localized areas of surface displacement that are unrelated to the magmatic system of Mount St. 

Helens. All three regions are patches of subsidence located 

on the May 18, 1980, debris-avalanche deposit (fig. 2). The 

area with the highest subsidence rate, as much as 15 mm/yr 

(LOS), is directly north of the crater (site 1 in fig. 2), near the 

base of Johnston Ridge, and had been recognized previously in 

individual interferograms by Diefenbach and Poland (2003). 

A second, less distinct area located near the outlet of Coldwater Lake (site 2 in fig. 2) is subsiding at a rate of slightly less 

than 5 mm/yr (LOS). The third patch of subsidence, which 

is sinking at an LOS rate of approximately 6 mm/yr, is in the 

North Fork Toutle River valley, just southeast of Elk Rock and 

immediately upstream from a valley constriction (site 3 in fig. 

2). The latter two patches had not been recognized previously.

The stacks presented in figures 2 through 6 give average 

deformation rates. To assess whether or not the subsidence 

rates of the three localized patches changed during 1992–

2005, we used the individual unwrapped interferograms to 

reconstruct the temporal evolution of the deformation following the method of Berardino and others (2002). At Mount St. 

Helens, all five preeruption stacks suggest a constant subsidence rate of the area near the base of Johnston Ridge (fig. 7A). 

The four ERS-1/2 tracks are directly comparable in these time 

series, because they have similar look angles (table 1), and we 

assume that the deformation is dominated by vertical motion. 

In addition, the ENVISAT mode 2, track 156, time series confirms that the subsidence was continuing at least into 2004 at 

a similar rate. Subsidence near the mouth of Coldwater Lake 

(fig. 7B) and below Elk Rock (fig. 7C) may have decayed 
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 365

Satellite Mode Track A or D Incidence 

angle (deg)

Years 

spanned

Number of 

igrams

Figure 

No.

Preeruptive SAR data

ERS-1/2 - 156 D 23.2 1992–2001 36 2

ERS-1/2 - 163 A 23.2 1993–2000 7 3

ERS-1/2 - 385 D 23.2 1992–2001 38 4

ERS-1/2 - 392 A 23.2 1995–2000 7 5

ENVISAT IS 2 156 D 22.8 2003–2004 6 6

Coeruptive SAR data

ENVISAT IS 2 156 D 22.8 2004–2005 7 8

ENVISAT IS 2 163 A 22.8 2004–2005 13 9

ENVISAT IS 2 385 D 22.8 2004–2005 31 10

ENVISAT IS 2 392 A 22.8 2004–2005 6 11

ENVISAT IS 6 20 A 40.9 2004–2005 13 12

RADARSAT S 2 - A 27.6 2004–2005 9 13

RADARSAT S 5 - A 39.1 2004–2005 6 14

1

 Note that all three input interferograms span 70 days or less.

Table 1. Characteristics of SAR data from Mount St. Helens, Washington, used in stacks and 

corresponding figures.

[Mode not shown for ERS-1/2 satellite data, which have only one mode. RADARSAT does not use track numbers. 

A, ascending; D, descending. Incidence angle is measured in degrees from vertical. Number of interferograms (igrams) 

indicates those used to construct the stack. Figure numbers correspond to this chapter.]

-15 -10 -5 0 5 10

LOS displacement (mm/yr)

Elk Rock

1

2

3

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

AREA

OF MAP

WASH.

OREG.

5 km

Spirit Lake

North Fork Toutle River

Johnston

Ridge

S. Fk. Toutle River

Coldwater L.

Figure 2. ERS-1/2, track 156, interferometric stack of Mount St. Helens, Washington, composed of 36 interferograms 

spanning preeruptive 1992–2001 period. Inset shows location of stack. Locations of major features referred to in 

the text are labeled. Image shows three areas of subsidence in debris-avalanche deposit—1 (Johnston Ridge), 2 

(Coldwater), and 3 (Elk Rock).
366 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

-20 -10 0 10 20

LOS displacement (mm/yr)

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

5 km

S. Fk. Toutle River

Figure 4. ERS-1/2, track 385, interferometric stack of Mount St. Helens, Washington, composed of 38 

interferograms spanning preeruptive 1992–2001 time period. 

-20 -15 -10 -5 0 5 10

LOS displacement (mm/yr)

122 ̊30 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

Figure 3. ERS-1/2, track 163, interferometric stack of Mount St. Helens, Washington, composed of seven 

interferograms spanning preeruptive 1993–2000 time period.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 367

20 km

-28 0 28 56 84

LOS displacement (mm/yr)

122 ̊30' 122 ̊00' 121 ̊30' 121 ̊00'

46 ̊00'

46 ̊30'

47 ̊00'

5 km

AREA

OF MAP

WASH.

OREG.

5 km

Mount Rainier

Mount Adams

Mount St. Helens

Figure 5. ERS-1/2, track 392, 

interferometric stack of Mount St. 

Helens, Washington, composed 

of seven interferograms spanning 

preeruptive 1995–2000 time period. 

Very little signal is associated with 

Mount Adams, suggesting that 

atmospheric artifacts are minimal.

slightly over the observation period, although the data are too 

noisy to be certain. Both regions, however, are apparent in 

coeruptive stacks that span 2004–5 (figs. 8–14), so neither area 

had ceased deforming by 2005.

Discussion

Lack of Preeruption Deformation

Dzurisin (2003) postulated that deep magma accumulation should precede volcanic unrest and eruption and might 

have gone undetected prior to the May 18, 1980, eruption of 

Mount St. Helens. He also suggested that InSAR is perhaps 

the best tool for detecting such deformation, as exemplified by 

discoveries of aseismic inflation at South Sister (Wicks and 

others, 2002) and Peulik (Lu and others, 2002b) volcanoes. 

Our results rule out surface displacements on the order of 1 

cm or greater due to magma accumulation beneath Mount St. 

Helens between 1992 and late 2001 and between 2003 and 

the start of the 2004 eruption. It is unlikely that the period not 

covered by the InSAR stacks was the only time that displacements occurred; therefore, it is probable that no volcanowide 
368 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

-33 -22 -11 0 11 22

LOS displacement (mm/yr)

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

Figure 6. ENVISAT mode 2, track 156, interferometric stack of Mount St. Helens, Washington, composed of six 

interferograms spanning preeruptive 2003–2004 time period. 

displacements occurred during 1992–2004. How can this 

result be reconciled with Dzurisin’s (2003) proposal?

Seismic and geodetic evidence from the late 1980s and 

early 1990s favor repressurization and possible resupply of the 

magma body located at about 6–10 km depth. Moran (1994) 

examined patterns of seismicity between 1987 and 1992 and 

concluded that the approximately 6- to 10-km-deep reservoir 

was being pressurized, probably by volatiles that exsolved during magma crystallization and were trapped beneath a seal in 

the shallow conduit. Seismicity, gas-emission events, and small 

explosions in 1989–91 may have been related to rupturing of this 

seal and escape of volatiles (Mastin, 1994). A seismic swarm 

in 1998 was accompanied by the emission of CO2 (Gerlach and 

others, this volume, chap. 26), but whether this seismicity was 

related to magma migration or only the ascent of magmatic gases 

is unclear. Magma accumulation below 5.5 km depth is proposed 

by Musumeci and others (2002) to explain patterns of relocated 

seismicity. Lisowski and others (this volume, chap. 15) analyzed 

trilateration data collected in 1982 and 1991 and GPS data from 

2000 and found evidence for areal dilatation between 1982 and 

1991 that may indicate magma recharge. No significant changes 

were observed during the 1991–2000 time period. It is possible 

that magma was accumulating beneath Mount St. Helens only 

during the late 1980s and early 1990s and was manifested by 

deep seismicity, areal dilatation, and phreatic explosions (Mastin, 

1994; Moran, 1994; Musumeci and others, 2002; Moran and others, this volume, chap. 2).

It seems that there is ample evidence for a pressure 

increase—either due to exsolution of volatiles, intrusion of new 

magma, or both—at ~6–10 km depth beneath Mount St. Helens 

in the years between the end of eruptive activity in 1986 and the 

start of the 2004 eruption. Changes in pressure or volume at this 

depth, if large enough, should result in displacements observable 

by InSAR because they would extend well outside the incoherent crater area without being overly broad or diffuse. The fact 

that no inflation was observed by InSAR during 1992–2004 

(assuming that no deformation occurred during late 2001 to early 

2003, when no InSAR data are available) suggests that either 

(1) the magnitude of volume increase (if magma accumulation 

occurred) or pressure increase (if only gas exsolution occurred) 

was insufficient to produce measurable surface displacements, 

(2) the increase in pressure or volume was accommodated by 

inelastic processes that did not deform the surface, or (3) the 

increase in pressure or volume occurred prior to 1992. Distinguishing among these mechanisms is not possible at present, 

although seismic and geodetic data (discussed above) that were 

collected after the end of the 1980–86 eruptive period and before 

InSAR results became available in 1992 favor option 3.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 369

-80

-60

-40

-20

0

20

40

1990 1992 1994 1996 1998 2000 2002 2004 2006

B

-100

-80

-60

-40

-20

0

20

40

60

1990 1992 1994 1996 1998 2000 2002 2004 2006

CULUMATIVE LOS DISPLACEMENT,

IN MILLIMETERS

CULUMATIVE LOS DISPLACEMENT,

IN MILLIMETERS

CULUMATIVE LOS DISPLACEMENT,

IN MILLIMETERS

A

ERS-1/2, track 156

ERS-1/2, track 163

ERS-1/2, track 385

ERS-1/2, track 392

ENVISAT mode 2, track 156

-100

-80

-60

-40

-20

0

20

40

60

80

1990 1992 1994 1996 1998 2000 2002 2004 2006

C

Figure 7. Time series showing line-of-sight (LOS) displacements of subsidence patches at (A) Johnston Ridge (site 1 in fig. 2), (B) 

Coldwater (site 2 in fig. 2), and (C) Elk Rock (site 3 in fig. 2). LOS displacements are relative to one another, so absolute value of Y axis 

is less important than changes in value over time. Note that more than 35 interferograms were used to generate time series for ERS 

tracks 156 and 385, whereas only 6–7 interferograms were used for ERS tracks 163 and 392, and ENVISAT track 156. Deformation 

rates determined from tracks with relatively few input interferograms could therefore be biased by atmospheric delay anomalies.
370 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Subsidence Patches in the Debris-Avalanche 

Deposit

The origin of the localized subsidence patches on the 

debris-avalanche deposit is difficult to constrain and may 

be a result of several different processes. All three patches 

occur on the 1980 debris-avalanche deposit. Pyroclasticflow deposits emplaced during explosive eruptions in 1980 

bury the debris-avalanche unit at the base of Johnston Ridge 

near patch 1 but are not present at the other two patches 

(Rowley and others, 1981). The displacement patterns 

resemble poroelastic deformation caused by groundwater 

withdrawal (for example, Bawden and others, 2001), but 

there has been no deliberate removal of water from the 

debris-avalanche deposit. The subsidence has been occurring since at least 1992, and probably much longer, so subsidence mechanisms that occur on short time scales (days 

to months), including release of gas (Matthews and others, 

2003), hydrocompaction due to rainfall (Hoblitt and others, 1985), and poroelastic deformation of the substrate (Lu 

and others, 2005a) cannot account for the longevity of the 

observed deformation and the lack of significant changes in 

subsidence rate over time (fig. 7). Instead, the deformation 

must be related to long-lived activity associated with the 

May 18, 1980, debris-avalanche deposit. Processes that have 

the potential to cause years-long subsidence of a loosely 

consolidated volcanic deposit like the Mount St. Helens 

debris-avalanche unit include cooling and contraction of an 

initially hot deposit (Masterlark and others, 2006), loading 

of a viscoelastic substrate (Briole and others, 1997; Stevens 

and others, 2001; Lu and others, 2005a; Masterlark and others, 2006), consolidation of a saturated or unusually dilated 

deposit (Major, 2000), and melting of buried ice (Branney 

and Gilbert, 1995; Everest and Bradwell, 2003).

The debris-avalanche deposit was emplaced onto a 

substrate of Tertiary volcanic rocks, with volcaniclastic sediments along river terraces of, and in, the upper part of the 

North Fork Toutle River valley and other drainages (Crandell, 1987; Evarts and others, 1987). The loosely consolidated debris-avalanche deposit has a volume of 2,500×106

 m3

and is composed mostly of cold dacite, andesite, and basalt, 

with probably less than 50×106 m3 of the volume taken up 

by the hot (temperature greater than several hundred degrees 

Celsius) cryptodome that intruded the volcanic edifice in 

March–May 1980 (Voight and others, 1981; Glicken, 1996). 

The deposit reaches its maximum thickness, about 200 m, 

near the base of Johnston Ridge (but not at the exact location of the Johnston Ridge subsidence patch; Glicken, 1996). 

Glacial ice is thought to represent at least 100×106 m3 of 

the volume of the unit (Brugman and Meier, 1981; Glicken, 

1996). Much of this ice was buried by debris, and localized 

melting may have contributed to the formation of the North 

Fork Toutle River lahar on the afternoon of May 18, 1980 

(Fairchild, 1987). Water derived from groundwater in the 

precollapse volcano and glacial ice incorporated into the 

debris avalanche amounted to approximately 250×106

 m3, 

about 12 percent of the total volume of the deposit (Glicken, 

1996). A ground-water system developed in the unit following emplacement, as confirmed by numerous springs and 

seeps in the deposit, and seasonal changes in pond levels 

(J. Major, oral commun., 2006). Some parts of the debrisavalanche deposit were also saturated on emplacement, 

including two of the three patches of subsidence observed 

by InSAR (at the outlet of Coldwater Lake and below Elk 

Rock). The area below Elk Rock may have been the source 

of the May 18, 1980, North Fork Toutle River lahar (Fairchild, 1987; Glicken, 1996), and mud volcanoes caused by 

dewatering were observed near the outlet of Coldwater Lake 

in 1981 (J. Major, oral commun, 2006). In the weeks following May 18, 1980, Banks and Hoblitt (1996) measured 

the temperature of the deposit in shallow holes and found 

temperatures to be less than 100°C, with no increase in 

temperature below about 1.5 m. These data suggest that the 

maximum temperature of the deposit was no greater than 

boiling (Voight and others, 1981; Banks and Hoblitt, 1996), 

although temperatures probably varied widely depending 

on deposit composition (for example, shattered cryptodome 

versus cold volcanic edifice).

The initial emplacement temperature of the deposit 

would have to have been unreasonably high to cause subsidence due to cooling and contraction, as suggested by analogy with pyroclastic flows emplaced in 1986 at Augustine 

Volcano, Alaska. Masterlark and others (2006) used the 

known 3 cm/yr subsidence rate of Augustine’s 1986 deposits 

(measured 13 years after the flows were emplaced) and the 

known maximum 126-m deposit thickness to model an initial 

temperature of 640°C. Decreasing the initial temperature to 

500°C required a maximum pyroclastic-flow thickness of 

336 m to produce the same subsidence rate. By analogy, the 

debris-avalanche deposit at Mount St. Helens would need to 

have an emplacement temperature of over 600°C to produce 

the subsidence observed by InSAR. Considering the small 

amount of hot cryptodome incorporated in the deposit and 

relatively low measured temperatures, it is unlikely that the 

temperature of significant volumes of the unit approached 

the required several hundred degrees.

Subsidence in response to a recently emplaced load 

occurs as the substrate relaxes viscoelastically beneath the 

load. The magnitude of the subsidence should be directly 

proportional to the size of the load and the thickness of the 

substrate (Briole and others, 1997). Subsidence patches on the 

Mount St. Helens debris-avalanche deposit, however, do not 

correlate with deposit thickness, although heterogeneity in pre1980 surficial geology could still allow for subsidence due to 

loading. Loosely consolidated, relatively low-density sediment 

in and around the North Fork Toutle valley may be compacting under the load of the debris-avalanche deposit. Subsidence 

would occur only on parts of the debris-avalanche deposit that 

overlie accumulated sediment, explaining the patchiness of the 

subsidence and the fact that deformation occurs mostly where 

the unit lies near or above the buried channel and floodplains 

of the North Fork Toutle River. Testing this hypothesis is 
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 371

difficult without additional knowledge of the thickness and 

extent of sedimentary deposits prior to the catastrophic debris 

avalanche of May 18, 1980.

Consolidation of granular deposits is a function of water 

content and diffusivity; compaction and subsidence occur in 

saturated deposits as water is removed (Major, 2000). Localized subsidence undoubtedly resulted as water seeped from 

saturated parts of the debris-avalanche deposit, and dewatering 

persisted through 2005 on the basis of springs and seeps found 

throughout the unit (J. Major, oral commun., 2006). Depending on the diffusivity and saturated thickness of the deposit, 

gravitational consolidation resulting from dewatering and diffusion of fluid pressure in excess of hydrostatic pressure could 

last for years, causing subsidence for decades after the deposit 

was emplaced (Major, 2000).

Melting of buried ice as a long-term subsidence mechanism could only occur if ice within the debris-avalanche 

deposit melted slowly over decades. In cold deposits (for 

example, the terminal moraine of a retreating glacier), buried 

ice can persist for hundreds of years (Everest and Bradwell, 

2003). Could ice have survived more than 25 years buried in 

the Mount St. Helens debris-avalanche deposit? The answer 

depends on the deposit temperature and whether ice survived 

as large blocks or pulverized grains. If the entire deposit were 

initially about boiling (Banks and Hoblitt, 1996), there is little 

chance that even large blocks of ice (25–50 m in diameter) 

could have survived more than a few months after May 18, 

1980 (Fairchild, 1987). Some parts of the debris-avalanche 

deposit were probably at ambient temperature, however, 

because much of the collapsed edifice, including most of the 

first landslide block, was composed of the cold outer skin of 

the volcano (including glacial ice; Glicken, 1996). Evidence 

for long-term survival of buried ice at Mount St. Helens is 

suggested by the discovery of subsurface ice during excavations near Spirit Lake by the U.S. Army Corps of Engineers 

almost 2 years after the May 18, 1980, eruption (Glicken, 

1996). Slow melting of buried ice would account for the 

patchiness of the subsidence, as well as the lack of correlation 

between debris avalanche thickness and subsidence magnitude. There is, unfortunately, no way to know whether or not 

buried ice is the source of subsidence of the Mount St. Helens 

debris-avalanche deposit, but it remains a viable mechanism.

All of the possible subsidence mechanisms are asymptotic processes, meaning that subsidence rates of the debrisavalanche deposit patches should decay over time. No significant rate changes are apparent from the InSAR time series (fig. 

7), so the mechanism of subsidence for the debris-avalanche 

deposit patches must be a long-term process, with changes in 

the rate of deformation requiring decades to manifest. Based 

on the temperature and composition of the debris-avalanche 

deposit and substrate, the observed patches of subsidence may 

be caused by differential compaction of sedimentary substrate along the buried North Fork Toutle River, gravitational 

consolidation of locally saturated or unusually dilated areas of 

the debris-avalanche deposit, or melting of buried ice. Using 

current knowledge, the relative contributions of these mechanisms cannot be determined; thus all three processes are viable 

causes for localized areas of subsidence on the May 18, 1980, 

debris-avalanche deposit.

Coeruptive stacks

Results

Unlike the preeruptive stacks, coeruptive stacks are 

inconsistent and heavily affected by atmospheric artifacts 

(figs. 8 through 14). The seven coeruptive stacks show all 

possible deformation patterns—no surface displacements, 

volcanowide inflation, and volcanowide deflation. Deflation of Mount St. Helens is suggested by ENVISAT mode 2, 

track 163 (fig. 9), ENVISAT mode 6, track 20 (fig. 12), and 

RADARSAT mode 2 (fig. 13). No deformation is apparent 

in the ENVISAT mode 2, track 156 (fig. 8), ENVISAT mode 

2, track 385 (fig. 10), and ENVISAT mode 2, track 392 (fig. 

11) stacks. Inflation, at least in two quadrants of the volcano, 

is indicated in the RADARSAT mode 5 (fig. 14) stack. These 

contradictory results can be reconciled by examining the 

character of the subsidence patch located at the base of Johnston Ridge, which was probably active during the coeruptive 

time period based on its persistence during the preeruptive 

interval and lack of decay over time (fig. 7A). This feature 

is only apparent in ENVISAT mode 2, track 163 (fig. 9), 

ENVISAT mode 6, track 20 (fig. 12), and the two RADARSAT beam modes (figs. 13 and 14). Stacks that do not show 

the subsidence patches are not sensitive to displacement rates 

of 2–3 cm/yr, on the same order as the expected volcanowide 

deflation signal measured by GPS (Lisowski and others, this 

volume, chap. 15).

The three interferograms that show volcanowide 

subsidence (figs. 9, 12, 13) are the stacks produced from at 

least nine input interferograms (table 1). We suspect that 

the averaged range changes in figures 8, 10, 11, and 14 are 

biased by atmospheric delay anomalies due to the limited 

number of interferograms used in the stacks. These stacks 

should not be considered for further analysis. Atmospheric 

delay artifacts in the stacks may also be recognized by 

topography-correlated fringes at presumably undeforming 

mountains also covered by the scenes (Beaudecel and others, 2000). Unfortunately, no mountains other than Mount 

St. Helens are coherent within, or covered by, the ENVISAT 

mode 2, track 163 (fig. 9) and ENVISAT mode 6, track 20 

(fig. 12) stacks, and their quality cannot be assessed independently. Mount Adams, 60 km east of Mount St. Helens, has no associated phase-correlated topography in the 

RADARSAT mode 2 stack (fig. 13) which, together with 

the debris-avalanche subsidence patches also visible in that 

image, suggests that the stack is sensitive to deformation on 

the order of 1 cm/yr. As a result, the stack that is most representative of the coeruptive deformation state of Mount St. 

Helens is RADARSAT mode 2 (fig. 13). This stack suggests 
372 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

-32 -16 0 16 32

LOS displacement (mm/yr)

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

-40 -20 0 20 40 60

LOS displacement (mm/yr)

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

Figure 8. ENVISAT mode 2, track 156, interferometric stack of Mount St. Helens, Washington, 

composed of seven interferograms spanning coeruptive 2004–5 time period. 

Figure 9. ENVISAT mode 2, track 163, interferometric stack of Mount St. Helens, Washington, 

composed of 13 interferograms spanning coeruptive 2004–5 time period.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 373

122 ̊00' 121 ̊30' 121  ̊00'

46 ̊00'

46 ̊30'

-300 -150 0 150

LOS displacement (mm/yr)

AREA

OF MAP

WASH.

OREG .

20 km

5 km

Mount St. Helens Mount Adams

Figure 10. ENVISAT mode 2, track 385, interferometric stack of Mount St. Helens, Washington, composed of three 

interferograms spanning coeruptive 2004–5 time period. All three interferograms cover times of 70 days or fewer, so stacked 

phase has low signal-to-noise ratio and is characterized by phase that correlates with topography. This is especially apparent 

at Mount Adams. 
374 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊ 10'

46 ̊ 20'

LOS displacement (mm/yr)

-13 0 13 26 39 52

Spirit Lake

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

-14 0 14

LOS displacement (mm/yr)

-21 -7 7

122 ̊30' 122 ̊20' 122 ̊10'

46 ̊10'

46 ̊20'

North Fork Toutle River

AREA

OF MAP

WASH.

OREG.

S. Fk. Toutle River

5 km

Spirit Lake

Figure 11. ENVISAT mode 2, track 392, interferometric stack of Mount St. Helens, Washington, 

composed of six interferograms spanning coeruptive 2004–5 time period. 

Figure 12. ENVISAT mode 6, track 20, interferometric stack of Mount St. Helens, Washington, 

composed of 13 interferograms spanning coeruptive 2004–5 time period.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 375

LOS displacement (mm/yr)

-30 0 30 60

123 ̊00' 122 ̊30' 122 ̊00' 121 ̊30' 121 ̊00'

AREA

OF MAP

WASH.

OREG.

46 ̊00'

46 ̊30'

47 ̊00'

20 km

Mount Rainier

Mount St. Helens Mount Adams

10 km

Figure 13. RADARSAT standard mode 2 ascending interferometric stack of Mount St. Helens, Washington, composed of 

nine interferograms spanning coeruptive 2004–5 time period. Stack clearly shows line-of-sight lengthening (subsidence) 

centered on Mount St. Helens which, as discussed in text, most likely reflects deformation of the ground surface. 

Atmospheric artifacts are probably minimized, as suggested by the lack of a significant signal associated with Mount Adams.
376 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

-15 0 15 30 45

LOS displacement (mm/yr)

123 ̊00' 122 ̊30' 122 ̊00' 121 ̊30'

WASH.

OREG.

AREA

OF MAP

46 ̊00'

46 ̊30'

45 ̊30'

20 km

10 km

Mount Hood

Mount St. Helens Mount Adams

Figure 14. RADARSAT standard mode 5 ascending interferometric stack of Mount St. Helens, Washington, composed of six 

interferograms spanning coeruptive 2004–5 time period. Topography-correlated phase at Mount Hood suggests the presence 

of atmospheric artifacts in the stack.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 377

LOS lengthening (subsidence) of at least 40–50 mm/yr centered on Mount St. Helens (the maximum is probably greater 

but not recoverable given the incoherence in the crater) and 

extends radially outwards beyond the flanks of the volcano.

Discussion

We modeled the RADARSAT mode 2 dataset (fig. 13) 

by inverting the LOS displacements for a buried point source 

of volume change (Mogi, 1958). Other types of sources (for 

example, an ellipsoid or circular crack) might yield a better fit 

to the data, but it is difficult to justify the use of more complicated source geometries in light of the limited coherence 

around Mount St. Helens, especially in the crater area where 

the deformation appears to reach a maximum. The best-fitting 

point source is located at a depth of 12 km directly beneath the 

crater of Mount St. Helens and has a volume loss of 27×106

m3/yr (fig. 15). The depth compares favorably with a point 

source model based on GPS data (Lisowski and others, this 

volume, chap. 15), but it is deeper than more-complicated 

ellipsoidal source models (Lisowski and others, this volume, 

chap. 15; Mastin and others, this volume, chap. 22) and the 

6–10-km-depth range for a proposed magma reservoir based 

on seismic data (Scandone and Malone, 1985; Barker and 

Malone, 1991; Moran, 1994; Musumeci and others, 2002). 

The greater InSAR-derived model depth is probably due to the 

lack of InSAR data in the crater area and the availability of 

only a single component of displacement from one interferometric stack.

We also modeled the data using a horizontal dislocation 

(Okada, 1985) constrained to uniform opening only, which 

approximates a sill. The best-fitting dislocation was located 

at a depth of 18 km, much deeper than the point source, but 

it had a similar volume change of about 30×106

 m3/yr. The 

difference in depths between the point source and dislocation 

models is unsurprising. A similar relation, for example, was 

found between point source and horizontal dislocation models 

of subsidence at Medicine Lake volcano, northern California 

(Dzurisin and others, 2002). The great variability in depths 

between the models is an indication of the strong dependence 

of model fits on model geometry (Delaney and McTigue, 

1994). More data, especially from the crater region where 

displacement magnitudes are greatest, and varied look angles, 

which would provide additional components of displacement, 

are required to better constrain the geometry, volume change, 

and depth of the subsidence source, as well as model error 

estimates. Using only a single interferometric stack, results are 

unavoidably more ambiguous.

The modeled volume change is much smaller than the 

approximately 70×106 m3 extruded over the first year of 

the eruption, as determined from analyses of digital elevation models (Schilling and others, this volume, chap. 8). 

Some of the difference may be explained by the fact that 

the stack averages the initially greater rates of subsidence 

and extrusion that occurred during the first few weeks 

of the eruption, but this is probably a minor effect. It is 

tempting to speculate that most of the discrepancy between 

modeled and extruded volumes is evidence for replenishment of the midcrustal magma reservoir by an even deeper 

source. Delaney and McTigue (1994), however, pointed out 

that the volume of injection should only equal the volume 

of eruption if the host rock is incompressible—clearly an 

unrealistic condition. They also demonstrated that modeled 

volume changes at depth are highly dependent on source 

geometry. In addition, our models assume a homogenous, 

isotropic, elastic rheology, which is almost certainly not the 

case beneath Mount St. Helens, where temperatures will be 

10 km

A B C

-30 -15 0 15 30

LOS displacement (mm/yr)

Figure 15. Observed (A), predicted (B), and residual (C ) line-of-sight (LOS) displacements for RADARSAT standard mode 2 

stack of Mount St. Helens, Washington (fig. 13), resulting from model that assumes point source of volume decrease (red circle) 

at depth of 12 km. LOS displacements scaled to maximize displacements around Mount St. Helens and differ from those in figure 

13 for same interferogram.
378 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

elevated above the local geotherm due to the presence of the 

active magmatic system. Nonelastic rheologies can have a 

profound effect on the relation between surface deformation 

volume and source volume or pressure change (for example, 

Newman and others, 2001, 2006).

Nevertheless, it remains possible that the midcrustal 

magma reservoir was being replenished during the eruption, and that the discrepancy between modeled and extruded 

volumes was due to magma extrusion outpacing influx into 

the reservoir. Any surface inflation that might have resulted 

from recharge of the reservoir would have been masked by the 

greater deflation signal resulting from magma withdrawal and 

depressurization. Mastin and others (this volume, chap. 22) 

address this possibility in greater detail.

Conclusions

Interferograms from several independent tracks have 

been examined for signs of pre-and coeruptive deformation 

at Mount St. Helens. We can find no evidence for volcanowide displacements during time periods covered by stacks of 

radar interferograms (1992 to late 2001 and early 2003 until 

the onset of eruptive activity in September 2004). Seismic 

and geodetic evidence suggests repressurization, and possible 

resupply, of the 6- to 10-km-depth magma reservoir during 

the late 1980s and early 1990s, but this activity did not result 

in surface deformation detectable by InSAR during 1992–

2004. We did find several small (1–2-km-diameter) patches 

of subsidence on the 1980 debris-avalanche deposit, one of 

which was recognized previously from individual interferograms of Mount St. Helens, and attribute this deformation 

to viscoelastic relaxation of loosely consolidated substrate 

under a load, differential consolidation of the deposit, or 

melting of ice buried within the deposit. Tracking the continued motion of these areas will serve as input for models of 

the post-depositional behavior of rapidly emplaced, unconsolidated deposits.

Coeruptive interferometric stacks are dominated by atmospheric noise, and only one of the seven stacks we assembled 

is demonstrably sensitive to deformation on the order of 1 cm 

or less. This stack indicates volcanowide deflation centered on 

the crater of Mount St. Helens and can be modeled by a point 

source of volume loss at a depth of 12 km beneath the edifice. 

This depth is at the lower boundary of a magma reservoir 

inferred from seismicity before 2004.

Both the pre- and coeruptive InSAR results have important implications for magma resupply, or lack thereof, before 

and during the 2004 eruption. Several papers in this volume 

address this question from differing perspectives and reach 

differing conclusions (for example, Moran and others, chap. 

2; Dzurisin and others, chap. 14; Lisowski and others, chap. 

15; Mastin and others, chap. 22; Gerlach and others, chap. 

26; Pallister and others, chap. 30). Neither InSAR nor any 

other single data set can resolve the issue with certainty, but 

the weight of evidence seems to favor at least a minor amount 

of magma influx or pressurization of the midcrustal magma 

reservoir in the years before, and perhaps during, the eruption. 

The InSAR results, although not directly supportive of the 

idea, can be reconciled with it if the preeruption changes were 

small enough, accommodated by inelastic processes that did 

not measurably deform the surface, or occurred before 1992, 

and if the coeruption changes reflect a net deflation of the 

volcano due to extrusion outpacing recharge.

It is worth noting that, in general, the RADARSAT results 

are more coherent than those from ERS-1/2 and ENVISAT. A 

possible explanation for this discrepancy is that RADARSAT’s 

H/H polarization provides better coherence in vegetated areas 

than V/V, which is the standard mode for ERS-1/2 and ENVISAT. Additional investigations into this possibility are warranted, as the result may suggest that one polarization mode is 

preferred over others for InSAR studies.

Our work has demonstrated the importance of including numerous interferograms in stacks. Coeruption stacks 

with less than nine input interferograms showed no signs of 

deflation that is known to be occurring from GPS. That our 

preeruption stacks span many more years and make use of 

many more interferograms than the coeruption stacks is the 

reason for the better signal-to-noise ratio in the preeruption 

stacks. The addition to the coeruptive stacks of InSAR data 

that extend through 2006 and beyond, assuming the eruption 

continues, will significantly reduce the magnitude of the atmospheric noise, allowing for more detailed analysis of surface 

deformation and associated source mechanisms.

Acknowledgments

Summer Miller provided invaluable assistance in 

preparing the ERS-1/2 and ENVISAT data sets for stacking. 

Paul Lundgren provided code for time series analysis, which 

Sarah Menassian modified for use with this project. Rick 

Hoblitt and Don Swanson offered much-needed guidance 

with respect to the emplacement of the Mount St. Helens 

debris avalanche, and discussions with Jon Major, Dick 

Iverson, and Joe Walder stimulated a closer look at the subsidence patches in the debris-avalanche deposit. Jon Major’s 

advice concerning the behavior of saturated sedimentary 

deposits was especially valuable. Oh-Ig Kwoun assisted 

with processing RADARSAT data. We are indebted to 

Chuck Wicks, Daniel Dzurisin, and Seth Moran for thoughtful reviews and Jane Takahashi for thorough edits. Figures 

1–6 and 8–15 were made using the Generic Mapping Tools 

(GMT) software (Wessel and Smith, 1998). ERS-1, ERS2, and ENVISAT SAR data are copyrighted 1992–2005 by 

ESA and provided by Alaska Satellite Facility (ASF) and 

ESA Category-1 grants 2648 and 2765. RADARSAT-1 data 

are copyrighted by Canadian Space Agency and were provided by ASF. We thank the ASF for supporting RADARSAT-1 data acquisitions.
18. Radar Interferometry Observations of Surface Displacements During Pre- and Coeruptive Periods at Mount St. Helens 379

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384 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 385

Models and Mechanics of Eruptive 

Processes 

Empirical data from seismology, geophysics, field geology, and petrology 

give us the precursory signals and ongoing history of the eruption. However, 

to understand how volcanoes work—and, consequently, to have a volcanomonitoring program with predictive value—requires deeper insight into an 

eruption’s physical mechanisms. A “sticky piston” model of solid-state extrusion was suggested by the drumbeat aspect of the 2004–6 eruption’s seismicity. Geologists made a renewed effort to catch the stick-slip process in action, 

using fixed-camera photography to monitor its visual manifestation and borehole tiltmeters to measure the associated inflation-deflation cycles expected 

in local ground deformation.

Laboratory testing of the gouge that encases the extruding dome indicates 

material properties suitable for the sticky-piston model to predict episodic, 

seismogenic slip, namely plug stiffness that exceeds that of the rising fluid 

magma and rate-weakening gouge friction at low slip rates. Textural evidence indicates that crystallization and solidification of the ascending magma 

preceded gouge formation. Comparison of groundmass textures from dome 

lavas and fault gouge suggests that brittle fracture was confined to the upper 

1 km of the conduit. Textural similarity between gouge and the ash emitted 

by explosions from spine margins in January and March 2005 suggests that 

fragmentation preceded rather than accompanied these explosions.

Eruption rates and modeling of the magma reservoir allow reasonable fits 

to either exponential or logarithmic decay curves. For each of these, however, 

the history of reservoir decompression forms an imperfect match to the rapid, 

radially inward motion of the ground surface during the first month of the 

eruption, as deduced from displacement data from the continually transmitting GPS station located 9 km from the vent. That movement suggests that 

erupted magma was replaced by recharge, but that the recharge rate remained 

less than the eruption’s characteristic effusion rate.
386 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Spine 4 on February 22, 2005. View to east-southeast. Shoestring notch is low 

point on east crater rim. USGS photo by S.P. Schilling.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 19

1

 Department of Geological Sciences, 1272 University of Oregon, Eugene, 

OR 97403

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

From Dome to Dust: Shallow Crystallization and 

Fragmentation of Conduit Magma During the 2004–2006 

Dome Extrusion of Mount St. Helens, Washington

By Katharine V. Cashman1, Carl R. Thornber2, and John S. Pallister2

Abstract

An unusual feature of the 2004–6 eruptive activity of 

Mount St. Helens has been the continuous growth of successive spines that are mantled by thick fault gouge. Fault gouge 

formation requires, first, solidification of ascending magma 

within the conduit, then brittle fragmentation and cataclastic 

flow. We document these processes through field relations, 

hand samples, and thin-section textures. Field observations 

show that the gouge zone is typically 1–3 m thick and that it 

includes cataclasite and, locally, breccia in addition to unconsolidated (true) gouge. The gouge contains multiple slickenside sets oriented subparallel to each other and to the striation 

direction, as well as surface striations parallel to extrusion 

direction. Hand specimens show the cataclasite and gouge to 

be composed of a wide size range of broken dome and wallrock fragments. This grain-size heterogeneity is even more 

pronounced in thin section, where individual samples contain 

fragments that span more than four orders of magnitude in size 

(from more than 10 to less than 10-3 mm).

Textures preserved within the gouge zone provide 

evidence of different processes operating in time and space. 

Most individual fragments are holocrystalline, suggesting 

that crystallization of the ascending magma preceded gouge 

formation. Cataclasite samples preserve a wide range of clast 

sizes; pronounced rounding of many clasts indicates extensive abrasion during transport. Within the gouge, crystals and 

lava fragments adjacent to finely comminuted shear zones 

(slickensides) are shattered into small, angular fragments that 

are either preserved in place, with little disruption, or incorporated into shear trains, creating a well-developed foliation. Together, evidence of initial grain shattering, followed 

by shear, grinding, and wear, suggests extensive transport 

distances (large strains). Textural transitions are often abrupt, 

indicating extreme shear localization during transport. Comparison of groundmass textures from dome lavas and fault 

gouge further suggests that brittle fracture was confined to the 

upper 400–500 m of the conduit. Observed magma extrusion 

(ascent) rates of ~7 m/d (8×10-5 m/s) permit several weeks for 

magma ascent from ~1,000 m (where groundmass crystallization becomes important) to ~500 m (where solidification 

nears completion). Brittle fracture, cataclastic flow, and shear 

localization (slickenside formation) probably dominated in the 

upper 500 m of the conduit.

Comparison of eruptive conditions during the 2004–6 

activity at Mount St. Helens with those of other spine-forming 

eruptions suggests that magma ascent rates of about 10-4 m/s 

or less allow sufficient degassing and crystallization within 

the conduit to form large volcanic spines of intermediate 

composition (andesite to dacite). Solidification deep within 

the conduit, in turn, requires transport of the solid plug over 

long distances (hundreds of meters); resultant large strains are 

responsible for extensive brittle breakage and development of 

thick gouge zones. Moreover, similarities between gouge textures and those of ash emitted by explosions from spine margins indicate that fault gouge is the origin for the ash. As the 

comminution and generation of ash-sized particles was clearly 

a multistep process, this observation suggests that fragmentation preceded, rather than accompanied, these explosions.

Introduction

The 2004–6 eruption of Mount St. Helens has been characterized by steady lava effusion in the form of a succession of 

smooth-surfaced dacitic spines (Scott and others, this volume, 
388 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

chap. 1) that are remarkable for the striated fault gouge that 

forms a carapace on each of the seven spines erupted between 

October 2004 and February 2006. This fault gouge is unusually well preserved and well sampled, thanks to helicopter 

dredging operations (see Pallister and others, this volume, 

chap. 30). Additional geophysical constraints on dome extrusion include photogrammetry (Schilling and others, this 

volume, chap. 8), thermal observations (Schneider and others, 

this volume, chap. 17), and seismic observations (Moran and 

others, this volume, chap. 2). Notable in these observations is 

the remarkable steadiness of both the extrusion and accompanying “drumbeat” earthquakes, all of which appear to originate at depths <1 km directly below the growing dome (Moran 

and others, this volume, chap. 2) and are interpreted to record 

steady stick-slip behavior of the ascending magma plug (Iverson and others, 2006; Iverson, this volume, chap. 21).

Individual outcrops document an abrupt transition from 

competent flow-banded dacite to a zone of breccia and cataclasite to finely comminuted and variably consolidated gouge. 

In this paper we use field, hand-specimen, and thin-section 

observations of these cataclastic dome facies to constrain 

some of the physical processes related to dome extrusion and 

gouge formation. Analysis of observations over this range 

in scale allows us to (1) describe progressive fragmentation 

of the solid dome rock to fine-grained fault gouge, (2) trace 

decompression/crystallization paths within the conduit, and 

(3) provide depth constraints on both solidification and onset 

of brittle deformation. We then link processes of fragmentation and gouge formation to relative rates of ascent (shear), 

degassing, and crystallization. This linkage is important both 

for recognizing conditions that may lead to changes in the 

emplacement conditions of the current dome at Mount St. 

Helens and for generalizing observations at Mount St. Helens 

to other recent effusive eruptions of intermediate composition 

magma. In addition, the spectacular preservation of fault textures allows us to extend analysis of tectonic gouge formation 

to volcanic environments and to compare brittle deformation 

processes in highly crystalline dacite with those of tuffisites 

in obsidian flows. Finally, the similarity of clasts preserved 

within the gouge to those found within ash from dome explosions (Rowe and others, this volume, chap. 29) allows us to 

speculate on the nature of explosions that have accompanied 

dome formation.

Background

As detailed descriptions of lava dome emplacement are 

provided in other papers in this volume (Scott and others, 

chap. 1; Schilling and others, chap. 8; Vallance and others, 

chap. 9), we briefly review only the elements of the dome’s 

emplacement history that are directly relevant to the conditions 

of gouge formation. We then provide an overview of spine formation at other volcanoes, with a particular emphasis on both 

observations of spine emplacement and eruption conditions 

that produced these features. Finally, we discuss observational 

and experimental constraints on decompression-driven crystallization required to form spines.

2004–2006 Activity at Mount St. Helens

After about three weeks of precursory seismic unrest 

and phreatic explosions in late September and early October 

of 2004, a dacite lava spine emerged within the central crater 

of Mount St. Helens, to the south of the 1980–86 lava dome. 

Extrusion has been continuous since mid-October 2004, 

with emplacement and partial to complete disintegration of 

seven individual spines. Extrusion rates were highest, about 6 

m3/s, during the first month of the eruption, after which they 

dropped to 1–2 m3/s. Linear rates of spine emergence from 

the vent have also been constant at 3–6 m/day (3.5–7×10-5

m/s), except during the early stages of extrusion when rates 

were about twice as fast. Extrusion has been accompanied by 

only two major explosions—one on January 16, 2005, and 

another on March 8, 2005. Ash produced by those explosions 

is described by Rowe and others (this volume, chap. 29).

Lava Dome Emplacement and Morphology

Howel Williams (1932) presented the first comprehensive 

description of volcanic dome emplacement. He used the term 

“dome” to describe viscous lava protrusions that could be 

subdivided into (1) “plug domes” (spines), (2) “endogenous 

domes” that grow by expansion from within and (3) “exogenous domes” built by surface effusion, usually from a central 

summit crater. He noted that effusion rate (Q) provided the 

strongest control on growth style, with the lowest effusion rate 

leading to plug dome (spine) formation. As a type example for 

plug-dome behavior he used the famous spine extruded from 

Mont Pelée, Martinique, in 1903. This massive spine reached 

a maximum height of >300 m, grew as rapidly as 25 meters 

in a day (2.9×10-4 m/s), and was disrupted continually by both 

collapse and explosions. The spine was described by Jaggar 

(1904) as “ * * * a most extraordinary monolith, shaped like 

the dorsal fin of a shark, with a steep and almost overhanging 

escarpment on the east, while the western aspect of the spine 

was curved and smooth in profile. The field glass showed 

jagged surfaces on the steeper eastern side, and long smooth 

striated slopes on the western” (fig. 1). These smooth, striated 

slopes also preserved slickensides and breccias produced by 

differential movement of already solid lava. Lacroix (1904) 

noted both vertical striations parallel to the extrusion direction and horizontal ridges indicating “successive stages in its 

upheaval” (Williams, 1932).

Recent eruptions at Mount St. Helens, Washington; 

Unzen, Japan; and Soufrière Hills, Montserrat, confirm and 

extend these early concepts. Dome growth at Mount St. Helens 

from 1980 to 1986 had both exogenous and endogenous phases 

and erupted a total volume of 0.077 km3

 of dacite magma 

(62.5–63.5 weight percent SiO2; Swanson and Holcomb 1990). 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 389

April 3,

1903

March 9,

1903

Oct. 4,

1902

July 31,

1902

Before eruption

June 13,

1903

Sept. 30,

1903

1593 m 1350 m

1270 m

1230 m

1230 m

1582 m

1447 m

1365 m

1380 m

Figure 1. Sketches showing evolution of the great spine 

of Mont Pelée, Martinique, erupted in 1903. Note the 

difference between smooth, curved fault surface and rough, 

near-vertical face. From Williams (1932).

Exogenous growth of individual flow lobes was the dominant 

dome growth mechanism through 1981, with lobes emplaced 

at rates of 9–24 m3/s during time periods of less than a week. 

Exogenous dome lobes were also emplaced from March 1984 

until October 1986, although accompanying endogenous 

growth became increasingly important during this time period 

(Fink and others, 1990). Times of anomalous activity included 

(1) March 1982, when a protracted effusive event (24 days at 

an average rate of 1.6 m3/s) was preceded by deep seismicity 

and a vulcanian explosion; (2) February 1983–February 1984, 

when effusion was continuous at a rate of 0.7 m3/s (approximately equivalent to the long-term magma supply rate from 

October 1980 to October 1981) and dome growth was primarily endogenous; and (3) May 1985, when effusion lasted 17 

days (at an average rate of 2.9 m3/s). Only two spines were 

emplaced during the 1980–86 effusive activity, one in February 1983, during the period of continuous endogenous growth, 

and one in May 1985, at the end of a protracted (and largely 

endogenous) dome growth episode. Magma batches feeding 

individual exogenous dome lobes ascended at 0.1–1 m/s (Chadwick and others, 1988; Endo and others, 1990; Geschwind 

and Rutherford, 1995), corresponding to decompression rates 

of ~0.0025–0.025 MPa/s. Because effusion rates were 10–20 

times lower during 1983, we infer that ascent rates during periods of continuous dome growth were correspondingly lower 

(at least by a factor of 2–5), with magma ascent rates for spine 

production even lower (probably <10-3 m/s).

Links between eruption rate and dome morphology can 

be further constrained by examining two other eruptions of 

slightly different bulk composition. Unzen volcano, Japan, 

active from 1990 to 1995, constructed a 0.2-km3 dome of 

dacitic composition (64.5–66 wt. percent SiO2). Initial effusion (Q) at ~1 m3

/s produced a spine. Shortly thereafter a rapid 

increase in Q to ~5 m3/s triggered a succession of pyroclastic 

flows accompanied by tremor spurts and two vulcanian eruptions (Nakada and Fujii, 1993; Nakada and others, 1995b, 

1999; Nakada and Motomura, 1999). After this time the dome 

grew nearly continuously, with exogenous growth dominating 

during periods of higher effusion rates (>4 m3/s) and endogenous growth dominant when effusion rates were low (<1 m3

/s). 

Magma ascent rates were estimated to be 13–40 m/d (1–5×10-4

m/s; Nakada and others, 1995b), more than an order of magnitude lower than during the time period of continuous magma 

ascent at Mount St. Helens in 1983. Effusion rates declined 

gradually until the end of the eruption, which was marked by 

the emergence of a spine from the endogenous dome (Nakada 

and others, 1995a).

Soufrière Hills volcano, Montserrat, intermittently active 

since 1995, has also shown a close correspondence between the 

rate of lava effusion and dome morphology. By 1999 the volcano 

had produced almost 0.3 km3

 of silicic andesite (57–61.5 wt. 

percent SiO2), about two-thirds of which traveled as pumiceous 

and dense pyroclastic flows into valleys around the volcano. 

Low effusion rates (Q ~0.5–2 m3/s) produced a range of spine 

and whaleback features, most of which grew in a matter of days. 

Moderate effusion rates (Q ~2–7 m3/s) produced shear lobes that 

were active for days to weeks to months. High effusion rates 

(Q ~7–9 m3/s) produced pancake-shaped lava lobes with flat 

tops and scoriaceous surfaces. Very high effusion rates (Q >9 

m3/s) resulted in vulcanian explosions (Watts and others, 2002). 

Of particular interest are features that bear close similarity to 

current activity at Mount St. Helens, specifically vertical spines 

and megaspines that characterized magma emplacement at low 

volumetric effusion rates (<2 m3/s; Watts and others, 2002).

Together these case studies present a consistent story of 

spine growth under conditions of slow magma effusion (typically <1–2 m 3/s). Descriptions of both smooth and striated spine 

surfaces are common; Watts and others (2002) inferred that these 

characteristic curved surfaces were controlled by shear faults (for 

example, Donnadieu and Merle, 1998) in magma made highly 

viscous by late-stage decompression-driven crystallization. As 

noted above, spine morphology generally has been interpreted to 

reflect extrusion of a solid lava plug, which probably solidified 

more as a consequence of decompression-induced crystallization than of cooling on the way to the surface. However, there 

has been no detailed characterization of these bounding-fault 

surfaces or comparison of striated surfaces with tectonic fault 

textures, and no attempt has been made to use either faultsurface characteristics or crystallization textures to constrain the 

conditions of magma ascent and decompression that create these 

volcanic spine features.
390 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The Importance of Decompression-Driven 

Crystallization

During the early part of the 20th century, there was great 

debate about the solidification mechanism that could generate 

the large Mont Pelée spine. Lacroix (1904) favored rapid cooling of lava on exit from the vent. Gilbert (1904) suggested that 

solidification of the spine might result from “escape of gases 

and formation of bubbles” rather than loss of heat by conduction. Gilbert’s concept was extended by Shepherd and Merwin 

(1927), who argued that gas pressures required to support the 

spine (~10 MPa) resulted from crystallization; they went on to 

suggest that these high gas pressures also caused the explosive 

generation of nuées ardentes that accompanied spine extrusion. Williams (1932) extended the concept of crystallizationinduced gas pressurization to explain late-stage explosions at 

several other domes, an association that Newhall and Melson 

(1983) find in most historical eruptions.

Dome growth during the 1980–86 eruption of Mount 

St. Helens provided the first quantitative documentation of 

decompression-driven crystallization in silicic magmas (Cashman, 1992). Decompression-driven crystallization has since 

been documented during effusive and intermittent explosive 

phases of numerous recent eruptions (for example, Nakada 

and others, 1995b; Wolf and Eichelberger, 1997; Gardner 

and others, 1998; Nakada and Motomura, 1999; Hammer and 

others, 1999, 2000; Cashman and Blundy, 2000; Martel and 

others, 2000; Sparks and others, 2000; Metrich and others, 

2001; Blundy and Cashman, 2005; Cashman and McConnell, 

2005) and is generally considered to be the dominant solidification mechanism in slowly ascending hydrous magma of 

intermediate composition. Petrographic observations show 

that plagioclase is the most abundant shallow-crystallizing 

phase, a consequence of the profound effect of dissolved H2

O 

on its stability. Quartz (or some other silica phase) joins the 

crystallizing assemblage when the quartz-feldspar cotectic 

is intersected, which typically occurs at low pressures, and 

rapidly increases the total crystal abundance. Both the number 

and mode of plagioclase crystals vary with decompression 

path (effective undercooling; ∆Teff).

Decompression experiments on H2O-saturated melts 

(Geschwind and Rutherford, 1995; Hammer and Rutherford, 

2002; Martel and Schmidt, 2003; Couch and others, 2003) 

show that plagioclase number density, size, and morphology are controlled by the rate of decompression and the 

final pressure of equilibration. The extent of crystallization 

is dependent on both equilibration pressure and experiment 

duration (Hammer and Rutherford, 2002; fig. 2). Crystallinity increases as equilibration pressure decreases and 

with increasing time at a given pressure. At low pressure 

(≤25 MPa), extensive crystallization requires time scales 

of several weeks. Crystal nucleation rates also are highest 

at pressures between 25 and 10 MPa, although nucleation 

diminishes in importance at very low pressure (≤5 MPa) 

owing to the high viscosity of silicic melts with low H2O 

(Hess and Dingwell, 1996). Extensive co-precipitation of 

quartz and feldspar appears to require equilibration at very 

low pressure (5–10 MPa).

An additional consequence of extensive late-stage decompression-driven crystallization is the release of latent heat. 

Evidence of late-stage heating in slowly erupted dome lavas 

includes both Ti-zoning of magnetite, which increases calculated Fe-Ti oxide temperatures, and reverse zoning of orthopyroxene (Devine and others, 2003; Pallister and others, 2005). 

These signs of heating have been interpreted as evidence of 

basaltic input at depth (for example, Devine and others, 2003). 

However, a recent study of samples erupted from Mount St. 

Helens in the early 1980s shows a strong correlation between 

late-stage heating and shallow crystallization, suggesting 

instead that this is a signature of shallow release of latent heat 

(Blundy and others, 2006). Because numerical models of conduit-flow dynamics show that even small changes in magma 

temperature have a large effect on magma discharge rate 

and eruptive behavior (Melnik and Sparks, 2005), localized 

crystallization-related heating may be an important component 

of shallow eruption dynamics.

In summary, the morphology of extruded spines—their 

asymmetric shape, smooth fault-bounded surfaces, and often 

imposing height—is controlled by effusion conditions. Formation of these spine features requires that magma ascent is 

sufficiently slow to allow near-complete volatile exsolution, 

degassing, and groundmass crystallization within the conduit. 

Outstanding questions include the relative timing of groundmass 

crystallization and fault formation, the potential role of latent 

heating in shear localization, the strain accommodated along 

these fault surfaces, the deformation mechanisms by which 

CRYSTALLINITY, IN PERCENT

FINAL PRESSURE,

IN MEGAPASCALS

250

200

150

100

50

0

0 20 40 60 80

Long-duration experiments

Single-step decompression

EXPLANATION

Long-duration experiments

Short-duration experiments

Multistep decompression

Figure 2. Relative effects of decompression time and equilibration 

pressure on crystallization of hydrous silicic melts. Single-step 

decompressions are followed by annealing at the final pressure; 

short-duration experiments, <30 hours; long-duration experiments, 

168–931 hours. Multistep decompressions occur at constant rate of 

1.2 MPa/hr. Data from Hammer and Rutherford (2002). Arrow shows 

equilibrium decompression-crystallization trend.
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 391

resulting fault gouge may affect plug extrusion, and the relation 

between observed seismic signals and spine emplacement.

Methods

Field observations and digital photography of the 

2004–6 Mount St. Helens dome, made during two short 

ground-based sampling missions and numerous close-in 

helicopter flights, provide an outcrop-scale context for 

describing the character and origin of the gouge. Because 

of safety concerns, sampling of the dome and the dome 

carapace were accomplished primarily by helicopter dredging (Pallister and others, this volume, chap. 30). Although 

this remote technique did not allow detailed mapping of 

field relations, it did allow access to steep and unstable parts 

of the spines that could not have been reached on foot. The 

source areas for debris and talus samples are, in most cases, 

well constrained by field documentation. Moreover, both 

photographic and seismic records document the locations 

and timing of the specific rockfall events sampled. Finally, 

data collected during two days of field work in the summer 

of 2006 confirmed field relations inferred from remote techniques (Pallister and others, this volume, chap. 30). Distinct 

differences in sample lithology can be correlated with dome 

rock facies observed in the scree slopes on the dome flanks. 

A complete catalogue of sample information, including 

locations, brief descriptions, and bulk chemistry, is provided 

by Thornber and others (2008b).

All dome samples have been described and curated at the 

Cascades Volcano Observatory (Thornber and others, 2008b). 

Samples of unconsolidated fault gouge (SH303 and SH307) 

were processed in the Cascades Volcano Observatory Sediment 

Laboratory. Samples were dry sieved to 63 μm; size fractions 

<63 μm were analyzed using a Micrometrics Sedigraph III 

particle size analyzer. Additionally, both dome lava and fault 

gouge were characterized by petrographic analysis. Whole thin 

sections were scanned at high resolution for general textural 

characterization and quantitative analysis of plagioclase phenocryst and microphenocryst textures (that is, crystals >30 μm 

in diameter). Microscopic evaluation of sample textures was 

accomplished using a JSM-6300V scanning electron microscope at the University of Oregon. Thin sections were imaged 

in backscattered electron (BSE) mode using 10-kV accelerating 

voltage and 5-nA beam current. Both scanned and BSE digital 

images were processed and analyzed using Adobe Photoshop 

and ImageJ 1.34s (NIH Image for Mac OS X) software. Most 

techniques that were used are similar to those described elsewhere (for example, Cashman and McConnell, 2005); however, 

we also experimented with ways of analyzing the size distribution of clasts in the cataclasite for comparison with sieve data. 

Analysis of larger (>30 μm) dome fragments was performed on 

low magnification (20×) BSE images. Gray-scale thresholding 

of individual clasts provided binary images that could be easily 

processed to allow clast size analysis (fig. 3). All image-analysis data were then analyzed using Microsoft Excel.

Figure 3. Examples from gouge sample SH326-2A illustrating 

the image-processing technique used to acquire grain-size data 

on cataclastic samples. Apparent circular concentration of very 

fine particles is an artifact of the image-acquisition technique; 

for this reason, only larger particles are used in final analysis. 

Images are 6.1 mm across. A, Backscattered electron image. B, 

Processed binary image.

Results

In this section we document processes related to spine 

extrusion. Because the observational record of spine extrusion is unique in the level of detail permitted by the sampling 

techniques, our primary goal is to provide comprehensive 

descriptions of textures produced as competent dacite lava 

is transformed into ultracataclastic fault gouge. We begin by 

reviewing details of field observations, followed by hand-specimen descriptions and thin-section observations. To describe 

the textural characteristics of fault rock samples, we follow 

Snoke and others (1998) in using the nomenclature conven-
392 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

tion of Sibson (1977), as modified by Scholz (1990), where 

“gouge” refers to noncohesive material with visible fragments 

constituting <30 percent (by volume) of the rock mass, in 

contrast to fault “breccia” that has >30 percent visible fragments. Although this naming convention would apply the term 

“cataclasite” to all cohesive brittle fault rocks, we distinguish 

cohesive breccia from cataclasite on the basis of apparent clast 

versus matrix support. Moreover, for simplicity we refer to the 

“gouge zone” when discussing the entire suite of cataclastic 

material that overlies massive dome dacite.

The View from the Field

The gouge carapace was first observed on November 7, 

2004, on the surface of spine 3. At that time a large lava spine 

(300–400 m north-south by 50–150 m east-west) had been 

thrust to ~100 m above the surface of the already-uplifted 

glacier in the southeastern quadrant of the crater (figs. 4A, 4B). 

On the east face of the spine, a slope of gouge was exposed 

above a basal apron of wet, sandy debris containing angular lava blocks. The lower 10–20 m of the slope’s exposed 

east face was composed of gouge that displayed prominent 

striations that plunged steeply east (parallel to the 45o slope 

of the spine) in exposures along the southeast face and more 

northerly along the northeast face. Along both faces a resistant 

surface layer was underlain by ≤1 m of white, powdery gouge 

(sample SH303) that was slightly warm to the touch (~30°C), 

dry, and friable. Sparse rock fragments within the gouge 

accumulated fine-grained material on their downslope sides, 

producing features akin to pressure shadows (Pabst, 1931) 

with a preserved sense of shear that was consistent with uplift 

of the spine core relative to marginal rocks (fig. 4B). Several 

curvilinear bands of discolored gouge and trails of embedded 

lithic fragments (“bathtub rings”; figs. 4A, 4B) paralleled the 

contact between the gouge and the debris apron at the base of 

the spine, recording earlier positions of this basal contact that 

shifted upward with growth. Disintegration of the south face of 

spine 3 (figs. 4C, 4D) allowed sampling of the (still hot) spine 

interior by helicopter dredging. The dominant rock type is 

dense hornblende-hypersthene-plagioclase-phyric dacite with 

a microcrystalline groundmass (SH304; Pallister and others, 

this volume, chap. 30).

The detailed descriptions of spine 3 provided above can 

be generalized to all other spines. A narrow band of mantling 

gouge typically forms an intact annulus at the vent margin (fig. 

5A). Thermal data from a Forward-Looking Infrared Radiometer (FLIR) indicate temperatures of 100–200°C at this margin 

but show that the gouge surface cools to near ambient temperatures after about 10 m of uplift. A FLIR temperature maxima of 

730°C (Schneider and others, this volume, chap. 17) indicates 

that temperatures remain high in the spine interior (within a 

few meters of the surface). Throughout the 2004–6 eruption, 

the thickness of the gouge zone has ranged from 1 to 3 m, 

always with surface striations oriented parallel to transport (fig. 

5B). Most commonly, dense flow-banded dacite is abruptly 

overlain by 1–2 m of reddish cataclasite, which is overlain 

by poorly consolidated white to tan powdery gouge (fig. 5C). 

Occurring less frequently are examples of foliated gouge with 

parallel striations on the surface that form a penetrative shear 

foliation (fig. 5D). Samples of foliated gouge show multiple 

subparallel slickensides that indicate slight variations in slip 

direction at different depths. In some outcrops, a capping 

resistant (and striated) surface layer protects the unconsolidated 

gouge that, at times, posed challenges to sampling efforts. On 

one occasion, a 75-kg dredge was dragged along the surface of 

spine 4 without penetrating this surface layer. Bulk chemical 

analyses of gouge samples show that early gouge included a 

substantial fraction of older material from the conduit margin, 

whereas later gouge samples are mostly juvenile dacite (Rowe 

and others, this volume, chap. 29).

An unusual, well-exposed section of the lava-to-gouge 

boundary within a graben that formed in spine 4 was photographed in December 2005, more than a year after spine extrusion (fig. 6). Exposed cross sections of the gouge-lava contact 

showed both breccia and pink (altered) cataclasite separating 

massive flow-banded dacite from overlying powdery gouge. 

The abrupt contact between massive dacite and the overlying cataclasite indicates that this is an important mechanical boundary. The progression from breccia to gouge was 

well displayed, with blocks of intact dacite sheared from the 

underlying lava and progressively fragmented to produce breccia, cataclasite, and gouge. The sense of brittle shear derived 

from the inclination of shear surfaces within the breccia is 

consistent with upward transport from the vent. Photographic 

interpretations were confirmed by field observations made in 

the summer of 2006. The observations showed that the gouge 

formed over a vertical distance of 1–2 m by shear between the 

upward-thrusting dacite within the conduit and the conduit 

margins. Reidel shears within the cataclasite provide further 

evidence of brittle shear failure (Pallister and others, 2006).

Hand Samples

Rock samples collected from the talus of the growing 

2004–6 Mount St. Helens dacite dome include both competent 

dacite (igneous rock) and a variety of fault rocks (cataclasite). 

Detailed descriptions of all samples can be found in Thornber 

and others (2008b). Here we briefly describe the varieties of 

pristine dome dacite before describing the range of fault rocks. 

Dacite samples vary in color from light gray to reddish or pink 

(a term used in deference to the first unequivocally juvenile 

samples of “hot pink” dacite SH304). The reddish alteration 

results from oxidation that is localized around hornblende, 

oxide, and hypersthene phenocrysts and does not noticeably 

affect the otherwise igneous fabric of the rocks.

In hand sample, the igneous facies of dome rocks may 

be broadly classified as (1) dense, light-gray (to pink) dacite, 

(2) flow-banded gray (to pink) dacite, or (3) vesicular darkgray (to pink) dacite. Most samples also contain 1–5 volume 

percent of coarse-grained hornblende-gabbronorite xenoliths 

that are similar in abundance and character to those in 1980 to 

1986 eruption products (Heliker, 1995). Lithic inclusions of 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 393

A

C D

B

i

ii

iii

Figure 4. Spine 3 in early to mid-November 2004. USGS photographs by J.S. Pallister. A, West-looking view on November 4 

showing, from the base upward: (i) an apron of sandy debris, snow, and glacial ice; (ii) 10–15-m-high wall of white slickensidebearing gouge with “bathtub rings” that mark former contact levels between gouge surface and debris apron; (iii) dark tan 

(weathered) gouge or cataclasite with steeply dipping erosional furrows. Blocky outcrops (uppermost spine rough-weathering 

area near center-frame) are uplifted material from the debris apron. Steam is venting from fractures in gouge carapace. B, 

Close-up photograph on November 4 of gouge outcrop showing slickensides and “bathtub” rings that parallel contact between 

the gouge and the debris apron. Note shadows on the downslope sides of rock fragments within the gouge that suggest upward 

and westward movement of the spine relative to the debris fan. Sample SH303-1 was collected from the white powdery gouge at 

site marked by X, about 1 m above the debris apron 10 m north (right) of the prominent steam vent. C, Oblique aerial photograph on 

November 4 of the southeast collapsed terminus of spine 3 showing blocky outcrops of dense dacite of the spine interior. View to 

the north-northwest. D, Photograph on November 10 of the southwest face and south terminus of spine 3 with a dark (ash covered) 

“bulldozed” glacier on right. White gouge carapace is visible above the blocky interior of the collapsing southwest face; powdery 

gouge eroded from the carapace has been deposited on the debris apron on the east flank of the spine. View is to the northeast.

dacite, andesite, and amphibolite are less abundant and typically of smaller size.

Dense light-gray dacite samples dominate the dome 

talus suite. Because these samples are both angular and dense 

(porosities <10 percent), and because several of these samples 

were dredged from collapsed faces of the spines, we infer that 

they are derived primarily from the dome interior. Flow fabric 

within the dome facies is not well developed and is indicated 

only by (1) alignment of acicular hornblende phenocrysts in 

some samples and (2) entrainment patterns of xenolithic fragments and vesicular pockets around them. Rare flow-banded 

dacite samples have distinctly light-colored or reddish bands 

(ranging from less than 10 to 20 mm in width) that are subparallel to foliation evident in dense, dacite host rock. Bands lack 

evidence of granular flow or other forms of brittle deformation 

and instead are characterized by high vesicularity relative to 

the host rock.

Vesicular dacite samples have 25–40 percent porosity 

and range in color from dark gray to red. In some vesicular 

samples, igneous flow is evidenced by fanned vesicle-distribution patterns and corresponding flow alignment of xenolith 

fragments and feldspar megacrysts. Proximity of the most 
394 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A B

C D

Figure 5. Features of the new lava dome. USGS photos by J.S. Pallister. A, Oblique aerial photograph of crater and 2004–2005 

lava dome taken from the northwest on June 29, 2005. Spine 3 is in a shadow on left; spine 4 is in the center (highly fractured and 

altered); and spine 5 is on the right, with white, striated gouge carapace on the north and west flanks, collapse scar on northwest 

face, and sharp, arcuate contact between hot, powdery gouge and talus marking the periphery of the vent. Crevassed and brown, 

ash-covered Crater Glacier surrounds the dome. The following three photos were taken on January 14, 2005. B, Striated surface 

of spine 4; direction of transport is from lower right to upper left. Orange box on instrument package at center of frame is 1 m long. 

C, Contacts between gouge (gray, weathered surface lacking distinct striations), underlying cataclastic breccia, and dense dacite 

exposed on the collapsed south face of spine 4. Note flow foliation in lower right of photograph and in some outcrops immediately 

below the gouge. Dacite lava is overlain abruptly by pink to yellow consolidated gouge, which in turn is mantled by powdery white to 

gray gouge. Preserved thickness of gouge zone above contact is 1–2 m, although original thickness was greater. D, Parallel striations 

on surface of spine 4; transport direction from lower right to upper left. Surface striations mark penetrative fabric that creates planar 

shear lamination of the cataclasite. Block of dacite collected from the interior of nearby spine 4 is flow foliated and displays a strong 

lineation of elongate amphiboles.

vesicular, glassy, and reddish dacite to the outer dome margins 

is inferred from in-place sampling (SH305-1; fig. 7) and is 

documented in the exposed cross sections of collapsed spines. 

Helium-pycnometer measurements of SH305-1 yield a Heaccessible porosity of 30.4 percent, nearly identical to the 30.3 

percent porosity determined from the measured bulk density 

of 1,829 kg/m3 and measured solid density of 2,627 kg/m3. 

The similarity of the connected and bulk porosity indicates 

that the pores are fully connected and, therefore, might allow 

permeable flow of gas upward along the margin of the conduit. 

Preliminary permeability measurements of the same samples 

yield permeabilities of 8–9×10-13 m2, confirming the possible 

ease of gas flow through the more vesicular dome lavas.

Both abrupt and gradational changes from competent 

dacite to finely powdered gouge are found within the sample 

suite. Some fractured and foliated dacite samples reveal 

a subtle gradation between vesicular dacite and marginal 

granular gouge. For example, samples from 2–5 m below 

surface gouge collected in the SH309 dredge along the southern end of spine 4 are somewhat altered and show moderate 

crushing and compaction associated with postsolidification 

shearing. Cataclastic rocks are distinguished by visible frag-
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 395

A B Residual gouge

Altered cataclasite

Angular breccia

Flow-banded dacite

Figure 6. Dome-breccia relations. USGS photo by J.S. 

Pallister, December 15, 2005. A, Erosional remnant atop spine 

4 preserves a breccia interval between massive to flowbanded dacite lava at base of outcrop and a thin surface 

layer of white powdery gouge. Brittle fractures (Reidel 

shears) are present in breccia and cataclasite. B, Trend lines 

indicate shear couple with basal dacite moving left (away 

from vent) relative to the upper (external) part of outcrop, 

consistent with shear between the dacite and the conduit 

wall during extrusion. Height of outcrop is approximately 3 m.

Figure 7. Outcrop and hand-specimen photographs of sample 

SH305, collected January 3, 2005, from spine 3.

ments of dome lava suspended within a finer grained matrix 

(figs. 8A, 8B) and range from well consolidated to friable and 

from coarse to fine grained. Both the size and abundance of 

macroscopic lava fragments decrease toward the outer surface of fine-grained noncohesive gouge. Friable and foliated 

gouge from spine 4 (SH313-1) shows pronounced slickensides forming subparallel striations that permeate the loosely 

consolidated sample (fig. 8C). Powdery fault gouge is light 

in color and unconsolidated.

Samples in Thin Section

The petrography and petrology of dacite lava dome 

samples are described by Pallister and others (this volume, 

chap. 30) and will be reviewed only briefly here. The dacite 

lava has a uniform bulk composition (65 wt. percent SiO2) and 

is highly crystalline, having 41–45 percent phenocrysts and 

microphenocrysts of plagioclase, hypersthene, amphiboles, and 

Fe-Ti oxides set in a microcrystalline matrix (fig. 9A). Irontitanium oxide temperatures in early samples are 840–850°C; 

apparent temperatures increase in later samples in response to 

latent heat produced by extensive shallow crystallization (for 

example, Blundy and others, 2006).

Thin sections provide important insight into fine-scale 

cataclasis of competent dome lava. An unusual sample 

(SH324-4A) preserves an abrupt boundary between coherent 

dome lava and highly fragmented fault gouge. In thin section, 

the boundary between gouge and coherent lava is marked by 

thin, oxidized bands that are subparallel to the boundary (fig. 

9B). Approximately parallel to these bands, more subtle variations in matrix density mark a throughgoing shear fabric within 

the boundary zone. More common are samples of cataclasite 

that are characterized by rounded fragments of dome lava 

within a finer grained matrix (SH326; fig. 9C). In thin section, dome fragments typically range in size from greater than 

10 mm to less than 0.1 mm. The matrix comprises primarily shattered phenocrysts, although pieces of holocrystalline 

groundmass can also be identified. In friable, foliated gouge 

samples with obvious slickensides, the grain-size reduction vis-
396 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

B

D

A

C

10 mm

Figure 8. Hand specimen and cut chips showing textural 

gradation from cataclasite to intensely comminuted fault gouge. A, 

Lithified breccia sample SH326-2A (spine 7); Scale in centimeters. 

B, cataclasite sample SH315-2 (spine 4); Scale in millimeters. C, 

striated gouge sample SH313-1 (spine 4). Scale in millimeters.

100

100 200

200
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 397

ible in thin section is extreme (SH313-1; fig. 9D), with typical 

cataclasite (fig. 9C) cut by an extraordinarily fine-grained, 

foliated shear zone (an ultracataclasite, with grain size 

<1 μm). Other thin sections of the same gouge sample show 

anastomosing branches of foliated ultracataclasite that isolate 

areas of coarser cataclasite.

Backscattered Electron Imaging Observations

Field, hand-specimen, and thin-section observations 

of dome and gouge samples indicate that competent (and 

flow-banded) dome lava achieved threshold conditions for 

wholesale fragmentation, cataclastic flow, and strain localization to accommodate the large strains achieved during spine 

emplacement. Backscattered electron (BSE) imaging provides 

additional textural evidence for (1) the crystallization history 

of the dome magma that may have led to fragmentation, (2) 

fragmentation mechanisms, and (3) cataclastic flow.

Detailed characterization of phenocryst textures and composition is provided elsewhere in this volume (Pallister and 

others, chap. 30; Rutherford and Devine, chap. 31; Streck and 

others, chap. 34), so we limit our descriptions of phenocrysts 

to characteristics that may help to constrain shallow conduit 

processes. A striking feature of virtually all samples we have 

imaged is the presence of some plagioclase phenocrysts (>100 

μm) with irregular margins on one or more growth faces that 

indicate rapid late-stage rim growth (fig. 10). In contrast, 

microphenocrysts (30–100 μm) typically are euhedral and 

prismatic. Phenocryst rim-growth irregularities range from a 

few to several tens of microns in width (fig. 10A) and include 

either multiple (fig. 10B) or single (fig. 10C) growth zones. 

In some samples, irregular growth protrusions partially entrap 

small pockets of groundmass glass that are substantially 

less crystalline and less vesicular than adjacent matrix (fig. 

10C). The combined observations of rapid rim growth in both 

early-erupted (fast) and late-erupted (slow) dome lavas and 

clean glass preserved within the growth irregularities suggest 

that these features reflect relatively fast magma ascent (and 

overstepping of the plagioclase liquidus) at deeper levels in 

the conduit system. The prevalence of growth irregularities in 

large (and often outsized) phenocrysts further suggests that 

these crystals may have a unique decompression history.

Although the groundmass of all samples is highly crystalline, groundmass textures are variable. At one extreme are 

juvenile samples, collected by dredge in early November 2004 

(SH304), that include several unusually glassy clasts. One 

such clast has only pyroxene and oxides as groundmass phases 

 Figure 9. Full thin-section scans of samples showing 

transition from competent dome lava to highly comminuted 

fault gouge. Scale bar in D applies to all four panels. A, Sample 

SH316, from spine 5. B, Sample SH324-4A, from spine 7. C, Sample 

SH326-2A, from spine 7; box shows transect illustrated in figure 

14A. D, Sample SH313-1C, from spine 4; box shows transect 

illustrated in figure 16A.

A 50 μm

B 250 μm

C 50 μm

Figure 10. Plagioclase phenocryst textures indicative of rapid 

growth. A, Sample SH316. B, Sample SH315-2. C, Sample SH309-1. 
398 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

50 μm 10 μm

50 μm 25 μm

A B

C D 

30). More typical of early-erupted samples is SH305-1, from 

spine 3, emplaced in late November 2004 (for example, fig. 

7). Sample SH305-1 is distinct from SH304-2G3 because it 

has less-numerous, but somewhat more-euhedral, plagioclase 

microlites in a groundmass mottled with 30–40-μm patches 

of cotectic feldspar and quartz crystallization, intergrowths 

that commonly nucleate along plagioclase phenocryst and 

microphenocryst boundaries and then grow into the melt (figs. 

11C, 11D). Disequilibrium crystallization is a common feature 

of natural and experimental melts decompressed slowly to 

shallow levels (Hammer and Rutherford, 2002; Martel and 

Schmidt, 2003), the result of delayed quartz nucleation in 

these highly silicic melts (Cashman and Blundy, 2000; Blundy 

and Cashman, 2001). Quartz may have nucleated more easily along euhedral crystal margins if a thin, slightly SiO2

-

enriched, boundary layer helped to overcome the activation 

energy required for quartz nucleation in the rhyolitic liquid.

Figure 11. Backscattered electron images of groundmass textures from early-erupted samples (spine 3). A, Sample 

SH304-2G3 showing relatively high glass content and dominance of plagioclase microlites (medium- to light-gray crystals). 

B, Highly magnified view. C, Sample SH305-1 showing patchy development of cotectic quartz-feldspar intergrowths 

(blotchy dark-gray regions) with apparent nucleation on phenocryst margins. D, Highly magnified view.

and preserves ~ 2 weight percent H2O in the groundmass glass 

(Pallister and others, this volume, chap. 30). Another sample 

(SH304-2G3; figs. 11A, 11B) has a groundmass composed 

primarily of feldspar microlites (<20 μm in length) with only 

minor pyroxene, oxides, and quartz. This sample is unusual in 

having abundant clean glass (~25 percent) and only minimal 

precipitation of silica phases. Many of the larger plagioclase 

microlites in this sample have swallowtail terminations that 

suggest rapid growth (for example, Hammer and Rutherford, 

2002). High crystal number densities (~107/mm3) preserved in 

this sample would suggest that most of the growth was both 

shallow (<25 MPa; see compilation in Cashman and McConnell, 2005) and rapid. These pressure estimates are roughly 

consistent with the ~30 MPa indicated by preserved H2O 

contents of some other SH304 clasts, confirming that glassy 

SH304 samples appear to have been quenched at depths of 

~1 km (for example, Pallister and others, this volume, chap. 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 399

By April 2005, the rate of dome extrusion had diminished and groundmass textures were correspondingly more 

crystalline. Dome dacite sample SH316 (from spine 5; 

fig. 9A) has a highly crystalline groundmass with abundant plagioclase crystals and patchy cotectic intergrowths 

of plagioclase and quartz (fig. 12A). Also apparent in the 

groundmass of SH316 are larger (~50 μm in length) acicular 

(platy in three dimensions) silica-phase crystals that appear 

to be a form of cristobalite (fig. 12A). Other samples, such 

as SH309, show abundant tridymite as a groundmass phase 

(figs. 12B, 12C), a feature that was also observed by Blundy 

and Cashman (2001) in some samples from the continuous growth phase of the 1983 Mount St. Helens dome. The 

long axes of tridymite crystals (fig. 12B) are aligned locally, 

recording late-stage deformation.

Backscattered electron imaging also provides a spectacular record of gouge development. As described above, 

dome dacite samples (fig. 13A) show little sign of deformation 

associated with magma ascent and extrusion, except for some 

samples with pronounced flow banding. Brittle breakage associated with flow is manifested first as throughgoing fractures, 

particularly along phenocryst boundaries (fig. 13B). Cataclasite sample SH326-2A shows more extensive evidence of 

crushing. Thin sections illustrate the textural heterogeneity of 

SH326-2A, with small (≤10 mm in diameter) dome lava clasts 

distributed randomly in a matrix of smaller clasts, individual 

crystals, and pulverized matrix (fig. 9C). Backscattered electron images show that the matrix itself contains a continuous 

range of fragment sizes, from intact lava clasts greater than 

1 mm to tiny fragments smaller than 1 μm (fig. 14A), where 

A B

C 50 μm D 10 μm

50 μm 25 μm

Figure 12. BSE images of groundmass textures from later-erupted samples. A, SH316 (spine 5) showing nearholocrystalline nature of groundmass. Dark-gray phase is quartz, medium-gray phase is plagioclase, light-gray phase is 

quenched glass, and white phases are mafic minerals. B, High-magnification view of SH316 groundmass showing high 

number density of plagioclase microlites in remnant, clean glass patches. C, SH309 (spine 4) showing unusual abundance 

of acicular aligned tridymite crystals in the groundmass. D, high magnification view of SH309 groundmass showing the 

intergrowths of plagioclase and tridymite (dark gray) and remnant patches of clean glass. 
400 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

crude banding is suggested by regions of higher and lower 

fragment concentrations.

All samples of cataclasite and gouge show nearly complete 

groundmass crystallization; groundmass textures preserved within 

individual clasts are similar to the more crystalline parts of domelava groundmass. Larger fragments within the gouge typically 

are cored by plagioclase phenocrysts enclosed in holocrystalline 

(and microvesicular) groundmass (figs. 14B, 14C), although 

clasts composed entirely of microphenocrysts and microvesicular, 

holocrystalline groundmass are not uncommon. Commonly, the 

outer margin of clasts is distinctly rounded and is surrounded by 

very fine grained fragments, features that suggest extensive grinding, abrasion, and even rolling during transport.

Textural transitions from dome lava to gouge to ultracataclasite can be observed in more detail in samples SH324-4A 

(fig. 9B) and SH313-1 (fig. 9D). In SH324-4A (fig. 15), the 

pronounced boundary between dome lava and gouge seen in 

thin section marks an abrupt change from intact, holocrystalline, igneous groundmass to fragmented and disseminated 

matrix plus crystals and crystal fragments (fig. 15A). In detail, 

the boundary is composed of thin (~100 μm) slivers of intact 

holocrystalline groundmass separated by bands of loose crystal 

and groundmass fragments (fig. 15B). In figure 15B, the shearzone margin shows a single plagioclase crystal in the process 

of being shattered, providing the source of the aligned crystal 

fragments seen in figure 15D. Patterns of fracture may provide 

a sense of shear, as in figure 15C, where conjugate fracture 

sets propagate through a dome lava fragment containing both 

crystals and holocrystalline groundmass.

A traverse across foliated gouge sample SH313-1 (figs. 

8C, 9D) shows more spectacular grain-size variation (fig. 16A). 

Much of the thin section is similar to cataclasite sample SH326 

(fig. 14), with intact clasts of dome lava suspended in a finegrained matrix. However, as shown in BSE imagery, the narrow 

(≤1 mm) fault zone visible in thin section (fig. 9D) is composed 

of alternating bands (~100 μm wide) of fine-grained (tens of 

micrometers) and extremely fine grained (<1 μm) material 

(figs. 16B, 16C). By analogy with features observed in 

SH324-4A (fig. 15), we infer that banding develops by incorporation of shattered crystals (fig. 15B) and dome fragments 

A 1 mm 1 mm B 

A

B C

1 mm

250 μm 100 μm

Figure 13. Low-magnification BSE 

images of two spine 4 samples. A, 

Competent dacite sample SH308-3A. B, 

partially fragmented sample SH309-1C. 

Note extensive fracturing (black areas 

indicate void space).

Figure 14. Backscattered electron images 

of fault gouge textures preserved in sample 

SH326-2A (spine 7; see fig. 9C for smallerscale view). A, Composite transect illustrating 

textures characteristic of cataclasite samples. 

B, C, Typical gouge fragments composed of 

individual phenocrysts or broken dome matrix 

surrounded by holocrystalline matrix. Typically 

clasts are rounded and are surrounded by a 

halo of finely pulverized material.
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 401

A B

C D 

1 mm 100 μm

250 μm 100 μm

Figure 15. Backscattered electron images 

showing details of the lava-cataclasite 

boundary preserved in sample SH324-4A, 

from spine 7 (see fig. 9B for smaller-scale 

view of boundary). A, Low-magnification 

image across the boundary showing 

coherent dacite on left and cataclasite 

on right. B, Highly magnified view of 

boundary (from upper center of A) showing 

extensive fragmentation of a crystal 

next to the boundary and thin slivers of 

intact holocrystalline matrix parallel to 

the boundary. C, Lava fragment showing 

(Reidel?) fractures through plagioclase 

microphenocrysts and holocrystalline 

matrix within gouge. D, Broken slivers of 

holocrystalline matrix just below image of 

B that show lateral entrainment of broken 

crystal fragments into boundary zone. 

(fig. 16D) into adjacent shear zones. Foliation orientations 

are both parallel and subparallel to the shear-zone orientation, 

consistent with macroscopic evidence for slight orientation 

variations among slickensides (fig. 8C). Moreover, in some 

locations along the shear zone, individual fragments of foliated 

gouge have been broken and rotated (figs. 16C, 16D). In other 

locations, the shear zone itself is anastomosing, forming an 

ultracataclastic zone that varies in width and complexity (fig. 

16E). Here too, the orientation of individual foliated layers 

within the gouge is subparallel to the shear zone.

Grain-Size Measurements

Sieve data for two powdery gouge samples—SH303-1 

collected on November 4, 2004, from spine 3 and SH307-1 

collected on February 22, 2005, from a collapsed part of 

spine 4—confirm the wide range in grain size of the gouge 

samples (fig. 17A). The two samples differ, however, in the 

extent of comminution. Sample SH307-1 has a bimodal size 

distribution with peaks at 10 and 350 μm; it has little material 

<1 μm in size and measurable clasts >5 mm. In contrast, the 

granulometric analysis for SH303-1 describes a rather broad 

peak with a poorly defined mode at about 32 μm, a substantial 

proportion of very fine grains, and no clasts >2.8 mm. This 

difference is best illustrated in a plot of cumulative volume, 

which shows the pronounced difference in volumetric contribution of very small particles (fig. 17B).

Grain size of fragments can also be estimated by analysis 

of BSE images. Although measurements of the number of 

clasts of a given size in area fraction are not directly comparable to measurements of the weight percent of a given sieve 

A

B C D 

E

1 mm

400 μm 600 μm 750 μm

1 mm

Figure 16. Backscattered electron images showing details 

of fault gouge sample SH313-1 from spine 4 (see fig. 9D for 

smaller-scale view). A, Transect across part of thin section 

SH313-1C showing a wide range of clast sizes in the gouge 

and extreme grain-size reduction within the “fault core,” 

right of center. B, Detailed image of fine-scale foliated fault 

gouge showing slight offset in foliation orientation across 

the image. C, Breakage, disruption, and partial rotation 

of finely pulverized gouge. D, Shattered lava fragment 

adjacent to the fault core. E, Composite transect along the 

fault zone in thin section SH313-1A. Note fine-scale foliation 

oriented subparallel to the fault zone and anastomosing 

strands that isolate regions of less extensively pulverized 

gouge fragments. 
402 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006 CUMULATIVE VOLUME, IN PERCENT

PARTICLE DIAMETER, IN MICROMETERS

1

0.1

10

100

0.1 10 1,000

SH303-1

SH307-1

0.25 1.0 4 16 62.5 250 1,000 4,000

SIEVE DIAMETER, IN MICROMETERS

5

10

0

SH303-1

SH307-1

NORMALIZED FREQUENCY, IN PERCENT

A

B

16 32 63 125 250 500 1,000 2,000

0.25

0.15

0.05

0.20

0.10

0

DIAMETER, IN MICROMETERS

NORMALIZED AREA

LOG DIAMETER, IN MICROMETERS

LOG NUMBER

y = -1.97x + 6.09

R

2

 = 0.99

A

B

4

3

2

1

-1

0

0 1 2 3 4

Figure 17. Sieve data for unconsolidated fault gouge 

samples SH303-1 (spine 3) and SH307-1 (spine 4). A, 

Histograms. B, Power-law plots. Note the difference in 

abundance of very fine particles (<10 μm) in the two samples 

and the absence of single power-law trend.

Figure 18. Grain-size data obtained by image analysis of BSE 

image of SH326-4A from spine 7. A, Histogram of data plotted by 

area (related to but not directly equivalent to volume-based sieve 

data). B, Power-law plot, with y-axis showing cumulative number 

of fragments greater than a given size. Data fit a power-law 

distribution over the size range 10–1,000 μm.

sample (for example, fig. 9B; table 1). This range encompasses 

that of all other samples. Accompanying the decrease in crystal 

abundance is a decrease in average phenocryst size, which 

reflects breakage of large crystals to form numerous small 

crystal fragments. This breakage most severely affects the large 

phenocrysts, as shown by normalized cumulative plagioclase 

volume distributions (fig. 19). The median (50th percentile) 

plagioclase size decreases from about 300 μm in competent 

dome lava to about half of that in partially comminuted cataclasite sample 324-4A (fig. 19). More striking is the difference 

in maximum crystal size, which reaches several millimeters in 

the dome lava but only about 300 μm in SH324-4A cataclasite 

(table 1). If this analysis were extended to measurement of the 

finely laminated ultracataclasite zone of sample SH313-1 (fig. 

16), the plagioclase size reduction would be even more extreme.

In summary, we have experimented with different 

ways of quantifying the observed fragmentation of competent dome rocks. Grain-size changes in fault gouge typisize, the image-analysis data provide another view of grainsize characteristics that can be compared directly with data 

from tectonic faults. Illustrative measurements on images from 

SH326-2A (figs. 3, 8A, 9C) show that the number-based data 

are dominated by the smallest size classes, which are difficult 

to measure accurately. When viewed as areal abundance, the 

data show a mode at 500 μm, similar to the volume-based 

(sieve) data for SH307-1 (fig. 18A). These data lack an equivalent peak at 8 μm because this was below our detection limit 

(~30 μm) for the imaging magnification used.

Quantifying the transition from intact dome lava to powdery gouge is more difficult. We experimented with measuring 

the reduction in size and abundance of plagioclase phenocrysts 

in thin section, a measurement that is easy to make using 

standard image-analysis techniques. The overall abundance of 

plagioclase, as measured in full thin-section scans (for example, 

fig. 9), decreases from a maximum of 32 percent in sample 

SH324-4A to 9.5 percent in consolidated gouge from the same 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 403

SH324-4A(1)

SH324-4A(2)

SH324-5(2)

SH315-2

SH316(1)

SH316(2)

SH320-1A2

SH320-1C1(2)

SH307-2A1

NORMALIZED CUMULATIVE VOLUME

DIAMETER, IN MICROMETERS

0.8

1.0

0.6

0.4

0.2

0

0 200 400 600 800 1,000

Figure 19. Plagioclase size distributions plotted as normalized 

cumulative volume for comparison with sieve data. Distributions 

determined by image analysis of full thin-section images, as 

shown in figure 9. Minimum size resolution approximately 

30 μm (microphenocrysts and phenocrysts). Data show that 

plagioclase size distributions of undeformed dacite lava change 

progressively with increasing fragmentation toward smaller 

median and maximum crystal diameters. This change in size 

distribution is accompanied by a decrease in overall abundance 

of macroscopic plagioclase.

Sample No. Plagioclase

area fraction

Total

area (mm2

)

Total number

of crystals

Average

crystal 

number

Maximum

diameter 

(mm)

Minimum

diameter 

(mm)

SH315 26.2 355 1,692 4.77 3.32 0.1

SH316(1) 27.9 222.5 2,441 10.97 1.19 0.03

SH316(2) 25.2 219.37 2,389 10.89 0.85 0.03

SH320-1C1(1) 11.7 185.54 2,404 12.96 1.01 0.02

SH320-1C1(2) 12.1 174.67 2,253 12.90 0.624 0.02

SH320-1A2 21.5 250.75 3,068 12.24 1.32 0.02

SH324-4A(1)-a 24 40.58 324 7.98 0.73 0.04

SH324-4A(2)-a 9.5 27.19 289 10.63 0.34 0.04

SH324-4A(1)-b 32 143.83 1,772 12.32 1.43 0.04

SH324-4A(2)-b 19.3 87.41 1,324 15.15 0.608 0.04

SH324-5(1) 25.8 198.9 3,496 17.58 0.91 0.02

SH324-5(2) 24.2 125.37 1,974 15.75 1.22 0.02

SH307-2A1 22.4 479.98 3,330 6.94 2.27 0.05

Table 1. Plagioclase phenocryst and microphenocryst data obtained from image analysis of whole thinsection scans.

[An example scan is shown in figure 9. Reported are the measured plagioclase area fraction, the total area measured, the total 

number of plagioclase crystals measured, the average plagioclase number per area, and the maximum and minimum plagioclase 

diameter. Where abrupt textural changes were preserved within a single thin section (for example, fig. 9B), different textural 

regimes were measured separately (denoted a and b).]
404 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

cally are measured by sieving, although this is possible 

only for unconsolidated samples. In cataclasite, clast size 

distributions may be analyzed using BSE images, although 

complete sample characterization would require combining multiple images at different magnifications. A simple 

characterization of progressive cataclasis of dome lava can 

be made using thin-section measurements of plagioclase 

phenocryst size and abundance. These data provide evidence 

of extensive fragmentation and grain-size reduction of dome 

lava during transport to the surface. Moreover, the common 

observations of abrupt textural transitions indicate extreme 

shear localization during transport. When combined with 

BSE evidence of initial grain shattering (figs. 15B, 16D) followed by shear, grinding, and wear (figs. 14B, 14C), these 

data suggest extensive transport distances (large strains) 

after initial breakage.

Discussion

The results presented above illustrate the range of processes that contributed to the transformation of dome lava to 

fault gouge and, in some cases, to ash emitted by small dome 

explosions (Rowe and others, this volume, chap. 29). In the 

following sections, we examine those processes in detail, paying particular attention to the role of degassing-induced crystallization and to conditions of fault formation. We generalize 

these observations by placing them first within the perspective 

provided by past activity (1980–1986) at Mount St. Helens 

and then within the context of other spine-forming eruptions.

The Role of Decompression Crystallization

All clasts within both cataclasite and fault gouge are 

nearly holocrystalline. Thus we infer that fragmentation 

of dome dacite occurred after most of the groundmass had 

crystallized. Extensive crystallization is most likely caused 

by decompression and degassing during the slow ascent of 

hydrous magma (reviewed above), although we cannot rule 

out the possibility of minor late-stage cooling along the 

conduit margin. Extensive crystallization, in turn, changes the 

rheology of the rising magma, particularly when coupled with 

changes in melt viscosity resulting from volatile exsolution 

(for example, Hess and Dingwell, 1996; Melnik and Sparks, 

2002, 2005). At Mount St. Helens, these rheological changes 

transformed the deformation properties of the magma from 

plastic to brittle, fragmenting the rising magma plug and creating the resultant thick fault gouge. Constraining the depth at 

which this transformation initiates is important for accurate 

modeling of dome extrusion (Iverson and others, 2006), as 

well as for recognizing conditions that might lead to changes 

in the current eruptive activity.

Interpretation of the crystallization history of 2004–6 

Mount St. Helens dacite is complicated by abundant evidence 

of nonequilibrium crystallization, including (1) rapid late-stage 

growth of rims on plagioclase phenocryst (fig. 10); (2) spatial 

heterogeneity of groundmass textures (both vesicles and 

crystals; fig. 11); (3) the presence of two or, sporadically, three 

silica phases (fig. 12); and (4) local temperature increases in 

the groundmass as the result of latent heat released by rapid 

groundmass crystallization (Blundy and others, this volume, 

chap. 33; Pallister and others, this volume, chap. 30). Here 

we briefly review the evidence for each before discussing the 

likely importance of each process to gouge formation.

Phenocryst textures that suggest rapid late-stage growth 

include swallowtail-like extensions from the corners of large 

crystals, irregular growth boundaries, and skeletal overgrowths 

that either partially or completely trap residual melt (fig. 10). 

These irregularities preserve glass within the growth protrusions, suggesting entrapment at moderately high pressures (that 

is, during early phases of groundmass crystallization), perhaps 

because microlite nucleation was inhibited by local increases 

in H2

O rejected from the rapidly growing crystal. Therefore it 

appears that this crystallization, while rapid, was not directly 

responsible for cataclasis and gouge formation, although it does 

reflect the overall importance of decompression-induced crystallization during this remarkably steady effusive activity.

More important for determining conditions of marginal 

fragmentation is the spatial inhomogeneity of groundmass 

textures found within the early dome lavas. Lavas are heterogeneous in both groundmass crystallinity and vesicularity. Heterogeneity in groundmass crystallinity is particularly 

evident in the concentration of cotectic quartz-plagioclase 

intergrowths along euhedral phenocryst margins (fig. 11). 

However, as most of the groundmass within the cataclasite 

and gouge samples is holocrystalline, it is unclear what effect, 

if any, the heterogeneities in groundmass crystallization may 

have had on brittle deformation of the dome dacite. Heterogeneous vesicle textures are more likely to affect the material 

properties of the lava when deformed. A feature of many of 

the holocrystalline clasts within the gouge is the presence of 

numerous tiny microvesicles that are probably a feature of 

second boiling during the last stages of crystallization. It is 

possible that variations in vesicle structure may play a role in 

clast breakage, although our observations suggest that this, 

also, does not dominate the failure behavior.

Most important for gouge formation seems to be nearcomplete solidification of the ascending magma. The depth of 

this solidification may be estimated using both groundmass 

phases and crystal textures. Three silica phases occur within 

the matrix of different gouge clasts: cristobalite, tridymite, 

and quartz. Cristobalite typically forms large crystals; similar 

crystals in cryptodome dacite from the 1980 eruption were 

interpreted by Hoblitt and Harmon (1993) to result from 

vapor-phase crystallization. Quartz can form discrete groundmass crystals, but it most commonly grows as fine-scale cotectic aggregates with plagioclase and anorthoclase. Tridymite 

forms discrete acicular grains (fig. 12) that are common in 

some samples, sparse in others, and completely absent in some 

dome lava. In their study of Mount St. Helens lava samples 

from 1980–1986, Blundy and Cashman (2001) found tridymite 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 405

in a single lava sample erupted during the period of continuous slow effusion in 1983. They used experimental data to 

argue that equilibrium crystallization of tridymite in H2Osaturated, anorthite-free melts requires a minimum pressure 

of 11 MPa and temperature of 885oC. Addition of anorthite 

to the melt will increase the temperature of the high quartztridymite boundary. Subsequently Cashman and McConnell 

(2005) found additional examples of tridymite-bearing dacite 

from pyroclasts stored temporarily at shallow levels prior to 

explosive eruptions of Mount St. Helens during the summer 

of 1980. When analyzed for volatile content using FTIR, one 

such glassy sample preserved ~0.6 weight percent H2O, suggesting a minimum pressure of 5 MPa (A. Rust, oral commun., 

2005). Thus it appears that the presence of tridymite requires 

shallow (<25 MPa) but not surface (>5–10 MPa) pressures 

(depths between 200 and 1,000 m for an average magma density of 2,500 kg/m3

). Patterns of silica-phase precipitation thus 

suggest (1) substantial delays in quartz nucleation to produce 

the observed high degree of spatial variability; (2) varying P-T

paths during decompression, as illustrated by differences in the 

proportions of silica phases among samples; and (3) tridymite 

crystallization under a limited range of P-T conditions.

Experimental data (Hammer and Rutherford, 2002; 

Couch and others, 2003; Martel and Schmidt, 2003) further 

suggest that the common groundmass texture of holocrystalline quartz and feldspar intergrowths requires pressures >10 

MPa (400 m) and crystallization time scales of weeks. If 

magma ascends at a constant rate from a storage region at >4 

km (Rutherford and Devine, this volume, chap. 31), measured 

extrusion rates of ~7 m/d allow sufficient time for shallow 

crystallization (~3 months for ascent from 1,000 to 400 m). 

Pressures as low as 5 MPa were required to precipitate quartzalkali feldspar intergrowths in decompression experiments 

on Pinatubo magma compositions (Hammer and Rutherford, 

2002), whereas maximum rates of plagioclase nucleation 

typically require pressures between 25 and 5 MPa (Hammer 

and Rutherford, 2002; Couch and others, 2003; Martel and 

Schmidt, 2003). If gouge forms only after crystallization of the 

groundmass is nearly complete, these data suggest a shallow 

origin for the onset of brittle deformation (<400–500 m), consistent with the plug model of Iverson and others (2006). We 

note, however, that all of these estimates assume only minor 

cooling of magma during ascent, an assumption that appears 

reasonable given the high surface temperatures measured by 

FLIR (Schneider and others, this volume, chap. 17) and high 

Fe-Ti oxide temperatures from the dacite (Pallister and others, 

this volume, chap. 30).

Fault Gouge Formation

Extensive research on tectonic fault zones provides a 

framework for examining gouge formation during the extrusion of the 2004–6 Mount St. Helens spines. Tectonic fault 

zones commonly are characterized by narrow (millimeter to 

centimeter) slip planes composed of finely pulverized rock 

surrounded by broader cataclasite zones (typically meters in 

scale and locally foliated) mantled outward by a zone of less 

intensely deformed rock (for example, Chester and Logan, 

1986, 1987; Chester and others, 1993; Schulz and Evans, 

1998; Storti and others, 2003). These fault-zone characteristics are taken as evidence of strain-weakening behavior, with 

shear localization into narrow zones that act as the primary 

slip surfaces (for example, Ben-Zion and Sammis, 2003). 

Fault initiation in porous material commences with dispersed 

breakage, porosity reduction, and temporary strain hardening 

(for example, Mair and others, 2000; Cashman and Cashman, 2000). A change to strain-weakening behavior requires 

localization of shear through nucleation of interconnected 

slip surfaces (Biegel and others, 1989; Shipton and Cowie, 

2003). Average particle size decreases with increasing strain 

(Marone and Scholz, 1989; Mair and others, 2000; Hadizadeh 

and Johnson, 2003) until that strain can be accommodated by 

rolling (Morgan and Boettcher, 1999), contact creep (Prasher, 

1987), or development of new fault strands (if band formation 

results in local strain hardening; Mair and others, 2000).

In tectonic fault zones, gouge thickness shows a weak 

dependence on total displacement, a correlation that has been 

attributed to breakage of increasingly large asperities (Power 

and others, 1988). Sibson (1986) recognized three different 

types of natural fault gouge:

1. Attrition breccia resulting from wear. This gouge type 

is rare in tectonic faults, characterized by rolled clasts, 

and associated only with zones of significant fault slip.

2. Distributed crush breccia formed by destruction of 

local asperities. This gouge type is characterized by 

pervasive microfractures and jigsaw textures.

3. Implosion breccia. This gouge type has textures similar 

to those of distributed crush breccia.

Gouge within the cores of tectonic faults may contain 

ultrafine grains (micron to nanometer scale) whose origin 

has been variously attributed to quasi-static wear and attrition (Sammis and others, 1986; Chester and Chester, 1998), 

dynamic pulverization resulting from a single rupture event 

(Reches and Dewers, 2005; Wilson and others, 2005), and 

large strains (for example, Yund and others, 1990).

Our observations of textural characteristics of cataclasite 

and gouge samples from Mount St. Helens provide insight 

into fault gouge formation. Macroscopically, the structure of 

the fault zone along the conduit margin (for example, fig. 6) 

is similar to examples described in the fault literature, with an 

outer breccia that changes to cataclasite and, finally, into either 

fault gouge containing slickensides or areas of slip localization (figs. 8C, 16A) at the margin. Microscopically, localized 

extreme grain-size reduction along slickensides is consistent 

with experimental generation of ultrafine grains by large shear 

strains (Yund and others, 1990). Using the correlation for 

tectonic faults, a gouge thickness of 1–3 m suggests displacements of <1,000 m, in agreement with petrologic assessments 

of the depth of gouge formation presented above. Moreover, 
406 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Increasing fault displacement

Decreasing particle fragmentation

Increasing particle abrasion

Increasing fractal dimension

Figure 20. Diagram showing schematic change in size and 

shape of grains with increasing fault displacement. Although 

drawn for tectonic faults, this progression from coarse to fine 

grain sizes and angular to rounded shapes appears appropriate 

to fault behavior associated with spine emplacement at Mount St. 

Helens, Washington. Figure from Storti and others (2003).

these values suggest total strains (displacement/thickness) less 

than or about 500. Assuming that all of the slip is concentrated 

along the conduit margin, observed slip rates of ~8×10-5 m/s 

are orders of magnitude higher than those of creeping faults 

and orders of magnitude lower than slip events on stick-slip 

faults. These slip velocities suggest strain rates (velocity/thickness) of 3–8×10-5/s, intermediate to values typical for creeping 

and stick-slip tectonic faults.

The fault gouge itself appears to represent at least two of 

Sibson’s (1986) types. Extensive grain fragmentation adjacent to the slickenside zones (figs. 15B, 16D) fits Sibson’s 

description of distributed crush breccia, in that individual 

pieces of a grain have not been displaced from each other by 

shear. However, it is clear that these fragments are eventually 

distributed along the foliation bands by shear displacement 

after crushing (figs. 15D, 16B, 16C). In contrast, the common 

rounded fragments in cataclasite and gouge (figs. 14B, 14C) 

appear to be attrition breccia, with rounding that results from 

abrasion and wear during shear transport. The prevalence of 

these attrition features may reflect not only the large total 

displacements (shear strain) involved with dome emplacement but also the steadiness of movement. Whereas sporadic 

movement of tectonic faults allows healing or lithification of 

fault zones between slip events, steady movement of solidified 

magma along the conduit margin apparently prevents such 

healing. Instead, the observed gradual decrease in the percentage of larger clasts apparently allows them to be isolated from 

each other by the matrix, thus permitting individual clasts to 

roll and abrade. Continued brittle deformation of fragmented 

particles results in further grain-size reduction and shear localization along multiple slickensides.

An interesting discussion in the fault zone literature 

involves the mechanism by which fragmentation occurs. 

Patterns of grain-size reduction may be characterized by the 

fractal dimension D of particle grain size, where D relates 

cumulative particle frequency N(S) (the number of particles 

less than S, usually measured per volume) to size (S):

N(S) ≈ S −D. (1)

Thus D describes the relative proportion of small and large 

fragments; higher D values reflect samples where the grainsize distribution is dominated by smaller particles. Within this 

framework, the “pillar of strength” model (Allegre and others, 

1982) predicts D = 1.97, the “plane of fragility” model (Turcotte, 1986) predicts D = 2.84, and the constrained comminution model (Sammis and others, 1986; Sammis and others, 

1987; Sammis and Biegel, 1989; Turcotte, 1992) predicts D = 

2.6. The latter model is based on sequential breakage of equalsize particles into a discrete (and relatively small) number 

of particles. Key to the constrained comminution model is 

that fragments of similar size are evenly distributed in space. 

Experimental studies of sheared quartz sand support a model 

of constrained comminution under conditions of high confining pressure (100 MPa) and particle size range of 10 μm to 

1 mm (Biegel and others, 1989; Marone and Scholz, 1989). 

However, data from natural cataclastic samples are less easily 

interpreted (for example, Blenkinsop, 1991; An and Sammis, 

1994; Storti and others, 2003; Cashman and others, 2007). 

In many fault zones, D values change from low in brecciated 

damage zones to high in intensely comminuted narrow shear 

zones (fig. 20). These high D values (high degrees of crushing) may result from slip-enhanced surface abrasion during 

shear transport, as suggested by Hooke and Iverson (1995) for 

glacially abraded sediment. High degrees of crushing also create the matrix support required to allow rolling and abrasion of 

remaining larger particles.

The fragmentation schematic shown in figure 20 provides 

a description of the process of gouge development at Mount 

St. Helens. Imaging data, which can be used to characterize 

the larger clast sizes in the cataclasite, yield fractal dimensions 

~1.97 for particles from 10 to 1,000 μm in diameter (fig. 18B) 

when measured as number per area; this is equivalent to D ~ 

2.97 on a volume basis. This high D value is identical to that 

measured by Chester and others (2005) in cataclasite and ultracataclasite of the Punchbowl Fault, a large-displacement fault 

in the San Andreas Fault system. Sieve data (fig. 17), however, 

show that grain size data are self similar (fractal) only within 

limited size ranges. As seen in other fault rocks, fits to straightline segments suggest that D increases with decreasing grain 

size (for example, Blenkinsop, 1991; An and Sammis, 1994; 

Storti and others, 2003). High D values at very small grain sizes 

result from extensive comminution of individual clasts, as evidenced by the shattered large crystals and dome fragments that 

border the slickensides (figs. 15B, 16D). Thus we interpret the 

grain-size data to reflect (1) cataclasis and particulate flow over 

most of the observed gouge thickness and (2) strain localization 

and intense comminution within the slickensides, which are 

themselves distributed within the outer gouge zone.
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 407

The frictional properties of 2004–6 Mount St. Helens 

fault gouge have been investigated by Moore and others (this 

volume, chap. 20) at room temperature, shear rates of 1.5–

5,000×10-6 m/s, and low normal stresses (5–195 kPa). They 

find rate-weakening behavior at shear rates <1×10-4 m/s and 

rate-strengthening behavior at >5×10-4 m/s, with stick-slip 

oscillations observed at the highest normal stress and lowest 

shear rates. However, although they observe shear localization 

into a narrow (1 mm) zone, they find no evidence of comminution or crushing of gouge, perhaps because the small 

cumulative strains and low normal stresses in the experiments 

were insufficient to produce the grain breakage observed in 

natural samples.

In summary, the unusually well developed fault gouge 

that has mantled all of the 2004–6 Mount St. Helens spines 

provides evidence of profound change in the rheological 

behavior of dome lava during ascent as a consequence of nearcomplete crystallization and subsequent brittle breakage and 

gouge formation. The effect of this extensive brittle deformation on dome extrusion is less clear. Iverson and others (2006) 

suggest that rate-weakening behavior is required to produce 

the stick-slip behavior suggested by the repetitive (drumbeat) 

earthquakes. Alternatively, Mastin and others (this volume, 

chap. 22) suggest a gouge zone that is, overall, rate strengthening but that fails in a patchwork fashion during extrusion. 

Field evidence for localized failure includes both observed 

Reidel shears in the cataclasite and in throughgoing ultracataclasite (slickenside) bands within the gouge.

Comparison with Other Dome Eruptions

The morphology and emplacement conditions of other 

recent domes and spines of intermediate composition are 

reviewed above. Here we use observations of the 2004–6 

activity at Mount St. Helens to evaluate processes that may 

have affected extrusive behavior at other volcanoes.

Dome emplacement during the 1980–86 eruption of 

Mount St. Helens provides the most obvious comparison 

with current (2004–6) activity. Importantly, although the 

erupted composition in the 1980s was similar to that of the 

2004–6 dacite and average effusion rates were comparable (or 

slightly lower), the 1980–86 magma had a higher temperature 

(880–900°C in the early 1980s compared with 840–850°C in 

2004–6), and its extrusion was dominantly episodic rather than 

continuous. Spines were produced on only two occasions: (1) 

accompanying endogenous growth during continuous effusion in 1983 and (2) following a protracted, 17-day episode 

of (mostly endogenous) dome-lobe formation in May 1985. 

Backscattered electron images of a sample from the February 

1983 spine look similar to early samples of the 2004 spine 

(compare fig. 21A with fig. 11C). In both samples the distribution of groundmass crystals is heterogeneous, as is the 

distribution of clean glass. In contrast, a marginal sample of 

the May 1985 spine is holocrystalline (fig. 21B), with the characteristic cotectic crystallization fabric seen in the more slowly 

extruded 2005–6 spine samples.

However, there are two differences between the spine 

sample from May 1985 and those of the current eruption. First, 

the extremely fine-scale “myrmekitic” textures occupying 

interstices between microlites in the May 1985 sample (and not 

observed in the 2004–6 spine samples) suggest that preeruptive 

crystallization occurred at shallow levels, most likely within the 

dome itself. Second, the boundary between the extruded spine 

and the dome lobe had a different character in 1985. In contrast 

to the thick gouge zones described above, the spine-dome 

lobe boundary in 1985 comprised a thin (~10 mm) cataclasite 

zone that, although showing some of the grain-size reduction characteristics of the recent gouge, was narrower and less 

A

B 

50 μm

50 μm

Figure 21. Backscattered electron images showing 

groundmass textures of 1980s lava spines of Mount St. Helens. 

Textures are similar to those observed in spines produced 

during current Mount St. Helens eruptive activity; for example, 

compare with figures 11C, 11D, 12A, or 12B. A, February 1983. 

B, May 1985.
408 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1 cm

Figure 22. Full thin-section (4.5 cm x 2.5 cm) scan of sample 

collected across the margin of the May 1985 dome lobe (right) and 

spine (far left). The darker area of reduced phenocryst size toward 

the left side of the figure is a thin (<1 cm) cataclasite zone that 

marks the lava-spine boundary.

comminuted (compare thin-section image of fig. 22 with those 

of fig. 9). These features suggest that the May 1985 dome was 

formed by late-stage solidification within the growing dome, 

with extrusion resulting in only small cumulative displacements 

(strain), consistent with its small size.

More similar are larger spines produced at Mont Pelée 

(1903), Unzen (1991–95), and Soufrière Hills (1995–present), which have macroscopic features similar to the 2004–6 

spines at Mount St. Helens, including smooth curved forms 

and pronounced striations parallel to the extrusion direction 

(reviewed above). Descriptions in the literature suggest that 

gouge was present on the striated surface of the giant Mont 

Pelée spine (Jaggar, 1904); gouge and cataclasite also mantle 

some large dome fragments from Soufrière Hills, Montserrat. 

Conditions leading to spine formation also appear similar at all 

of these locations. Most important are estimated ascent rates 

of 1–3×10-4 m/s, rates that are sufficiently slow for the extensive crystallization necessary for both plug and fault gouge 

formation (see above). The comparison between the 1985 and 

current spines at Mount St. Helens further suggests that formation of large spines requires that magma solidify well within 

the conduit. This depth of solidification, in turn, means that 

solid dome rock must be transported large distances before 

exiting the vent, with resulting large strains responsible for the 

development of thick gouge zones.

From Dome to Dust

Brittle deformation is not unique to holocrystalline spines 

of intermediate composition. Numerous small-scale brittle 

deformation features have been described along shallow conduit margins and within basal shear zones of glassy obsidian 

domes, including anastomosing tuffisite veins and fractures, 

cataclasite zones, and trails of broken phenocrysts (for example, Stasiuk and others, 1996; Tuffen and Dingwell, 2005). 

Fracture and grain-size reduction in these systems have been 

explained by shear-strain rates at conduit margins exceeding 

experimentally determined critical rates for shear-induced 

fragmentation (for example, Dingwell and Webb, 1989) of 

high viscosity (1010–1012 Pa s) melts at the low strain rates 

(10-3–10-5/s) anticipated for dome extrusion (Gonnermann and 

Manga, 2003; Tuffen and others, 2003).

Deformation textures in obsidian domes, however, differ 

dramatically from those described here for the Mount St. Helens gouge zone. First, fracture networks in obsidian are limited to thin (<20 mm) fault and injection veins with limited to 

no displacement and thicker, irregular “injection voids” (≤80 

mm in width; Tuffen and Dingwell, 2005). The largest fault 

complex observed by Tuffen and Dingwell (2005) is ~5 m in 

length, with a maximum offset of 0.13 m. Thus these features 

are substantially smaller than the zones of breccia, cataclasite, 

and gouge that line the shallow Mount St. Helens conduit, 

with the possible exception of basal breccia zones in some 

obsidian flows (Gonnermann and Manga, 2003). Moreover, 

both tuffisite veins and basal breccias show abundant evidence of annealing and ductile deformation (Gonnermann and 

Manga, 2003; Tuffen and others, 2003; Tuffen and Dingwell, 

2005). In contrast, gouge zones at Mount St. Helens are dominated by brittle fracture and cataclastic flow, with no evidence 

of re-annealing between fracture events. Finally, fracture 

zones within obsidian domes appear to lack the distinct zones 

of breccia, cataclasite, and gouge seen at Mount St. Helens; 

instead, fragment sizes within the tuffisite veins are very small 

(much less than 1 mm), with no clasts of intermediate size.

What are the implications of these differences in fragmentation style? Annealing textures preserved within tuffisite 

veins and basal breccias of obsidian flows result from the 

strain-dependent rheology of viscous silicate melts; at the 

same temperature, the melt will flow at low strain rates but 

will fracture at higher strain rates. For this reason, a model 

of flow emplacement by repeated fracture and healing of the 

same magma has been suggested for obsidian (Tuffen and 

others, 2003; Tuffen and Dingwell, 2005). This model has 

been extended to explain repetitive seismic events accompanying andesitic dome growth at Soufrière Hills, Montserrat 

(for example, Neuberg and others, 2006). Our observations of 

fault gouge formation, however, would suggest caution in the 

direct application of deformation mechanisms in obsidian to 

the interpretation of seismicity associated with the emplacement of highly crystalline lava. Most importantly, we see no 

evidence of annealing within the gouge zone between fracture 

events. Instead, fault gouge forms by fracture and fragmentation of crystals and holocrystalline matrix into successively 

smaller fragments. For this reason, we suggest that the origin 

of the repetitive (drumbeat) earthquakes at Mount St. Helens most likely lies in progressive fragmentation of dome 

dacite rather than by process of repeated fracture and healing. 

Although definite assignment of individual slip features to 

specific earthquake characteristics is beyond the scope of this 

paper, we note that several types of slip zones are apparent 
19. From Dome to Dust: Shallow Crystallization and Fragmentation of Conduit Magma During the 2004–2006 Dome Extrusion 409

at the outcrop to hand-specimen scale, from Reidel shears in 

the cataclasite (fig. 6) to slickensides in the gouge (fig. 8C). 

If individual drumbeat earthquakes represent slip events of ~5 

mm (Iverson and others, 2006), then they may represent the 

composite effect of fracturing multiple smaller grains (as suggested by the dominance of fine (<10–20 μm) material within 

the cataclasite and gouge zone. Alternatively, perhaps only the 

initial fragmentation of larger dome fragments is seismogenic.

This leads to a final point—the role of fragmentation 

during dome explosions. A common assumption is that dome 

explosions result from sudden decompression of pressurized dome interiors and resultant brittle fragmentation (for 

example, Alidibirov and Dingwell, 1996). However, Rowe and 

others (this volume, chap. 29) demonstrate that ash released 

during the two substantial dome explosions associated with 

the current eruption (in January and March of 2005) bears a 

striking textural resemblance to the fault gouge described here. 

The implications of this observation are that the explosions 

themselves transported, but did not generate, much of the ash 

emitted during each event, an interpretation consistent with the 

source of the explosions at the junction of the spine and the 

conduit rather than from within the dome itself. We suspect 

that gouge formed along conduit margins may commonly 

supply ash to explosions that accompany spine growth (for 

example, Jaggar, 1904), a hypothesis that suggests that further 

studies of the physical properties of volcanic gouge (especially 

relations between deformation behavior and permeability) 

could improve our understanding of hazards related to the 

slow extrusion of lava spines.

Conclusions

We have documented field, hand-sample, and thinsection evidence for cataclasis and gouge formation along 

the margins of the smooth-surfaced spines extruded from 

Mount St. Helens from 2004 to present. These observations 

show brittle breakage and subsequent cataclastic flow to be 

important in the slow extrusion of intermediate-composition 

hydrous magma. Detailed textural observations document 

both crystallization accompanying magma ascent (decompression) and the sequential fragmentation of competent 

dome rock to produce fine-grained fault gouge. Taken 

together, these observations provide new constraints on interactions between rates of magma ascent, degassing, crystallization, and brittle fragmentation, interactions that control the 

dynamics of lava extrusion.

The groundmass of dacite lava samples is variably crystalline, with the extent of groundmass crystallinity apparently 

increasing with time (decreasing mass eruption rate). Most, 

if not all, of the crystallization is driven by decompression 

and degassing rather than cooling; Fe-Ti oxide temperatures 

suggest that the extensive crystallization may even help to heat 

the magma (for example, Blundy and others, 2006; Pallister 

and others, this volume, chap. 30). Within the gouge zone, 

dome fragments are holocrystalline, indicating solidification 

along conduit margins prior to brittle fragmentation. Detailed 

textural observations show that the transformation of competent dacite to fine powder occurred by sequential breakage 

of rock fragments to form zones of breccia, cataclasite, and 

gouge. Further concentration of slip along narrow slickenside planes produced thin (1 mm) zones of pulverized and 

foliated ultracataclasite. Both the extreme grain-size reduction along these slip zones and the shattering of individual 

grains observed adjacent to narrow slip surfaces are similar 

to textures observed in near-surface gouge produced by large 

tectonic faults (for example, Cashman and others, 2007) and 

suggest large strains. This inference is reasonable given a 

solidification depth of 400–500 m, above which deformation apparently is concentrated along the conduit margin (for 

example, Iverson and others, 2006). However, we see no 

evidence of subsequent annealing, as is commonly observed 

in tuffisite veins and obsidian flows (for example, Tuffen and 

others, 2003; Gonnermann and Manga, 2003), suggesting 

that earthquake-generating mechanisms based on repeated 

fracturing and healing do not apply to the extrusion of highly 

crystalline domes. Additionally, the two largest explosions at 

Mount St. Helens (in January and March of 2005) occurred at 

the contact between the spine and conduit wall and ejected ash 

that appears to have originated as fault gouge.

The physical characteristics of dome extrusion at Mount 

St. Helens from 2004–6 can be generalized to other recent 

examples of andesite and dacite extrusions. As surmised by 

Williams (1932), the complete decompression-driven solidification prior to magma extrusion necessary for spine formation requires magma ascent to be slow relative to the kinetics 

of crystallization and the dynamics of gas loss. Spine formation appears limited to magma ascent rates of <1–5×10-4 m/s, 

which allows several weeks for magma to transit from ~1,000 

m (the point at which crystallinity increases dramatically as 

a function of decreasing pressure; Blundy and others, 2006) 

to ~400 m, where extensive cotectic crystallization of quartz 

and feldspar occurs. Comparison of marginal shear behavior 

of a small spine produced at Mount St. Helens in 1985 and 

the large spines produced from 2004 to 2006 also suggest a 

further point—that formation of large spines requires solidification to occur deep within the conduit, a condition that leads 

to extensive brittle deformation and gouge formation during 

transit to the surface.

Acknowledgments

The authors would like to acknowledge all CVO personnel who helped with field observations and sample collection, 

as well as a host of volunteers who helped to sort and catalogue the samples. We thank reviewers Rob Watts and Nick 

Beeler for their thoughtful comments on the first draft of the 

manuscript. Cashman acknowledges support from NSF EAR0207362 and EAR-0510437.
410 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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414 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 20

1

 Department of Geological and Atmospheric Sciences, Iowa State 

University, Ames, IA 50011

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Frictional Properties of the Mount St. Helens Gouge

By Peter L. Moore1, Neal R. Iverson1, and Richard M. Iverson2

Abstract

Frictional properties of gouge bounding the solid dacite 

plug that extruded at Mount St. Helens during 2004 and 2005 

may have caused stick-slip upward motion of the plug and 

associated seismicity. Laboratory experiments were performed 

with a ring-shear device to test the dependence of the peak 

and steady-state frictional strength of the gouge on shearing rate and hold time. A remolded gouge specimen (~0.012 

m3) was sheared under constant normal stresses ranging from 

5 to 200 kPa and at rates ranging from 10-6 to 10-3 m/s. The 

gouge exhibited rate-weakening behavior at rates lower than 

1×10-4 m/s and rate-strengthening at rates above 5×10-4 m/s. 

Peak strengths occurred during the onset of shearing, when 

displacements were generally less than 0.5 mm. In slide-holdslide tests, the peak strength of the gouge increased logarithmically as hold times increased from 3 s to almost 105 s.

Rate-weakening friction is a requirement for stick-slip 

behavior that is satisfied by the Mount St. Helens gouge. 

Indeed, regular stick-slip oscillations were observed in two 

experiments performed at the highest normal stress and lowest rates of shear. The conditions under which this stick-slip 

motion occurred indicate that the gouge also satisfies a second 

criterion for stick-slip behavior of materials exhibiting rateand-state dependent friction—gouge stiffness exceeds that of 

the ascending magma that drives upward motion of the plug. 

The presence of highly compliant magma as a driving element 

may be crucial for generating stick-slip instabilities at the shallow earthquake focal depths observed during the eruption.

Introduction

Lava dome formation during the 2004–2005 eruption 

of Mount St. Helens was accompanied by abundant, nearly 

periodic, shallow-focus earthquakes. These “drumbeat” 

earthquakes had magnitudes <2, focal depths <1 km, and they 

typically recurred about every 1–2 minutes. This seismicity 

is thought to have resulted from incremental uplift of a solid 

dacite plug driven by magma ascent from below at rates of 

roughly 1–2 m3/s (Iverson and others, 2006). Wear along the 

margin of the plug produced a layer of crushed rock, or gouge, 

of the same lithology as the extruding rock (fig. 1). Observed 

at the surface, this gouge formed a coating 0.1–1 m thick 

on freshly exposed faces of the lava dome. Moreover, these 

freshly exposed faces were conspicuously striated in directions mostly parallel to the direction of plug motion, indicating 

that gouge deformation probably accommodated much of the 

upward displacement of the plug mass. Interpretation of the 

source of drumbeat seismicity, therefore, requires knowledge 

of the frictional properties of this gouge.

Gouge is generated by the wear and cataclasis of rock 

during frictional shear and develops progressively with 

shear displacement (for example, Engelder, 1974). Although 

cataclasis may occur episodically in rapidly strained magma 

near its solidus (for example, Tuffen and Dingwell, 2005), 

granular gouge dredged from the margin of the new Mount St. 

Helens lava dome early in 2005 was cool, noncohesive, and 

bore evidence primarily of mechanical wear and comminution. Therefore, the frictional behavior of the gouge is likely 

similar to that of low-temperature gouge in shallow crustal 

faults. Consequently, frictional behavior observed in roomtemperature laboratory experiments is relevant to interpreting 

dome extrusion processes during the 2004–2005 eruption of 

Mount St. Helens.

Theoretical Background

Rate- and State-Dependent Friction

Gouge behavior closely approximates that of a Coulomb material, such that its shear strength  is described by 

 = μ ́

n

+ c, where  ́

n

 is the effective normal stress, c is cohe-
416 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 1. Photograph of striated gouge near base of extruding spine at Mount St. 

Helens, July 28, 2005. Field of view approximately 20 m. 

USGS photo by S.P. Schilling.

sion and μ is the coefficient of friction. In most granular media 

at low temperatures, μ can be considered nearly constant, but 

laboratory experiments have routinely shown that μ is slightly 

sensitive to both shear rate and the history of deformation. 

These deviations from a constant coefficient of friction are 

small but of great interest because under the appropriate conditions they give rise to an instability that may be responsible 

for earthquakes. In fault mechanics literature, these shear-rate 

and deformation-history effects are termed “rate” and “state” 

dependencies, respectively (see Marone, 1998, for a review).

Stemming from the work of Dieterich (1979) and Ruina 

(1983), the most common expression of the rate- and statedependence of friction is

, (1)

where μ0 is a reference coefficient of friction applicable at a 

reference shear displacement rate V0

, V is shear displacement 

rate (with dimensions of length/time),  is a state variable (with 

dimensions of time), Dc is a characteristic slip distance over 

which friction adjusts to a steady state in response to a velocity 

change, and a and b are positive coefficients that are typically of 

order 10-3 to 10-2. Individually, the parameters a and b describe 

the relative magnitude of the so-called “direct” and “evolution” 

effects, respectively. The direct effect is a near-instantaneous 

response of friction to a change in shear rate, whereas the 

evolution effect is a transient response that takes place over the 

characteristic length scale Dc. Note that in this paper, we use the 

term “shear rate” to denote shear displacement rates such as V. 

Where we wish to distinguish such rates from shear strain rates, 

which have dimensions of inverse time, we do so explicitly.

The state variable  in equation 1 describes a time-dependent effect that can be viewed as the length of time that a set 

of grain contacts persists at a given velocity (Scholz, 2002). 

According to Dieterich (1979), the evolution of  with time 

may be described by

1

c

d V

dt D

    = −    

. (2)

Although the characteristic slip distance can be determined 

from stress-displacement data from experiments with bare 

rock, it is difficult to determine in experiments with gouge 

because this displacement depends on the strain distribution in 

the gouge (Marone and Kilgore, 1993). In a steady state 

( d d t = 0 ),  = Dc /V and equation 1 reduces to

, (3)

where μss is the steady-state friction coefficient. The sensitivity of steady-state frictional strength to shear rate, therefore, is 

given by the difference between the coefficients a and b for the 

direct and evolution effects:

. (4)

If this difference is negative, the material weakens with 

increasing slip rate, and conversely, if a − b is positive, the 

material strengthens with increasing slip rate (Scholz, 2002). 

The coefficients a and b are usually determined experimentally 

by conducting velocity-stepping triaxial or direct-shear experi-















 +













 +=

c

ln l

D

V b

V

V

μμ a 0

0

0 n

















 = + −

0

0 V

V ( ba ) ss   ln

( ) VV a b

0

− 0 =− ss

ln

 
20. Frictional Properties of the Mount St. Helens Gouge 417

ments in which shear rate is varied over a wide range during 

continuous shear.

Consistent with Dieterich’s original reasoning (Dieterich, 

1972, 1978, 1979), Beeler and others (1994) demonstrated 

that the parameter b can also be determined independently by 

measuring the variation in peak friction as a function of the 

time that shear is stopped (“hold time”), such that

. (5)

Peak shear strength typically increases logarithmically with 

time as sliding surfaces are held in stationary contact. The 

mechanisms responsible for this aging effect are thought to be 

associated with compaction and growth of the area of solidto-solid grain contact across the shear zone (Losert and others, 

2000; Scholz, 2002).

Relation Between Gouge Frictional Behavior 

and Seismicity

If gouge becomes weaker as shear rate increases 

(a - b < 0), slip instability and stick-slip motion are possible, but according to the rate-and-state dependent friction 

model, rate weakening alone is not sufficient to induce 

stick-slip behavior (Rice and Ruina, 1983; Ruina, 1983). 

The stiffnesses of the gouge and the system that loads the 

gouge are also important. The decline from peak (static) 

frictional strength to steady-state (dynamic) strength of 

the gouge requires a finite displacement that exceeds but is 

proportional to the characteristic slip distance, Dc. If elastic 

strain in the materials loading the shear zone is relaxed over 

a displacement larger than Dc, the applied driving stress 

momentarily exceeds the resisting stress and slip accelerates until driving and resisting stresses become equal (fig. 

2). As defined by Ruina (1983), the “critical stiffness” of 

the gouge, kcrit, is the ratio of shearing-resistance decline 

to characteristic slip distance upon a change in slip velocity, and kcrit has dimensions of stress per unit length. If the 

loading system requires a slip distance much greater than 

Dc to reduce the applied stress from the peak to steady-state 

strength, the loading system’s stiffness is smaller than kc and 

stick-slip behavior is possible (fig. 2).

According to the arguments summarized above, repetitive stick-slip motion along the subsurface margins of the 

extruding dacite plug at Mount St. Helens can occur if 

two conditions are met—(1) the gouge must exhibit rateweakening behavior at the imposed loading rate, and (2) the 

most compliant component of the loading system (that is, the 

magma) must have a stiffness smaller than the critical stiffness of the gouge. The first requirement is readily evaluated 

with our laboratory experiments. The second requirement is 

more difficult to address, but our experiments and inferences 

about the magma-plug system at Mount St. Helens provide 

some constraints.

)( hold

peak

td

d

b = ln



|slope| = kcrit

|slope| = ksys

Dc

DISPLACEMENT

SHEAR STRESS

A

B

C

Rapid slip

Applied shear stress

Shear stress

supported by gouge

Experimental Methods

Gouge was obtained by dredging the surface of the 

newly extruded spine at Mount St. Helens by helicopter in 

February 2005. Ring-shear tests were performed at room 

temperature with dry, remolded specimens. Particles with 

diameters larger than one-tenth the smallest dimension of the 

specimen chamber (70 mm) were removed from the gouge, 

consistent with normal geotechnical testing procedure (Head, 

1989). Shearing rates and normal stresses were varied among 

tests through the ranges permitted by the device (5×10-3 to 

1.5×10-6 m/s and 5 to 195 kPa respectively). These experiments differed from velocity-stepping tests (for example, 

Blanpied and others, 1987; Biegel and others 1989; Marone 

and others, 1990) in that shear-rate changes were made only 

after first stopping shear.

The ring-shear device (detailed in Iverson and others, 

1997) has been used to study the mechanical properties of a 

variety of granular geological materials (for example, Iverson and others, 1998; Iverson and others, 2000; Moore and 

Iverson, 2002; Scherer and others, 2004). Its large capacity 

permits inclusion of larger particles that are common in gouge, 

but the test specimen’s large surface area limits the magnitude 

Figure 2. Stress applied by the loading system and shear 

stress supported by gouge during a stick-slip cycle. When 

slip commences at A, strength of the gouge initially declines 

more rapidly than the stress applied by the loading system, 

and acceleration occurs. When driving and resisting forces 

are equal again at B, slip begins to decelerate and stops at 

C. Slope of the decline in applied stress with displacement 

is the stiffness of the system (ksys). When this stiffness is 

smaller than the critical stiffness of the gouge kcrit as defined 

by the shear stress decline and the gouge’s characteristic 

slip distance Dc

, the system is susceptible to stick-slip. Dc is 

the characteristic distance of exponential decay of gouge 

shearing resistance during rapid slip (after Scholz, 2002).
418 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

of the normal stress that can be applied. The ~70-mm-thick 

specimen rests in an annular chamber (annulus OD = 0.600 m, 

ID = 0.370 m) and is gripped at its base by a toothed platen 

(fig. 3). This platen is screwed to a base plate, and it and the 

lower walls bounding the specimen rotate. The toothed lid 

of the specimen chamber is prevented from rotating with the 

base by a pair of stationary, horizontally oriented load cells, 

causing the specimen to shear. The force exerted by the lid on 

these load cells is approximately the shear force supported by 

the specimen. Shear is focused in a lens-shaped zone that is 

10–30 mm in thickness at the centerline of the specimen and is 

centered on the sliding interface between the upper and lower 

walls (fig. 3B) (for example, Hooyer and Iverson, 2000). A 

stress normal to the shearing direction is applied with a lever 

system that presses downward on the lid (fig. 3A). The lid is 

free to move vertically if the porosity of the specimen changes, 

and this vertical displacement is measured continuously with 

three equally spaced linear variable displacement transducers 

(LVDTs) that press on the perimeter of the lid.

Vertical and circumferential frictional forces between 

the gouge and the upper walls bounding the specimen are 

measured independently with a load cell and torque sensor, 

respectively, which are contained within a suspension called 

the “yoke” (fig. 3A). The yoke holds the upper walls stationary 

and is not coupled to the lid. The measured forces on the walls 

are used to correct both the normal stress applied to the shear 

zone, which is reduced by upward wall drag in a consolidating 

specimen, and the shear resistance on the lid, which is reduced 

due to circumferential drag between the specimen and walls. 

However, owing to extrusion of fine sediment along the interface between the upper and lower walls, a component of the 

measured circumferential wall force is due to friction between 

the upper and lower walls rather than between the upper wall 

and specimen. Thus, when applied normal stress is small, this 

friction represents an unknown fraction of the wall correction 

necessary to estimate the total resistance to shear. As a result, 

in experiments with the smallest applied normal stresses (5–23 

kPa), uncertainty in the measured shear stresses and associated friction coefficients is large. Despite the uncertainty, 

these low-stress measurements are, nevertheless, included to 

qualitatively illustrate patterns that are available only from the 

low-stress tests.

Two variable-speed electric motors (1/17 and 1/2 hp) 

were used to turn the base and thereby shear the gouge specimen. Used in combination with either one or two gearboxes, 

these motors enabled two ranges of shear rate—1.5×10-6

–2.5×10-5 m/s and 5.0×10-4–5.0 ×10-3 m/s. The higher-power 

motor coupled to a single gearbox allowed shearing at the 

higher range of rates, but this configuration could not deliver 

sufficient torque to shear the gouge under normal stresses 

greater than 23 kPa. When the smaller motor was used with 

both gearboxes normal stresses of as much as 200 kPa could 

be applied. At higher normal stresses this motor/gearbox configuration was not sufficient to turn the rotating base.

Two groups of experiments were conducted. In one group 

designed to investigate the rate dependence of gouge friction, 

Load cell (1 of 2)

Normal-load plate

Lever arm To weights

Yoke 

Load cell 

Rotating base

60 cm

Gear box

Variable-speed

 motor

LVDT

(1 of 3)

A

B

Torque

sensor

Specimen

Rotating lower wall 

and platen

Upper wall 

Upper platen Specimen

Lens-shaped

shear zone

Figure 3. Ring-shear device. Gouge specimen is shown in crosshachure. A, Diagrammatic cross section. Dark-gray parts rotate 

when motor is on. A second gearbox between motor and large 

gearbox was used in the experiments conducted at low shear 

rates (1.5×10-6 – 2.5 ×10-5 m/s). LVDT, linear variable displacement 

transducers. B, Oblique sketch of the specimen chamber. 

Shearing in gouge typically is focused in a lens-shaped zone, 

shown in black, centered on interface between stationary upper 

wall and rotating lower wall.

shear was driven at a constant rate until shearing resistance 

became steady, usually after 1–10 mm of total centerline displacement. Shear was then stopped to adjust the motor speed 

and restarted under the same normal stress, but at a new rate, 

until shearing resistance again became steady. This process 

was repeated at multiple shearing rates, and whenever possible 

tests were repeated in sequences of increasing and decreasing shear rate. In a second group of experiments, periods of 

shear, which were terminated once shearing resistance became 

steady, were separated by intervals of hold (no shear) that 

ranged from 3 seconds to more than one day. In these experiments, shear rate and normal stress were held constant so that 

the effects of hold time on peak friction could be isolated. 
20. Frictional Properties of the Mount St. Helens Gouge 419

After some of the experiments, the gouge was sieved to 

measure whether there had been a size reduction of the largest 

grains; no size reduction could be detected indicating that little 

comminution or crushing occurred in the experiments.

Results

Data obtained in two typical ring-shear experiments, conducted at different shear rates under an applied normal stress of 

159 kPa, are shown in figure 4. Following initial stress peaks 

related to the “direct effect” of changing shear rate, shear stress 

became approximately steady after displacements of less than 

10 mm. The key feature illustrated by these two data series is 

that the steady-state shear strength of the gouge was smaller 

at the faster shear rate. For rates smaller than 10-4 m/s, such 

rate weakening was always observed, except for experiments 

conducted at a normal stress of 23 kPa in which the shear-rate 

dependence was neutral or slightly positive, depending on how 

the data are regressed (fig. 5). In the two sets of experiments 

conducted at faster shearing rates and low normal stresses, rate 

strengthening was observed (fig. 5B). Unfortunately, because of 

the torque limitations of our motor/gearbox configurations, the 

shear rate at which the transition to rate strengthening occurs 

cannot be pinpointed, but it appears to be on the order of 10-4

m/s for the range of stresses considered. As described previously, measurements at the lowest applied normal stresses are 

uncertain, but the consistently larger apparent friction coefficient at normal stresses of 5 kPa may indicate that the gouge 

has some cohesion. (Cohesion of about 1 kPa would decrease 

the apparent friction coefficient at 5 kPa to about 0.4 or 0.5, 

comparable to values measured at higher normal stresses.) 

Alternatively, the larger friction coefficient at 5 kPa may reflect 

steepening of the Coulomb failure envelope that is commonly 

observed at low normal stresses in various laboratory tests on 

granular materials (Lambe and Whitman, 1969).

68

67

69

70

71

72

73

0 0.002 0.004 0.006 0.008

CENTERLINE DISPLACEMENT, IN METERS

SHEAR STRES, IN KILOPASCALS

Faster (3.0 x 10-6 m/s)

Slower (1.5 x 10-6 m/s)

Figure 4. Stress-displacement data comparing two 

experiments conducted at normal stress of 159 kPa but at 

different shear rates. Prominent peaks in curves at ~0.5 mm 

of displacement are peak shear stresses, after which shear 

stress falls to steady-state value in both experiments.

Values of the rate-dependence parameter a–b were 

estimated for each normal-stress level using equation 4 and 

a reference shear rate of 1.5×10-6 m/s, the smallest shear rate 

used in the experiments (table 1). Each value of a–b was 

obtained from the slope of the linear regression of the steadystate friction coefficient against ln(V/V0

). Data from the “slow” 

(1.5×10-6 m/s to 2.5×10-5 m/s) and “fast” (5×10-4 m/s to 5×10-3 

m/s) experiments were regressed separately. Values of a–b at 

slow rates and high normal stresses (86–195 kPa) range from 

−0.0013 to −0.0103, indicating various degrees of rate weakening (fig. 5A).

In most of the ring-shear experiments, after a few millimeters of displacement, the shearing resistance of the gouge 

did not vary systematically, much as in the examples shown 

in figure 4. However, in the two experiments conducted at the 

largest normal stress and lowest shear rates, no steady shear 

stress developed. Instead, shear proceeded by regular stickslip episodes (fig. 6). Major slip events were accompanied 

by abrupt consolidation, whereas dilation appeared to occur 

during “stick” intervals. Apparently the normal stresses and 

shear rates of these two experiments produced gouge stiffness 

and rate weakening that met both requirements for stick-slip 

behavior outlined previously.

In two sets of slide-hold-slide tests at a single shear rate 

and normal stress, peak strength increased logarithmically 

with hold time (fig. 7). Fitting these data to equation 5 indicates that b = 0.0079 (table 1). Hold time was also recorded 

during the variable-rate tests and, although these data have 

more scatter (perhaps due to a dependence of b on shear rate), 

they indicate a b = 0.0075. During hold periods compaction 

of the gouge was generally not observed, indicating either 

that growth of contact area between grains may not have been 

the key aging mechanism or that this effect was too subtle to 

be measured.

Discussion

Results of our ring-shear tests indicate that the Mount 

St. Helens gouge exhibits rate weakening at shear rates below 

about 10-4 m/s. At these rates, the magnitude of a – b is comparable to typical published values for fault gouges, which 

vary from –0.001 to –0.01 (for example, Blanpied and others, 

1987; Kilgore and others, 1993). Above 10-4 m/s and at the 

lowest normal stresses of these experiments, a – b is positive, 

indicating that the gouge is rate strengthening. The apparent 

transition from negative to positive values of a – b occurs at a 

shear rate similar to the same transition documented for other 

rock and gouge shear zones (for example, Shimamoto, 1986; 

Blanpied and others, 1987; Kilgore and others, 1993). 

The transition shear rate in our experiments is only moderately larger than the rate of extrusion estimated for much 

of the 2004–2005 eruption (3×10-5 m/s to 7×10-5 m/s), but 

two observations indicate that gouge shear rates at Mount St. 

Helens may be well within the rate-weakening regime. Kilgore 
420 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006 STEADY-STATE FRICTION COEFFICIENT

SHEARING RATE, IN METERS PER SECOND

SHEARING RATE, IN METERS PER SECOND

10-6 10-5 10-4 10-3 10-2

0.4

0.5

0.6

0.7

5 kPa

23 kPa

10-6 10-5 10-4

0.41

0.42

0.43

0.44

0.45

0.46

0.47

0.48

86 kPa

141 kPa

159 kPa

195 kPa

σn

σn

STEADY-STATE FRICTION COEFFICIENT

A

B

μss = (τss/σn) μss = (τss/σn)

Figure 5. Steady-state friction coefficient as a function of 

shear rate. Each data point represents results of one experiment. 

Steady-state friction coefficient is shear stress supported by the 

gouge divided by corrected normal stress, both averaged over the 

part of experiment when time-averaged steady state had been 

attained (see fig. 4). Data are grouped by magnitude of applied 

normal stress. Error bars represent ±1 standard deviation of 

observed variability in friction coefficient during the period of the 

time-averaged steady state. A, High applied normal stresses. B, 

Low applied normal stresses.

Normal stress 

(kPa)

Shear-rate range 

(m/s) a−b b

Fast

5 5 ×10-4 – 5×10-3 0.0150

23 5 ×10-4 – 5×10-3 0.0105

Slow

5 1.5 ×10-6 – 2.5×10-5 –0.0250

23 1.5 ×10-6 – 2.5×10-5 0.00099

86 1.5 ×10-6 – 2.5×10-5 –0.0022

141 1.5 ×10-6 – 2.5×10-5 –0.0013

159 1.5 ×10-6 – 2.5×10-5 –0.0024 0.0079

195 1.5 ×10-6 – 2.5×10-5 –0.0103

Table 1. Rate-dependence parameters.

[Value a–b describes sign and magnitude of rate dependence of gouge friction. 

The b value alone describes magnitude of gouge strengthening as a function 

of hold time.]

and others (1993) noted that the transition shear rate increases 

with increasing normal stress. Because normal stresses in our 

experiments were probably at least one order of magnitude 

smaller than those on the margins of the magma plug at depths 

of 100–500 m at Mount St. Helens (2–10 MPa), rate-weakening behavior at Mount St. Helens would likely be exhibited 

at higher shear rates than those of the experiments. Secondly, 

Marone and Kilgore (1993) showed that the characteristic slip 

distance Dc is more appropriately viewed as a characteristic 

shear strain (shear displacement divided by the shear-zone 

thickness) when considering shear within a gouge layer. In 

ring-shear experiments on the Mount St. Helens gouge in 

which we placed vertical columns of beads in the gouge to 

study the shear-strain distribution, more than 90 percent of the 

strain was distributed over a thickness of 10 mm. At Mount St. 

Helens, the shear-zone thickness is likely larger, given the ~1 

m total thickness of the gouge layer and presence of particles 

too large to include in the ring-shear experiments. Thus, shearstrain rates in the ring-shear experiments were likely larger at 

a particular shear rate than those at Mount St. Helens, indicating that the gouge there would likely rate-weaken at higher 

shear rates than it did in the experiments. Thus, one of the conditions required for stick-slip ascent of the dacite plug—rate 

weakening of the gouge—appears to be met.

We now consider the second condition—that the stiffness 

of the loading system must be smaller than the critical stiffness 

defined by the evolution behavior of the gouge. Assessment 

of this condition requires independent estimates of both the in 

situ critical stiffness kcrit of the gouge and the stiffness ksys of 

the magma that is driving extrusion.

The critical stiffness of the gouge was defined by Ruina 

(1983) as the post-peak reduction in shear stress divided by the 

characteristic slip distance over which this reduction occurs:

c

n

crit

)(

D

ba k −−  ′ = , for (a – b) < 0. (6)

If a lithostatic stress state is assumed at the nucleation depth of 

the drumbeat earthquakes (~500 m) and pore-water pressure is 

neglected, such that the total normal stress equals the effective 

normal stress, then  ́

n ~10 MPa. A reasonable value of a–b 

indicated by our experimental results is –0.003. The characteristic slip distance Dc

 can be determined from the transition 

from steady slip to persistent stick-slip behavior that occurred 

in our experiments with high normal stresses. At this transition the experimental normal stress was 195 kPa (table 1), and 

kcrit equaled the stiffness of the test device. This stiffness is 

approximately 1.7×104 kPa/m (see appendix 1). (Here, again 

following Ruina (1983), we employ a stiffness having dimensions of stress per unit length (that is, mass/(time2

×length2))). 

Substituting this value for kcrit, a–b = –0.003 and the experimental value of  ́

n

 (195 kPa) in equation 6 yields Dc = 34 μm, 

which is of the same order as typical laboratory values for 

gouge (for example, Marone and Kilgore, 1993). Assuming 

that Dc = 34 μm, a–b = –0.003, and  ́

n =10 MPa are appropri-
20. Frictional Properties of the Mount St. Helens Gouge 421

0.42

0.40

0.44

0.46

0.48

0.50

0 2,000 4,000 6,000 8,000 10,000 12,000 14,000 16,000 18,000 20,000

-0.1

-0.08

-0.06

-0.04

-0.02

0

TIME, IN SECONDS

SPECIMEN THICKNESS CHANGE, IN MILLIMETERS

Friction coefficient

FRICTION COEFFICIENT (

μ = τ/σn)

Specimen thickness change

Figure 6. Part of one record of stick slip from experiments conducted at the two lowest shear 

rates (1.5×10-6 and 3.0×10-6 m/s) at 195 kPa normal stress. No direct measurement was made of 

displacement of the rotating base during these tests, so stick-slip displacement was inferred 

from shear-stress and specimen thickness records.

ate for the in situ gouge at Mount St. Helens, then equation 6 

yields an in situ critical gouge stiffness of ~106 kPa/m.

To estimate the stiffness of the loading system, we 

assume that magma underlying the plug drives extrusion (fig. 

8), and that the magma contains about 12 percent exsolved 

gas by volume (T. Gerlach, oral commun., 2005) and has a 

compressibility  = 10-4 per kPa (Iverson, this volume, chap. 

21). The pertinent stiffness k is the magmatic “spring force” 

per unit area of plug-margin slip surface divided by the 

“spring displacement” associated with magma compressionrelaxation cycles (compare with Ruina, 1983). To estimate this 

k, we approximate the magma body as a right circular column 

with height H = 8,000 m and cross sectional area A=1/4(�d 2), 

where d is the mean column diameter, and we approximate 

the area of the slip surface on the margins of the plug as �dh, 

where h is the height of the plug in contact with the conduit 

walls. The spring force associated with magma compression is 

(A/)(/H), where  is the characteristic displacement associated with magma compression or relaxation, and /H is the 

accompanying longitudinal strain of the magma column. Thus, 

the spring force per unit area of slip surface is d/(4hH ), and 

dividing this expression by the displacement distance  yields

 k = d/(4hH). (7)

If h is 500 m, and d is 50 m (these values are probably of the 

correct order of magnitude for the geometry at Mount St. 

Helens), then inserting these values and the  and H values 

noted above into equation 7 yields k ~10-1 kPa/m, seven 

orders of magnitude smaller than the critical gouge stiffness 

kcrit ~106 kPa/m. Therefore, even if the estimated values of 

magma compressibility or the geometrical parameters d, h, 

and H are greatly in error, the gouge is likely much stiffer 

than the magma column that drives upward movement of the 

plug. Thus, in addition to rate weakening of the gouge, the 

second requirement for stick-slip—that the critical stiffness 

of the gouge exceeds the stiffness of some part of the loading 

system—is likely satisfied.

The finding that the magma-plug system can exhibit 

stick-slip behavior at shallow depths contrasts with the common observation that fault slip is aseismic within ~2 km of 

the surface (Scholz, 2002). The tendency for aseismic tectonic 

fault slip at shallow depths is commonly attributed either to 

rate strengthening in unconsolidated gouge or to the dependence of critical stiffness on normal stress, which is low near 

Earth’s surface (Marone and Scholz, 1988; Scholz, 2002). Our 

results indicate that the Mount St. Helens gouge is rate weakening and the loading-system stiffness is exceptionally small 

owing to the presence of a large body of highly compressible 

magma. Therefore, stick-slip behavior should be possible at 

much smaller depths at Mount St. Helens than would be predicted for tectonically driven crustal faults.

100 101 102 103 104 105 106

1.00

1.02

1.04

1.06

1.08

1.10

HOLD TIME, IN SECONDS

NORMALIZED FRICTION COEFFICIENT

μpeak / 

μss

μpeak / μss = 1+0.0079 ln(t h)

Figure 7. Normalized peak friction coefficient, μpeak, 

as a function of logarithm of hold time. Standard error 

of b-value (slope of the regression) is 0.0004. All tests 

were at a normal stress of 159 kPa and shear rate of 

2.5 ×10-5 m/s.
422 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

h

d

H

Plug

Zone of gouge

formation and

drumbeat

nucleation

Gouge

α

Magma

Figure 8. Schematic illustration of 

hypothesized magma-plug system at Mount 

St. Helens. Cool, rigid dacite plug is driven 

upward through upper conduit by ascending 

compressible magma. Gouge forms by 

frictional wear at interface between extruding 

plug and conduit. Because the gouge is stiffer 

than effective stiffness of the magma that is 

driving plug uplift, uplift occurs in stick-slip 

cycles, producing drumbeat earthquakes. 

Physical and geometrical variables (letters 

and symbols on figure) are discussed in text.

Conclusions

Under normal stresses less than about 200 kPa, gouge 

sampled from the surface of the newly extruded Mount St. 

Helens dome in February 2005 exhibits frictional behavior 

that is described well by the rate- and state-dependent friction model. In our ring-shear device the gouge weakens with 

increasing steady shear rates up to about 10-4 m/s and strengthens at larger shear rates. This transition from weakening to 

strengthening occurs at an experimental shear rate only moderately larger than the observed time-averaged rate of extrusion at Mount St. Helens, but in situ the transition probably 

occurs at higher shear rates because in situ normal stresses 

in the gouge are probably significantly larger than those that 

could be applied experimentally. Also, the in situ stiffness 

of the gouge, as estimated using the conditions under which 

stick-slip occurred in the experiments, is roughly seven orders 

of magnitude greater than the stiffness of the most compliant component of the system driving uplift of the plug: the 

underlying magma. The rate weakening and high stiffness of 

the gouge relative to the magma are requirements for stick-slip 

behavior when rate-and-state dependent friction is present, so 

stick-slip behavior provides a viable hypothesis for the source 

of the observed shallow drumbeat earthquakes at Mount St. 

Helens. In addition, experiments showed that the peak strength 

of the gouge increased logarithmically with hold time, indicating that deformation history of the gouge could influence the 

character of stick-slip behavior and associated seismicity.

Acknowledgments

We thank John Pallister for collection of gouge samples 

from Mount St. Helens. Technical reviews by Nick Beeler and 

John Power and editorial suggestions by Jim Hendley significantly improved this paper.

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424 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

66

68

70

72

74

76

78

0.70 0.75 0.80 0.85 0.90 0.95 1.00

EXPECTED SHEAR DISPLACEMENT,

IN MILLIMETERS

SHEAR STRESS,

IN KILOPASCALS

0

0.2

0.4

0.6

0.8

MEASURED SHEAR

DISPLACEMENT, IN

MILLIMETERS

Shear stress

Measured displacement

Figure 9. Shear stress and measured shear displacement during 

initial minutes of a ring-shear test conducted at a normal stress of 159 

kPa. Data are plotted as a function of expected shear displacement 

calculated from rotation rate of motor that drives rotating base (fig. 

3). Data begin when motor was actively driving rotation, but no 

displacement was measured at perimeter of the rotating base. During 

subsequent abrupt slip, there was a marked reduction in measured 

shear stress, well below steady-state strength of the gouge. This 

stress reduction and associated abrupt displacement are interpreted 

to reflect elastic compliance of the device.

Appendix 1. Stiffness of the Ring-Shear Device

The stiffness of the ring-shear device can be estimated from stress-displacement transients that were common when shearing was initiated in experiments conducted at the highest normal stresses. Particularly after extended hold periods, the start of 

shear was characterized by sudden stress reductions that accompanied abrupt slip (fig. 9). Abrupt slip displacement was measured with a horizontally oriented LVDT that pressed on an armature extending from the rotating base of the device. This abrupt 

displacement of the rotating base occurred despite the constant speed of the motor driving the rotation of the base. These events 

reflect the release of elastic strain energy stored in the drive mechanism of the device, and therefore the ratio of abrupt stress 

reduction to abrupt slip describes the device stiffness, as applicable to the rate-and-state model of Ruina (1983)—1.7×104 kPa/m.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 21

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Dynamics of Seismogenic Volcanic Extrusion Resisted by 

a Solid Surface Plug, Mount St. Helens, 2004–2005

By Richard M. Iverson1

Abstract

The 2004–5 eruption of Mount St. Helens exhibited 

sustained, near-equilibrium behavior characterized by nearly 

steady extrusion of a solid dacite plug and nearly periodic 

occurrence of shallow earthquakes. Diverse data support the 

hypothesis that these earthquakes resulted from stick-slip 

motion along the margins of the plug as it was forced incrementally upward by ascending, solidifying, gas-poor magma. 

I formalize this hypothesis with a mathematical model derived 

by assuming that magma enters the base of the eruption 

conduit at a steady rate, invoking conservation of mass and 

momentum of the magma and plug, and postulating simple 

constitutive equations that describe magma and conduit compressibilities and friction along the plug margins. Reduction 

of the model equations reveals a strong mathematical analogy 

between the dynamics of the magma-plug system and those of 

a variably damped oscillator. Oscillations in extrusion velocity 

result from the interaction of plug inertia, a variable upward 

force due to magma pressure, and a downward force due to 

the plug weight. Damping of oscillations depends mostly 

on plug-boundary friction, and oscillations grow unstably if 

friction exhibits rate weakening similar to that observed in 

experiments. When growth of oscillations causes the extrusion 

rate to reach zero, however, gravity causes friction to reverse 

direction, and this reversal instigates a transition from unstable 

oscillations to self-regulating stick-slip cycles. The transition 

occurs irrespective of the details of rate-weakening behavior, 

and repetitive stick-slip cycles are, therefore, robust features of 

the system’s dynamics. The presence of a highly compressible 

elastic driving element (that is, magma containing bubbles) 

appears crucial for enabling seismogenic slip events to occur 

repeatedly at the shallow earthquake focal depths (<1 km) 

observed during the 2004–5 eruption. Computations show that 

fluctuations in magma pressure accompanying such slip events 

are <3 kPa, indicating that deviations from mechanical equilibrium are slight and that coseismic force drops are <108

 N. 

These results imply that the system’s self-regulating behavior 

is not susceptible to dramatic change—provided that the rate 

of magma ascent remains similar to the rate of magma accretion at the base of the plug, that plug surface erosion more or 

less compensates for mass gain due to basal accretion, and that 

magma and rock properties do not change significantly. Even 

if disequilibrium initial conditions are imposed, the dynamics 

of the magma-plug system are strongly attracted to self-regulating stick-slip cycles, although this self-regulating behavior 

can be bypassed on the way to runaway behavior if the initial 

state is too far from equilibrium.

Introduction

The dome-building eruption of Mount St. Helens that 

began in October 2004 was remarkable in several respects. 

This paper describes formulation, analysis, and predictions of 

a mechanistic model that links three key aspects of the eruption. The first and perhaps most striking of these aspects was 

extrusion of solid rock that emerged from the crater floor as a 

sequence of spines with conspicuous fault gouge and striations 

on their freshly exposed bounding surfaces. Second, extrusion 

proceeded at a nearly constant long-term rate (~1–2 m3

/s) that 

was sustained from December 2004 through at least December 2005. Third, extrusion was accompanied by more than a 

million small earthquakes that occurred almost periodically 

at hypocentral locations <1 km beneath the extruding dome. 

These eruption characteristics motivate the hypothesis that 

extrusion was driven by a nearly constant influx of magma at 

depth and resisted by a plug of solidified magma that slipped 

incrementally and seismogenically against the wall rock 

forming the upper parts of the magma conduit. I refer to this 

hypothesis as SPASM, an acronym for Seismogenic Plug of 

Ascending, Solidifying Magma.
426 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

This paper formalizes the SPASM hypothesis mathematically and tests whether it is consistent quantitatively with 

behavior observed during the 2004–5 eruption of Mount St. 

Helens. The mechanistic framework of the SPASM model is 

simple, reflecting my belief that a relatively simple physical 

process (that is, one involving few special conditions and contingencies) is most likely responsible for producing persistent, 

repetitive, natural events—such as nearly periodic earthquakes. Although the SPASM model aims chiefly to explain 

the origin of these earthquakes, it can also provide insight into 

conditions under which the eruption style might significantly 

change, and the potential for such change has large implications for assessment of volcano hazards.

Below, following a brief overview of key features of the 

2004–5 eruption of Mount St. Helens, I describe the conceptual basis of the SPASM model. I then present the mathematical formulation of the nonlinear SPASM equations and 

analytical results obtained from exact solutions of approximate 

(that is, linearized) versions of these equations. The analytical 

results demonstrate the plausibility of several broad classes of 

eruptive behavior, including both stable and unstable behavior. 

I then use numerical results obtained from approximate solutions of the exact, nonlinear model equations to clarify some 

consequences of instability. In particular, the numerical results 

show how stick-slip motion arises as a natural consequence of 

plug extrusion dynamics. In the final sections of the paper, I 

discuss implications of these findings for interpreting solidstate volcanic extrusion and accompanying seismicity.

Eruption Overview

Despite Mount St. Helens’ famous explosive eruption in 

1980, the dome-building activity that began in 2004 is consistent with the volcano’s recent geologic history (Mullineaux 

and Crandell, 1981). Over the past ~4,000 years, Mount St. 

Helens has extruded rock at a mean rate of about 0.2 m3/s 

while constructing a 26-km3 modern edifice (defined here as 

the volume above 1,220 m altitude) composed primarily of 

andesite and dacite lava flows and domes and their detritus. 

From 1980 to 1986 a dacite dome grew episodically in the crater formed during the 1980 eruption, and its volume ultimately 

reached 7.4×107

 m3 (Swanson and Holcomb, 1990). From 

1987 to 2004 Mount St. Helens did not erupt, although at least 

six phreatic explosions occurred from 1989 to 1991 (Mastin, 1994). Recurrent seismicity at depths of 2–8 km in the 

late 1980s and 1990s may have been associated with magma 

recharge but did not lead to eruptions (Moran, 1994).

Renewed eruptive activity began on October 1, 

2004, when a small explosion formed a vent through the 

~150-m-thick glacier that had grown in the southern part of 

Mount St. Helens’ crater since 1986 (Schilling and others, 

2004; Dzurisin and others, 2005; Walder and others, 2005). 

The explosion was preceded by about 7 days of increasingly 

intense seismicity at depths <1 km, but deeper seismicity 

(such as might be indicative of magma-chamber pressurization 

or depressurization) did not occur then and has not occurred 

subsequently (Moran and others, this volume, chap. 2). By 

October 11, explosions had largely ceased, seismic energy 

release had decreased to a rate about one-tenth that of the 

preceding two weeks, and extrusion of a solid dacite plug had 

begun (Dzurisin and others, 2005). By December 2004, extrusion rates had become nearly steady, and by December 15, 

2005, the volume of the resulting new lava dome was ~7.3×107

m3 (Schilling and others, this volume, chap. 8). This volume, 

added to that of the 1980–86 lava dome, implies that the mean 

extrusion rate at Mount St. Helens from 1980 to 2005 was 

about 0.2 m3/s, similar to the mean rate for the past 4,000 

years. Thus, the 2004–5 activity of Mount St. Helens was by 

no means unusual.

The remainder of this section focuses on the quasi-steady 

eruptive behavior observed at Mount St. Helens from December 

2004 through December 2005, because understanding this behavior is the goal of SPASM. Other papers in this volume provide 

detailed descriptions of the findings briefly summarized below.

Extrusion Rates and Vent Size

After extrusion commenced, it appeared to occur continuously through December 2005, and it produced a sequence of 

monolithic dacite spines, with some reaching heights >100 m. 

The fourth of these spines had such a strikingly smooth, 

symmetrical, elongate form that it resembled a breaching 

whale and was accordingly dubbed the “whaleback” (fig. 1). 

Each spine emerged over a period of several weeks to several months and eventually disintegrated as a consequence of 

fracturing and avalanching of rock from its exposed surfaces. 

By late 2005, the sequence of spines had formed a composite 

dome with the appearance of a multicrested pile of rubble, and 

distinguishing individual spines would in retrospect have been 

difficult without knowledge of their emplacement history.

Despite the rather complicated details of spine emplacement, photogrammetric analysis showed that the rate of 

extrusion remained remarkably constant from about December 

2004 to December 2005 (fig. 2) and typically ranged from 

about 1 to 2 m3

/s (Schilling and others, this volume, chap. 8). 

During the same period, the linear extrusion rate (that is, the 

speed of spine emergence from the ground) estimated from 

far-field, ground-based photography typically ranged from 

about 3×10-5 to 7×10-5 m/s (3–6 m/day) (Major and others, this 

volume, chap. 12), and these rates were largely consistent with 

high-precision data transmitted by short-lived GPS receivers placed intermittently on the extruding spines (LaHusen 

and others, this volume, chap. 16). Similar linear extrusion 

rates were inferred from short-duration, short-range photography, which detected centimeter-scale plug emergence over 

durations as brief as a few minutes (Dzurisin and others, this 

volume, chap. 14). Thus, all evidence indicates that extrusion 

rates were essentially constant over time scales longer than the 

duration of small earthquakes (about 10 s or less) but shorter 

than the multiyear duration of the eruption as a whole (>107

 s).
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 427

Figure 1. Oblique aerial photograph of extruding 

“whaleback” spine 4 at Mount St. Helens, viewed from the 

northwest, on February 22, 2005. Horizontal length of the 

smooth whaleback is about 380 m. Arrow shows vent from 

which the spine emerged. USGS photo by S.P. Schilling.

According to the SPASM hypothesis, small, abrupt 

pulses in plug extrusion were responsible for generating small 

earthquakes. These pulses would have typically entailed 

upward plug displacements of ~5 mm (a value obtained by 

multiplying the typical 5×10-5 m/s linear extrusion rate by 

the typical 100-s interval between earthquakes), but abrupt 

movements this small were not resolvable by displacement 

measurements. Moreover, abrupt slip occurring at earthquake 

hypocentral depths may have been muted at shallower depths 

owing to inelastic deformation of weak near-surface materials. As a consequence, no direct measurements of coseismic 

pulses of plug extrusion were made during the eruption. 

Therefore, the SPASM model addresses the dynamics of 

stick-slip cycles with properties that have been inferred but 

not directly measured.

Inference also plays a role in estimating the size of the 

vent where extruding spines breached the crater floor, because 

the presence of fragmented glacier ice and accumulated talus, 

as well as previously extruded rock, partly obscured the vent 

margins. Moreover, different parts of the vent became visible as successive spines emerged and moved laterally as well 

as upward. From the standpoint of constraining the SPASM 

model, the most useful estimates of the effective cross-sectional area of the vent are obtained not from direct observations of spine geometry but from comparison of volumetric 

and linear extrusion rates measured over extended periods. 

Division of the typical volumetric rate of 1.5 m3

/s by the typical linear rate of 5×10-5 m/s yields an effective vent area of 

30,000 m2, which implies an effective vent diameter of ~200 

m if the vent geometry is assumed to be circular. A vent of this 

size (but not necessarily circular) is assumed for all calculations I present in this paper.

Gouge Properties

Where fresh surfaces of newly extruded spines were 

exposed, they were coated with granulated, striated dacite 

interpreted to be fault gouge. The gouge presumably formed as 

a consequence of mechanical wear during localized shearing 

along the margins of the dacite plug as it moved upward relative to the adjacent conduit walls (compare Tuffen and Dingwell, 2005). Observations and dredge sampling from hovering 

helicopters indicated that the gouge thickness was typically 

about 1 m (J.S. Pallister, U.S. Geological Survey, oral commun., 2005). Striations on the gouge surface were abundant and 

generally aligned with the direction of extrusion (fig. 3).

Frictional properties of the gouge were measured using a 

large-scale ring-shear apparatus, in which a remolded, annular 

specimen (~0.012 m3

) was deformed in simple shear at various 

imposed rates (Moore and others, this volume, chap. 20). The 

measurements demonstrated that the gouge typically exhibited 

peak strength at displacements <0.5 mm, steady-state strength 

after about 3 mm of displacement, and reduction of steadystate strength with increasing shear rate (for example, fig. 4). 

At rates ≥5×10-4 m/s (about 43 m/day), this rate-weakening 

Figure 2. Measured volume of extruded dome rock at Mount 

St. Helens as a function of time (after Schilling and others, 

this volume, chap. 8). The red reference line shows growth of 

dome volume occurring with a constant extrusion rate of 1.5 

m3

/s from December 2004 to December 2005.
428 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

B

Figure 3. Photographs of striated, gouge-coated surfaces of 

extruding spines at Mount St. Helens. Arrows show direction of 

spine motion. A, Crest of spine 4 viewed from east on February 

22, 2005. Hovering helicopter (circled) provides scale. B, Base of 

spine 5 viewed from north on July 28, 2005. Field of view is roughly 

30 m wide. USGS photos by S.P. Schilling.

behavior was supplanted by rate-strengthening behavior, but 

the relevance of this transition is questionable because the test 

equipment could shear the specimen at rates ≥5×10-4 m/s only 

under very low confining stresses, ≤23 kPa (Moore and others, 

this volume, chap. 20).

Overall, the ring-shear tests of Moore and others (this 

volume, chap. 20) showed that frictional behavior of the 

Mount St. Helens gouge was largely consistent with that 

expected from models of rate- and state-dependent friction 

(for example, Dieterich, 1979; Ruina, 1983; Marone, 1998). 

Such models posit that frictional strength varies in proportion to the logarithms of the imposed shear rate and hold time 

(that is, the time a specimen is held in a static state between 

successive shear events). In rate-weakening materials, the 

combined effect of shear rate and hold time causes a reduction 

of frictional strength as steady shear rates increase. In tests of 

the Mount St. Helens fault gouge under confining stresses of 

86–195 kPa, measured values of steady-state friction coefficients ranged from 0.42 to 0.47, and these values declined 

logarithmically as the imposed shear rate increased (Moore 

and others, this volume, chap. 20). Measured peak friction 

coefficients were 1–9 percent larger than steady-state friction 

coefficients, and peak values increased logarithmically with 

hold time. This hold-time effect is typical of rocks and densely 

packed granular materials (for example, Beeler and others, 

1994; Losert and others, 2000).

Testing by Moore and others (this volume, chap. 20) also 

revealed that the effective in-place shear stiffness of the gouge 

was probably orders of magnitude larger than the effective stiffness of the magma body that loaded the gouge as it pushed the 

extruding plug upward. This contrast in stiffness, along with 

rate-weakening steady-state friction, provides a sufficient condition for stick-slip behavior in materials exhibiting rate- and 

state-dependent friction (Rice and Ruina, 1983; Ruina, 1983).

Drumbeat Earthquakes

An extraordinary feature of the 2004–5 eruption of 

Mount St. Helens was persistence of small (Md ≤2), shallow earthquakes that recurred so regularly they were dubbed 

“drumbeats” (Moran and others, this volume, chap. 2). The 

period between successive drumbeats shifted slowly with time 

but was commonly ~100 s (for example, fig. 5) and nearly 

always in the range 30–300 s. Seismograms showed that 

Figure 4. Example of data collected during ring-shear tests 

of gouge friction by Moore and others (this volume, chap. 20). 

Measured shear stress as a function of displacement is shown 

for two steady-state shear rates. Inferred friction coefficients 

(right-hand axis) represent shear stress divided by a nominally 

constant normal stress of 159 kPa, but normal stress fluctuated 

slightly (<1 percent) during the tests, such that friction 

coefficients are approximate (after Moore and others, this 

volume, chap. 20).
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 429

drumbeat waveforms generally had impulsive, high-frequency 

onsets and low-frequency codas, similar to waveforms of 

other hybrid volcanic earthquakes (for example, Lahr and 

others, 1994; Neuberg, 2000). Over time scales of hours to 

weeks, drumbeats typically had consistent sizes and did not 

display a Gutenberg-Richter magnitude-frequency distribution typical of tectonic earthquakes (Moran and others, this 

volume, chap. 2). Precise location of drumbeat hypocenters 

was hindered by the geologic and topographic complexity of 

the Mount St. Helens crater and the low density of crater seismometers, but within resolution limits (~100 m), all drumbeats originated at depths <1 km directly around or beneath 

the growing dome. Accompanying the drumbeats at irregular 

intervals were smaller and larger earthquakes (as large as Md

3.4) with differing seismic signatures, but these earthquakes 

had little lasting effect on the drumbeats.

The recurrence and character of drumbeat earthquakes 

implies the existence of a nondestructive seismic source, as 

Figure 5. Example of 24-hour seismogram illustrating nearly periodic occurrence of “drumbeat” earthquakes at Mount 

St. Helens. Graph shows seismicity recorded at station YEL, located 1.5 km north of the 2004–5 vent. Time begins at 21:00 

UTC on December 1, 2005, and scrolls from left to right and then top to bottom. Earthquake magnitudes were roughly 

0.5–1 during this interval. Data courtesy of Pacific Northwest Seismic Network.

has been inferred for other repetitive volcanic earthquakes 

(Lahr and others, 1994; Goto, 1999; Neuberg, 2000; Neuberg 

and others, 2006). The most outstanding attribute of the drumbeat earthquakes at Mount St. Helens, however, was their periodicity. This periodicity, together with the presence of striated 

fault gouge bounding the extruding plug, is the key motivation 

for the SPASM hypothesis.

Although the SPASM model aims to link plug extrusion and earthquake generation, it does not address resulting 

seismic radiation. Radiation of seismic waves could result 

from rapid propagation of rupture that spreads along the 

fault surface after nucleating in a strong “keystone” patch of 

fault gouge (see Scholz, 2002). Alternatively, the force drop 

accompanying frictional slip that occurs uniformly along 

the plug margins might be so abrupt as to radiate seismic 

energy (see Marone and Richardson, 2006), and this force 

drop is calculated by the SPASM model. In relating SPASM 

mechanics to seismic radiation, however, it must be borne 

1800

0 1 2 3 4 5 6 7 8

TIME (minutes)

TIME (hours)

9 10 11 12 13 14 15

1200

0600

0000

YEL EHZ UW
430 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

in mind that only a small fraction (<10 percent) of the work 

done during fault slip typically results in seismic radiation 

(McGarr, 1999). Generally, more work is done in overcoming friction, and the mechanics of the total work cycle is the 

primary focus of SPASM.

Magma Solidification and Compressibility

Although genesis of the magma that erupted in 2004–5 

remains uncertain, petrologic data indicate that solidification 

occurred at depths <1 km. This inference derives from the 

fact that the composition of glass in the newly erupted dacite 

plots between the 0.1 and 50 MPa cotectics of the modified 

quartz-albite-orthoclase phase diagram for Mount St. Helens 

dacites (Blundy and Cashman, 2001) and from the presence 

of tridymite, which constrains the late stages of solidification 

to pressures of 10–20 MPa (equivalent to estimated lithostatic 

pressures at depths ~0.5–1 km) (Pallister and others, 2005, and 

this volume, chap. 30). These depths are consistent with the 

maximum hypocentral depths inferred for drumbeat earthquakes (Moran and others, this volume, chap. 2), and they 

imply that solid-state extrusion and earthquake generation 

were collocated.

The quantity and composition of volcanic gases emitted during the 2004–5 eruption demonstrate that the magma 

was gas poor in comparison to the 1980 Mount St. Helens 

magma but that the 2004–5 magma nonetheless contained 

sufficient exsolved gas to greatly influence its compressibility. At 8 km depth the gas volume fraction was probably 

<2 percent, but calculations using methods of Newman and 

Lowenstern (2002) indicate that the gas volume fraction 

grew during magma ascent and reached about 50 percent at 

~1 km depth, where solidification began (Gerlach and others, this volume, chap. 26). The same calculations indicate 

that at depths between 8 and 1 km, the gas volume fraction 

averaged ~12 percent. Allowing for inevitable gas separation 

from rock during extrusion, these results are consistent with 

observed vesicle volume fractions of 11–34 percent in samples 

of the 2004–5 dacite (Gerlach and others, this volume, chap. 

26). An exsolved gas volume fraction of 12 percent implies a 

magma compressibility ~10−7 Pa−1, according to the model of 

Mastin and Ghiorso (2000). Although this value is, of course, 

inexact, its mechanical significance is clear: the magma was 

almost certainly much more compressible than solid rock, 

which generally has compressibilities <10−10 Pa−1 (Hatheway 

and Kiersch, 1989). As a consequence, magma compression 

almost certainly dominated elastic strain as pressure within the 

magma-conduit-plug system increased.

Geodetic Inferences About Magma Influx

Measured displacements of the volcano flanks and adjacent areas preceding and during the 2004–5 eruption imply 

that the volume of magma evacuated from depths <10 km was 

considerably less than the volume of extruded rock (Lisowski 

and others, this volume, chap. 15). No evidence of systematic 

preeruption surface displacement was found by global positioning system (GPS) surveys in 2000 and 2003 of a 40-station 

network centered on the volcano, nor by continuous operation of GPS station JRO1, located 9 km north of the eruption 

vent. Seismicity that heralded the eruption in late September 

2004 was accompanied by only centimeter-scale downward 

and southward (that is, inward) surface displacements at JRO1 

(Lisowski and others, this volume, chap. 15). The displacement 

pattern measured at all stations corresponds well with that 

predicted by an elastic half-space model that assumes pressure 

decrease within a vertically oriented, prolate spheroidal cavity 

with a mean depth of 8 km and volume loss ~2×107

 m3 during the period from October 1, 2004, to November 25, 2005 

(Lisowski and others, this volume, chap. 15). This apparent 

volume loss is less than one-third the volume of rock extruded 

during the same period, and little of the apparent volume loss 

occurred after the onset of nearly steady extrusion in December 2004, implying that magma recharge from a deep (>10 

km) source accompanied this phase of the eruption.

Conceptual Basis of Mathematical 

Model

The basic mechanical elements of the SPASM model 

are shown schematically in figure 6, and table 1 defines all 

mathematical symbols used in development and analysis of 

the model. The model assumes that magma flows into the 

base of a feeder conduit at a steady volumetric rate Q. The 

conduit is assumed to originate about 8 km beneath the Mount 

St. Helens crater, a depth inferred from hypocentral locations 

of pre-2004 earthquakes apparently associated with magma 

movement (Moran, 1994; Moran and others, this volume, 

chap. 2). Ascent of magma at the top of the conduit is resisted 

by force exerted by a near-surface plug of solidified magma, 

owing to its weight mg and boundary friction F. Seismic and 

petrologic data collected during the 2004–5 eruption imply 

that the plug extends to a depth <1 km, and as a baseline value 

I assume that it extends to a depth ~500 m. The plug mass m

can change with time as a consequence of basal accretion of 

congealing magma at mass rate B , where  is the magma 

bulk density, and as a consequence of surface erosion by spalling and avalanching at mass rate rE , where r is the bulk 

density of the plug rock. The conduit volume can change with 

time as a result of motion of the base of the plug and changes 

in magma pressure that cause elastic deflection of the conduit walls. The magma pressure and density can change in 

response to the changing balance between the steady magma 

influx and changing conduit volume. The resulting mathematical model represents the simultaneous evolution of the upward 

plug velocity u, magma pressure against the base of the plug 

p, and conduit volume V, which are influenced by concurrent 

evolution of m and .
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 431

Implicit in the SPASM model is a “top-down” perspective 

of eruption dynamics. The model focuses on observed surface 

and near-surface phenomena associated with eruptive behavior 

but does not consider phenomena associated with unobserved 

changes that might occur in a deep magma reservoir. Instead, 

in the SPASM model, variations in extrusion rate are postulated to arise naturally as a consequence of the dynamics of 

the solid plug responding to steady forcing. Unsteady forcing 

due to unsteady magma influx would complicate behavior 

exhibited by the SPASM model but would not change its fundamental character.

The SPASM model is one-dimensional and does not 

explicitly consider the effects of conduit and plug geometry. 

This simplification poses both an advantage and disadvantage. 

The advantage derives from the fact that predictions of the 

SPASM model are independent of geometrical effects and are, 

in a general sense, applicable to any geometry. The disadvantage is that SPASM yields no insight concerning the effects of 

geometrical complications such as variations in the shape of 

the magma conduit, vent, or growing lava dome.

Many previous eruption models have used a one-dimensional approach similar to the one used here, and some models 

have invoked stick-slip motion as a phenomenon responsible 

for cyclical eruptive behavior (for example, Denlinger and 

Hoblitt, 1999; Voight and others, 1999; Ozerov and others, 

2003). The SPASM model, however, is the first to demonstrate 

how stick-slip behavior arises as a natural consequence of 

system dynamics. Indeed, a key feature of the SPASM model 

is that forces need not be balanced; therefore, the model can 

exhibit dynamical behavior not possible in eruption models 

that assume balanced forces (for example, Mastin and others, 

this volume, chap. 22).

Mathematical Formulation

The most fundamental equations used to derive the 

SPASM model express conservation of mass and linear 

momentum of the solid plug and conduit fluid. These conservation laws are supplemented by constitutive equations defining magma compressibility, conduit wall-rock compliance, and 

the frictional force acting where the plug contacts the conduit 

walls. In this section the conservation and constitutive equations are presented and reduced to a set of three simultaneous 

differential equations that describe behavior of the magmaconduit-plug system as a whole.

Conservation of Linear Momentum of Solid Plug

Changes in the upward momentum of the solid plug are 

described by Newton’s second law of motion, expressed as

du m u p A m g F

dt dt

+ = − −

dm , (1)

where m is the plug mass, u is the vertical (upward) plug 

velocity, g is the magnitude of gravitational acceleration, and p

is the magma pressure against the base of the plug, which has 

area A in horizontal projection. Upward motion of the plug is 

driven by the basal magma-pressure force pA and resisted by 

the plug weight mg and boundary friction force F. Implicit in 

equation 1 is the understanding that F would change sign (that 

is, friction would reverse its direction of action) if u were to 

change sign. A detailed specification of F is provided in the 

section on “Constitutive Equations,” below.

Conservation of Mass of Solid Plug

Mass change of the solid plug depends on the rate of 

mass accretion at the base of the plug, B , and the rate of 

mass loss at the surface of plug due to erosion, rE , where B

is the volumetric rate of magma conversion to solid rock and 

E is the volumetric erosion rate of the plug surface. Magma 

solidification may involve a change in mass density from 

to r . Summing these effects yields the mass-conservation 

equation

r

dm B E 

dt = − = , (2)

where  is a convenient shorthand for   rE. For the sake 

of simplicity,  is assumed constant, although this assumption 

is readily relaxed in numerical analyses if warranted. With 

constant , equation 2 yields the explicit solution

Figure 6. Schematic diagram illustrating the 

conceptual framework of the SPASM model.
432 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Symbol Dimensions Definition

A L2 Horizontal cross-sectional area of base of plug

B L3/T Volumetric rate of magma solidification at base of plug

C M/T Dimensional measure of rate dependence of plug friction force

c none Dimensionless measure of rate dependence of plug friction force

c1

, c2 none Arbitrary constants of integration

D none Dimensionless parameter summarizing effects of frictional damping

E L3

/T Volumetric rate of erosion of surface of plug

F ML/T2 Friction force on margins of plug

F0 ML/T2 Value of F at static limiting equilibrium

F(u0

) ML/T2 Value of F at the steady equilibrium extrusion rate u0

F* none F/F0

G none Dimensionless parameter summarizing effects of gravity

g L/T2 Gravitational acceleration

Hcon L Vertical height of magma-filled conduit

Hplug L Vertical height of extruding solid plug

K none Dimensionless parameter summarizing effects of plug mass change

m M Plug mass

m0 M Static or steady equilibrium value of m

p M/LT2 Magma pressure against base of plug

pb M/LT2 Magma pressure at base of conduit

pi M/LT2 Value of p at onset of a stick event

p0 M/LT2 Value of p at static limiting equilibrium

pref M/LT2 Reference value of p used in magma compression equation

Q L3

/T Volumetric rate of magma influx at base of conduit

R none Dimensionless parameter equal to 1− ρ/ρr

R0 none Value of R when ρ = ρ0

R* none R/R0

S none Dimensionless parameter summarizing rate dependent plug friction

ˆ

S none Dimensionless parameter summarizing static plug friction

T T Oscillation period

t T Time

t

0 T Natural time scale of oscillations, defined in equation 25

t

* none t/t

0

u L/T Vertical extrusion velocity

u0 L/T Value of u at steady equilibrium

uref L/T Reference value of u used in friction equation

u ́ L/T u−u0

u* none u ́/u0

ui

* none Initial value of u*

V L3 Volume of magma-filled conduit

V0 L3 Static or steady equilibrium value of V

V ́ L3 V−V0

V* none V ́/V0

Table 1. Definitions of mathematical symbols.
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 433

Symbol Dimensions Definition

Vi L3 Value of V at onset of stick event

W L3

/T Parameter defined in equation 53

X none Dimensionless parameter defined in equation 44

Y none Dimensionless parameter defined in equation 44

Z none Dimensionless parameter defined in equation 44

α1 LT2

/M Elastic bulk compressibility of magma

α2 LT2

/M Elastic compliance of walls of magma-filled conduit

η M/LT Magma viscosity

κ M/T Rate of change of plug mass

λ none Parameter that relates plug weight to boundary normal stress

μ0 none Static friction coefficient

ν none Dimensionless parameter defined in equation 34

ρ M/L3 Bulk density of magma

ρ0 M/L3 Value of ρ at static limiting equilibrium

ρr M/L3 Bulk density of plug rock

ξ none Rescaled time variable defined in equation 34

Table 1. Definitions of mathematical symbols.—Continued

m = t + m0 (3)

where m0 is the initial value of m.

Conservation of Linear Momentum of Conduit 

Fluid

Newton’s second law for upward motion of magma in the 

conduit takes a simple form if variations of magma properties 

and velocity with position are neglected:

2

b 8

con

dQ d p - p Q A g Q 

dt dt H A

  + = − −    

.

 

(4)

Here Q is the vertical (upward) volumetric flux of magma, 

b p is the magma pressure at the base of the conduit, and η

is the magma viscosity. Equation 4 is the fluid-mechanical 

equivalent of equation 1 and is also equivalent to the 

Navier-Stokes equation for one-dimensional laminar flow, 

integrated over the conduit cross-sectional area A and 

height Hcon . According to equation 4, upward motion of 

magma in the conduit is driven by the vertical pressure 

gradient ( )/ b con p - p H and is resisted by the magma unit 

weight g and viscous drag, represented by the last term 

in the equation. The form of this drag term is inferred from 

an elementary analysis of Poiseuille flow in a cylindrical 

conduit, although alternative drag terms (appropriate for 

other conduit geometries or magma rheologies) could be 

used without difficulty.

A simplified momentum equation is obtained by assuming that the magma flux Q is independent of time and making 

the substitution dQ dt / = 0 in equation 4. Rearrangement of 

the resulting equation yields an explicit expression for Q,

. 8

b

con

p p

 g H Q A d dt A



 

− −

=

+

 (5)

Equation 5 shows that maintenance of constant Q in the presence of changing Hcon , p, and  (all of which can occur in 

the context of the SPASM model) can imply that compensating changes occur in b p and/or η. The model assumes that 

such compensating changes may indeed occur, but it does not 

evaluate such changes explicitly. A complete evaluation could 

be accomplished by using equation 4 together with a massconservation equation (see below) to model the dynamics of 

transient magma flow in the conduit (for example, Melnik and 

Sparks, 2002). However, such a model also requires specification of a basal boundary condition (for example, magmachamber pressure) to drive magma inflow. Any such specification involves assumptions that are arbitrary, and the SPASM 

model minimizes use of arbitrary assumptions by specifying a 

constant basal magma influx Q.

Conservation of Mass of Conduit Fluid

The mass of the fluid magma in the conduit is  V , where 

V is the conduit volume. Changes in  V depend not only on 

changes in  and V but also on the influx of fluid mass at the 
434 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

base of the conduit Q and the loss of fluid mass at the top of 

the conduit B, which results from solidification at the base of 

the solid plug. These phenomena are summarized by the fluid 

mass-conservation equation

( ). dV d V Q B

dt dt

  + =−  (6)

Both Q and B are treated as constants.

Constitutive Equations

Although the conservation equations for Q and m reduce 

to the explicit forms shown above, the remaining two conservation equations (1 and 6) contain four dependent variables, 

u, p, V, and  and an as-yet-unspecified friction force F. Thus, 

three constitutive equations must be specified to attain mathematical closure.

Magma Compressibility

The first constitutive equation defines the compressibility 

of the fluid magma 1 as

1

1 , d

dp

 

 = (7)

and integration of equation 7 yields 0 1 exp[ ( )] ref   = p p − , 

where 0 is the magma bulk density at a reference pressure

ref p . Combination of equation 7 with the chain rule 

ddt (ddp)(dp dt) yields an equation that relates magma 

density change to pressure change:

1 . d dp dt dt

 =  (8)

Below, this equation is used to replace density derivatives with 

pressure derivatives where advantageous.

Conduit Compliance

A second constitutive equation defines the bulk elastic 

compliance of the conduit walls as

2

0

1 , dV

V dp



  =     (9)

where the subscript 0 denotes conduit volume change under a 

condition of zero plug velocity (u=0) and zero plug accretion 

(B=0). The utility of equation 9 is increased by embedding 

the equation in a definition of the total rate of conduit volume 

change that occurs when u and B are nonzero,

0

.

r

dV dV A u B 

dt dt





  =− +     

 (10)

Here again, the subscript 0 denotes the rate of volume change 

that would exist if u=0 and B=0, whereas the terms Au and 

(r) B describe conduit volume change due to upward plug 

motion and basal plug accretion, respectively. The factor 

r accounts for the influence of density change from  to 

r during magma solidification at the volumetric rate B.

To obtain a “systemic” constitutive equation for total conduit volume change, equation 9 is embedded in equation 10 by 

using the chain rule 0 0 [ / [ / ] /( / ) dV dp = dV dt dp dt ] , yielding

2 .

r

dV dp A u B V

dt dt

 



=− + (11)

The volume change described by equation 11 includes both an 

irreversible component and a reversible (elastic) component.

Plug Boundary Friction

The final constitutive equation defines the friction force 

F that acts where the plug contacts the conduit walls. This friction results from shearing of gouge, discussed briefly above in 

the section on “Gouge Properties” and in detail by Moore and 

others (this volume, chap. 20). Because friction might potentially exhibit diverse behaviors, and because this diversity has 

significant ramifications for extrusion dynamics, I represent F

with a functional form that is consistent with the key findings 

of Moore and others (this volume, chap. 20) but that compromises between precision, generality, and simplicity:

1

0 sgn( ) 1 sinh .

ref

u F u mg c 

u

 −   =  +     

 (12)

Here sgn(u) denotes the sign of u and stipulates that the 

frictional force always opposes motion; 0 is a static friction coefficient applicable when u=0; c is a parameter that 

describes the sign and magnitude of frictional rate dependence; and ref u is a reference velocity that specifies the extent 

of nonlinearity of rate dependence (fig. 7). In the simplest 

case, with c=0, equation 12 specifies that the friction force 

has a constant value 0 0 F F u mg = = sgn( )  . If c ≠ 0 and 

/ 1 ref uu << , then equation 12 implies that friction depends 

almost linearly on slip rate ( 0 sgn( ) [1 / ] ref F u mg ≈  c u u + ), 

whereas for / 1 ref u u >> , equation 12 implies that rate dependence of friction is essentially logarithmic 

( 0 sgn( ) [1 ln 2 / ] ref F u mg ≈  c u u + ) (see Abramowitz and 

Stegun, 1964, p. 87). This logarithmic dependence mimics 

behavior observed in the steady sliding experiments of Moore 

and others (this volume, chap. 20). The fact that nearly logarithmic behavior as well as other styles of frictional behavior 

may be represented by equation 12 is a significant advantage 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 435

in analytical studies. Another advantage is that the equation 

implies that the maximum friction force is finite at u=0 (that 

is, F F mg = =± 0 0 ), whereas purely logarithmic friction 

rules imply that friction is infinite at u=0.

As specified by equation 12, the friction force F is 

proportional to the effective normal force on the sides of the 

plug, mg , where λ is a numerical factor that scales this 

normal force to the plug weight, mg. This definition implies 

that the maximum plausible value of λ is about 2 (assuming 

μ0~0.5, as shown by the data of Moore and others, this volume, 

chap. 20), in which case sidewall friction suffices to support 

the entire plug weight. Realistic values of λ are likely to be 

considerably smaller than 2 and are dependent on the height of 

the plug in contact with the conduit walls and on the state of 

effective stress governing the normal traction on the plug margins. Although the effective stress state is unknown at Mount 

St. Helens, estimates of λ are constrained by the balance of 

forces implied by the right-hand side of equation 1 for the case 

of static limiting equilibrium: F F mg pA mg == = − 0 0 . 

Algebraic rearrangement of this balance shows that λ must satisfy 0  = (1/ )[( / ) 1]  pA mg − . Therefore, because the magma 

pressure p is unlikely to deviate much from lithostatic pressure 

(for if it did, it would cause hydraulic fracturing or conduit 

collapse), λ is largely determined by the plug geometry, which 

determines the plug mass m and basal area A, as well as the 

depth where p operates.

Although friction described by equation 12 represents 

both the peak-strength effect and shear-rate effect observed in 

the experiments by Moore and others (this volume, chap. 20), 

it includes no provision for the hold-time (or “state evolution”) 

effect also observed in those experiments. In the context of 

equation 12, inclusion of such evolution would entail making

0 a time- or state-dependent quantity. This complication 

would introduce additional constitutive parameters, but it 

Figure 7. Graph of equation 12, illustrating nonlinear variation 

of relative friction force as a function of normalized slip rate 

/ ref u u and rate-dependence parameter c. Parameter F0

 is 

friction force at static, limiting equilibrium, F mg 0 = 0 0 .

would add little to understanding the mechanism of regularly 

occurring drumbeat earthquakes at Mount St. Helens. Therefore, I have chosen to exclude state-evolution effects from the 

SPASM model.

Friction represented by equation 12 does include a 

simple yet fundamental type of state dependence, however. 

The factor sgn(u) in equation 12 stipulates that the static 

(zero-velocity) friction force can jump from a positive value 

mg0 to a negative value ≥ −mg0 if the extrusion velocity u changes from a positive value to a negative value. If u

subsequently becomes positive again, then the friction force 

again becomes positive. Such jumps ensure that friction 

opposes motion, and they have great implications for the 

dynamical behavior of the extruding plug.

Reduced Governing Equations

The equations described above can be reduced to a compact system of three equations governing simultaneous evolution of the dependent variables u, V, and p. In this system the 

magma density is eliminated as a dependent variable by using 

equation 8 to replace ddt in equation 6 with dpdt and then 

dividing all terms in the resulting equation by , yielding

1 . dV dp V Q B dt dt

+ =−  (13)

Equations 11 and 13 are then combined and rearranged algebraically to obtain explicit equations for ddt and dVdt. These 

two equations accompany the equation of motion obtained by 

combining equations 2 and 12 with equation 1, thereby forming 

a system of three first-order differential equations,

0

1 (/ ), refdu g pA u F u u dt m t





=− +  − −    +

 (14)

[ ] 1 2

1/ , and dp V Au RB Q dt  − = +−

+ (15)

[ ] 1

1 2

, dV A u RB Q Q B dt



 

= + − +−

+ 

(16)

where R is a nearly constant coefficient defined by

0

1 1 1 exp[ ( )]. ref

r r

R p p   

  =− =− − (17)

The system of differential equations (14–16) contains four 

types of nonlinearities: one involving the quotient A u /V in 

equation 15, one involving the nonlinear dependence of R on 

p shown in equation 17, one involving the potential for jumps 

in F described in the section above on “Plug Boundary Friction,” and one involving the dependence of F on 1 sinh / ref u u −

shown in equation 12.
436 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The Forced, Damped Oscillator Equation

 The physical implications of the governing equations 

become clearer when equations 14 and 15 are combined 

to form a single, second-order equation. Differentiating 

equation 14 with respect to t, employing the chain rule 

dF dt dF du du dt / ( / )( / ) = , and substituting equation 15 into 

the resulting equation yields

( ) 2 2

0 2

1 2 1 2

( ) . [] []

d u dF du A A Q RB m t 2 u g dt du dt + V + V

  

   

    − + + + +     =− +    

 

2

1 2 1 2

( ) . [] []F du A A Q RB 

u g 

u dt + V + V



   

   − 

+ 

   =− + 

  

 (18)

This equation has a form like that of equations governing 

behavior of forced, damped oscillators (Kreyszig, 1979, p. 82 

ff.), and it implies that the extrusion velocity u has a natural 

tendency to oscillate about equilibrium. This tendency arises 

from the interplay of plug inertia, an upward “spring force” 

due to compression of the magma and conduit, and the 

downward force due to gravity. Oscillations of u are complicated by the presence of a time-dependent mass term

m0 +t, variable damping implicit in dF du / , and variable 

forcing due to the presence of R and V in the last term of 

equation 18, as well as by coupling of equation 18 to equations 16 and 17. Of course, equation 18 is strictly valid only 

insofar as F is a continuously differentiable function of u, 

and this restriction must be borne in mind when interpreting 

solutions of equation 18.

Linearized Dynamics: Analytical 

Results

Analytical solutions of equation 18 are important for 

guiding interpretation and numerical solution of the full equation set, 14–16, despite the fact that analytical solutions can 

be readily obtained only for various special cases. Analytical 

results also aid identification of diverse eruptive styles that 

are represented by the governing equations but are difficult to 

identify through numerical solutions alone.

Static, Steady-State, and Pseudosteady-State 

Solutions

The most basic special cases involve assumptions about 

the extrusion velocity u. A very simple special case with great 

physical importance assumes a condition of static limiting 

equilibrium, in which u = 0 and plug boundary friction just 

suffices to resist the upward force due to magma pressure. 

This case also assumes that the plug mass is constant (that is, 

 = 0 , m m= 0 ) and that the magma influx rate Q and basal 

accretion rate B are zero. These assumptions lead to the static 

equilibrium solution,

0 0

0 0 0 , mg F u pp V V

A

+ = == =

 

(19)

which satisfies equations 14–16 as well as 18. Here 0 p is the 

limiting equilibrium magma pressure at the base of the plug, F0

is the friction force at static limiting equilibrium, F mg 0 = 0 0 , 

and V0 is an arbitrary but constant conduit volume.

An equally important special case assumes that the plug 

mass is constant ( = 0 , m m= 0 ) but that volumetric rate of 

basal plug accretion and volumetric rate of magma influx 

are finite and equal (B=Q). These conditions lead to an exact 

steady-state solution satisfying equations 14–16 as well as 18,

 

0 00

0 0

( ) , Q R B m g F u u u p V V

A A

− + == = =

 

(20)

in which 0 u is the steady-state upward plug velocity,

0 F(u ) is the steady-state friction force, 1

0 00 0 ( ) [1 sinh ( / )] F u m g c u u  ref

− = + , and R0

 is a constant 

value of R that applies when  = 0 (that is, 0 0 1( / ) R = −  r ). 

The equation group 20 represents dynamic equilibrium of the 

steadily ascending magma-plug system, whereas transient 

states represent departures from this equilibrium. Below, 

analyses of these departures show that steady states can be 

stable in some circumstances and unstable in others.

In addition to the exact steady state described by equation 

group 20, pseudosteady states can exist in which the extrusion velocity u remains constant but the magma pressure p

increases or decreases as the plug mass evolves according to

m m0 = +t . A pseudosteady-state solution that satisfies equations 14–16 and 18 is

0 0 1 2

2

1

( ) ( ) () (). Q RB g V mg Fu gt u A Q B u pV AA AA g  

 

− + + + −

=− = + =

 

 0 0 1 2

2

1

( ) ( ) () (). Q RB g V mg Fu gt u A Q B u pV AA AA g  

 

− + + + −

=− = + = (21)

Existence of this pseudosteady state requires that B≠Q and 

that R changes negligibly as p evolves. Despite these restrictions, the state described by equation group 21 has physical significance because it implies that essentially steady 

extrusion may occur in the presence of evolving plug mass 

and magma pressure, and it has mathematical significance 

because it provides a check on numerical results reported 

later in this paper. Additional pseudosteady states may, of 

course, exist if values of other parameters (for example, F0

, 

A, 1, 2) evolve, but such evolution is not addressed explicitly in this paper.
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 437

Linear Approximation of Transient States

Analysis of transient states is facilitated by linearization. As 

a first step, the dependent variables u and V in equation 18 are 

decomposed into sums of the steady-state values, defined in equation group 20, and transient deviations from steady state, denoted 

by primes: 

 0 0 u u u t V V V t =+ =+ ′( ) ( ). ′ (22)

Substitution of equation group 22 into 18 and elimination of 

terms that sum to zero yields a simplified, but still nonlinear, 

version of equation 18 that describes the behavior of transient 

deviations. One linearization of this equation results from the 

assumption that magma density changes are small in comparison to the steady-state density, which enables R to be approximated by its steady-state value 0 0 1( / ) R = −  r . A second 

linearization involves the assumption that deviations in the 

magma-conduit volume are small in comparison to its steadystate volume (that is, 0 V′/V <<1). Then neglect of small terms 

involving 0 V′/V decouples equation 18 from 16 and removes 

the associated nonlinearity. These two linearizations generally 

have little effect on model predictions because they involve 

physical effects that are typically very subtle.

The most significant linearization involves approximation 

of dF du / in equation 18. If u remains positive, an exact, nonlinear expression for dF du / follows from the definition of F in 

equation 12 and can be written as (see Abramowitz and Stegun, 

1964, p. 88)

 

1/2 2

0 0 1 .

 

ref ref

dF m g c u du u u



−

    =  +         

 (23a)

Substituting 0 u u u = + ′ in equation 23a and simplifying the 

result algebraically yields an approximation of 23a that is valid 

if velocity deviations from steady state are sufficiently small 

that 0 u /u ′ <<1 and 1 ref u /u ′ << :

 

1/2 2

00 0 1 .

ref ref

dF m g c u C

du u u



−

    ≈  +    =        

 (23b)

Here C is a constant with the same sign as c but with dimensions of mass/time. If friction is rate independent, then C=0, 

whereas C>0 indicates rate-strengthening friction and C<0 

indicates rate-weakening friction.

The linearized form of equation 18 results from making 

the substitutions shown in equations 21 and 23b and assuming

R = R0 , 0 V /V ′ <<1, and 0 u /u ′ <<1, which yields

 

 

Except for the plug mass-growth factor 0 t / m , all coefficients 

in equation 24 are constant, a property that facilitates analysis.

Natural Period of Oscillations

The form of equation 24 implies that if  = 0 (that is, 

the plug mass is constant), then u′ will oscillate freely with 

constant period 0 T t = 2 , where 0t is the natural time scale 

implied by the reciprocal of the coefficient that precedes u′ in 

equation 24: 1/2

00 1 2

0

[ ( )] . m V

t

A

 + = (25)

This result is demonstrated more formally in the section on 

“Solution for Undamped Free Oscillations” below, but T and 

the time scale 0t are introduced here as a basis for normalization of equation 24 and a first comparison of T and the typical 

interval between repetitive drumbeat earthquakes at Mount St. 

Helens. To facilitate this comparison, equation 25 is recast in a 

special form that is appropriate if the magma conduit and plug 

are approximated as right cylinders (not necessarily circular) with cross-sectional areas A and heights Hcon and H plug , 

respectively. In this case the oscillation period 0 T = 2t can be 

expressed as 1/2

1 2 2 [( ) ] . T =    + r con plug H H (26)

A graph of equation 26 for the values Hcon = 8 km and r =

2,000 kg/m3 is depicted in figure 8. The graph shows how the 

free oscillation period T varies as a function of the plug height 

H plug and lumped compressibility 1

+2. For reasonable values of H plug and 1

+2, the predicted T has values that range 

from about 10 s to several minutes. The similarity of these values to the observed recurrence period of drumbeat earthquakes 

during the 2004–5 eruption of Mount St. Helens helps support 

the hypothesis that the drumbeats were associated with oscillations of the extrusion rate u.

Normalized Oscillator Equation and 

Dimensionless Parameters

The quantity 0t defined in equation 25 provides the 

appropriate time scale for normalization of equation 24, and 

this normalization leads to identification of the dimensionless 

parameters that control the linearized dynamics of the magmaplug system. Substitution of the normalized time *

0 t = t / t and 

normalized velocity deviation *

0 u = u′/ u into equation 24 

reduces the equation to

 

(1+Κt∗) d 2 u∗

dt* 2 + (2Κ+GS) du∗

dt∗ + u∗ = −KG ,

 

(27)

in which K, S, and G are dimensionless parameters defined as

0

0

, t K

m

 = (28)

0

0

, and Cu S

m g = (29)

0

0

. gt G

u = (30)

(24)

1+  −−+ − +

t

m0

d 2u'

dt 2 dt  2 + C du'

m0

 A2

m0

V0

└

┌

│

└

┌

│ α1+α2

u'=− − m0

gj .
438 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 8. Graph of equation 26, illustrating predictions (from 

linearized theory) of variation of oscillation period T as a function 

of lumped compressibility 1 +2 and plug height H plug .

It is also useful to define a dimensionless damping factor D, 

which is half the coefficient in the second term of equation 27:

 

 D =1

2 (2 + GS )= t

0

m0

 + C

2 . (31)

If K=0, this damping factor plays a role like that of damping 

factors in textbook examples of linear oscillators, and D=1 

constitutes critical damping (for example, Kreyszig, 1979, p. 

82 ff.). This interpretation changes only subtly for cases with 

K0, as shown below.

A complete assessment of the magnitudes of the dimensionless parameters defined above is provided in the section 

on “Normalized Nonlinear Equations and Control Parameters” 

below, but for present purposes it suffices to note that typical 

magnitudes of K, G, and S imply that a satisfactory approximation of equation 27 commonly results from neglecting the 

plug growth term * Kt as well as the effect of K on D. An even 

simpler but still relevant approximation is obtained by setting 

both K and S equal to 0 in equation 27. On the other hand, 

because solutions of equation 27 for nonzero values of these 

parameters imply diverse eruptive behaviors, which can differ 

qualitatively as well as quantitatively, I analyze the full spectrum of these behaviors before considering numerical solutions of the nonlinear system of equations 14–16.

Solution for Undamped Free Oscillations

For the case in which K=0 and S=0 (implying constant 

plug mass and constant plug-margin friction), equation 27 

reduces to an elementary second–order equation describing 

undamped, free oscillations of the plug velocity. In this case 

the simplest nontrivial solution of equation 27 is

* * * cos , i u u t = (32)

which obeys the initial conditions * *

i u = u and * * du dt / = 0 . 

This solution demonstrates that the natural period of the plug’s 

velocity oscillations is 0 T = 2t , as inferred in the section on 

“Natural Period of Free Oscillations” above. According to 

equation 32, sinusoidal oscillations with period T and amplitude *

i u continue forever if an initial disturbance with magnitude *

i u causes them to begin, provided that K=S=0.

Solutions for Damped Free Oscillations—

Stability of Quasi-Steady Extrusion

Another class of solutions of equation 27 exists if 

K=0 but S0. Physically, these conditions imply that the plug 

mass is constant S0 and that plug-margin friction depends 

linearly on the extrusion rate. In such cases the behavior of 

solutions depends on the value of D(= S/2) relative to the 

transition value D=0 and critical values D=± 1, as summarized in table 2. In all cases the solutions imply stable eruptive 

behavior if D>0 (that is, if rate-strengthening friction exists), 

because * u decays with time essentially like Dt* e− . On the 

other hand, if D<0 (that is, if rate-weakening friction exists), 

solutions containing the factor Dt* e− indicate unstable growth 

of * u . The instability can be manifested as either runaway 

acceleration or oscillations that grow with time, as shown by 

the solutions listed in table 2.

The solutions listed in table 2 can also be used to infer 

whether the steady eruptive state represented by equation 

group 20 is physically accessible from an initial static state 

that exists before the onset of a volcanic eruption. For all 

cases with D>0, transient deviations in velocity ( * u ) decay 

toward zero as time proceeds, implying that any transient 

state will eventually give way to a steady state. On the other 

hand, for cases in which D<0, dynamic steady states are 

inaccessible from an initial static state, because any transient 

motion grows without bound as time proceeds. Of course, 

this simple picture can change if K0 or if friction is nonlinear, as described in the section on “Nonlinear Dynamics: 

Numerical Results” below.

Solution for Damped, Forced Oscillations—

Effects of Plug Growth on Extrusion Stability

The most complicated solutions of equation 27 apply 

to cases in which the plug mass changes with time such that 

 ≠ 0 and none of the parameters K, S, and G is zero. For such 

cases, solutions can be obtained by a multistep process that 

involves a simultaneous change of dependent and independent 

variables to transform equation 27 into Bessel’s differential 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 439

D General solution of equation 27 Type of behavior

0 * * *

1 2 u c t c t = sin cos + Undamped oscillations with

constant amplitude

 >0 but <1 * * 2 * 2*

1 2 [ cos ( ) sin ( )] Dt u c t 1 D + c t 1 D e− =− − Underdamped oscillations;

unstable growth if D<0

1 * * *

1 2 ( ) Dt u c c t e− = + Critically damped oscillation;

unstable growth if D<0

>1 2* 2* * ( 1) ( 1)

1 2

DD t DD t u c e + c e −+ − −− − = Overdamped, no oscillations;

unstable growth if D<0

Table 2. General solutions of equation 27 for cases with K=0 and D=S/2. All solutions contain arbitrary constants, 

c1 and c2

, which are constrained by specifying initial conditions.

equation (appendix 1). Transformation of the well-known 

Bessel-equation solution back to the original variables * u and *

t then yields the general solution

[ ] *

1 2 ( ) ( ), 4

u KG c J c Y



 

 =− + +   (33)

in which

*

2

2 1 2 . K D Kt K K

  − + = = (34)

Here 1 c and 2 c are arbitrary constants, and J and Y are 

Bessel functions of the first and second kind, of order ν (see 

Abramowitz and Stegun, 1964). Equation 33 describes diverse 

behaviors, partly analogous to those summarized in table 2. 

However, oscillations described by equation 33 have periods 

as well as amplitudes that evolve with time.

The relation between the solution for K0 (equation 33) 

and the solutions for K=0 (table 2) has important physical 

implications. The relation is clarified by considering asymptotic approximations of J andY that are valid for  >>1. 

(The condition  >>1 implies *

t K K >> −1/ , a criterion 

that is typically satisfied as time proceeds if K>0. If K<0, the 

criterion will not be satisfied as time proceeds, and equation 

33 then implies that a singularity develops in which * u → ∞ . 

Physically, this singularity represents a “catastrophe” in which 

the plug mass reaches zero, liquid magma reaches the surface, 

and the governing equations no longer apply.) For  >>1 the 

Bessel function approximations are (Abramowitz and Stegun, 

1964, p. 364)

2 J ( ) cos ( / 2 / 4), and        ≈ −− (35)

2 Y ( ) sin ( / 2 / 4).        ≈ −− (36)

Substituting these approximations into equation 33 and collecting terms containing powers of  shows that, over time, * u decays or grows according to

1/4 *

*

2

1 .

D/K

Kt

u

K

−

 +  ∝    

 (37)

The exponent 1/4 / − D K in equation 37 plays a physical role 

analogous to that of −D in solutions with K=0 (table 2). The 

physical meaning of equation 37 becomes clearer if equations 

29 and 31 are used to express the exponent in terms of physical parameters, yielding 1 3

* 2 2 *

2

1 .

C 

Kt

u

K



  −  +   +   

∝     (38)

The exponent in equation 38 shows that oscillations in u* 

will decay or grow depending on whether the criterion

C

─

>− 3

─

2 (39)

is satisfied. Several classes of behavior are implied by this 

result and are summarized in table 3.

The chief physical implication of the results summarized 

in table 3 is that extrusion stability can depend on a tradeoff 

involving the rate dependence of the frictional resisting force 

and the rate of change of plug mass. Stable eruptive behavior 

can occur in the presence of rate-weakening friction (C<0) 

if the plug mass increases at a sufficient rate. This behavior 

contrasts with that of a system with constant plug mass, which 

necessarily exhibits unstable behavior if C<0 and, therefore, 

D< 0 (table 2). Conversely, unstable behavior can occur in the 

presence of rate-strengthening friction (C >0) if the plug mass 

decreases at a sufficient rate.

Although further inferences can be drawn from the analytical 

results summarized by equations 38 and 39, for present purposes it 
440 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

κ C Behavior of equation 33

κ = −(2/3)C C = −(3/2)κ Undamped, constant-amplitude oscillations

κ > 0 C > 0 Oscillations necessarily decay toward a steady state

κ > 0 C < 0 Oscillations grow if κ < −(2/3)C and decay if κ > −(2/3)C

κ < 0 C > 0 Oscillations grow if κ < −(2/3)C and decay if κ > −(2/3)C

κ < 0 C < 0 Behavior becomes singular

Table 3. Summary of behavior of equation 33, which is general solution of equation 27 for cases with 

changing plug mass (K≠0, κ≠0).

[Behavior for both rate-strengthening friction (C > 0) and rate-weakening friction (C < 0) is summarized.]

is more useful next to consider numerical solutions of the nonlinear equation set 14–16. Nonlinearities produce important effects 

that are not revealed by analytical results that strictly apply only 

when transient disturbances are small (that is, 0 u′/ u <<1).

Nonlinear Dynamics: Numerical 

Results

Results of the linear theory point to several questions 

to be addressed through numerical solution of the nonlinear 

system of equations 14–16. Do the bounds of stable versus 

unstable eruptive behavior and character of oscillatory eruptive 

behavior change when nonlinearities exist? More specifically, can nonlinearities result in stick-slip instabilities like 

those inferred to produce drumbeat earthquakes at Mount St. 

Helens? What controls the magnitude and frequency of stickslip events? Can such events occur repeatedly (that is, forever) 

until some attribute of the system changes? Can evolution of 

stick-slip periods and amplitudes yield inferences about evolution of system properties? Do complications such as plug mass 

change have significant effects? How sensitive are the system’s dynamics to disequilibrium initial conditions? (Although 

the linear theory assumes that departures from equilibrium are 

always small, nature imposes no such constraint, and volcanic 

eruptions necessarily begin in disequilibrated states.)

Normalized Nonlinear Equations and Control 

Parameters

Guidance for investigating the behavior of numerical 

solutions comes from identification of dimensionless control parameters and their likely magnitudes. Behavior of the 

nonlinear equations is governed partly by the same dimensionless parameters that govern linearized behavior, but additional 

parameters also play a role. The additional control parameters 

are identified by normalizing the nonlinear system of firstorder equations 14–16 through use of dimensionless variables 

defined as * * * *

0 00 0 u uu p p p V VV t tt = = == /, / , /, / . 

Substituting these variables into equations 14–16 yields the 

normalized system

*

** *

* *

1 ˆ ˆ (1 ) , 1du G G S p Ku GSF dt Kt

=− + + − −     + 

(40)

* *

* *

*

1/ 1 ( 1) , and ˆ (1 )

dp V u ZR

dt G S

− = −+ −     + (41)

*

* *

* 1 ( 1) . dV X u ZR Y

dt

= −+ − +     (42)

Here, G and K are dimensionless parameters defined exactly 

as in the linear model (that is, in equations 28 and 30). The 

dimensionless parameter ˆ

S is related to S defined in the linear 

model (that is, in equation 29), but its definition is somewhat simpler because it involves F0 rather than the derivative

C dF du = / : 0

0

0

ˆ . F S

m g = =  (43)

The dimensionless parameters X, Y, and Z have no analog in 

the linear model and are defined as

 

1 00 0 0

0 0 0

( ) . mu t Q B RB X Y Z

At V u A

 − = = =

 

(44)

The variables R* and F* in equations 40–42 are normalized versions of the density variable R defined in equation 17 and the 

friction force F defined in equation 12:

*

0 *

0 0

ˆ 1 ( 1)exp (1 )( 1) , R XG S p R R

R R

+− + −     = =

 

(45)

 ( ) * * 1 *

0

0

sgn( ) 1 sinh [ ( / )] . ref

F F u c uu u

F

− == + (46)

In the nonlinear version of the SPASM model, equations 45 

and 46 must be satisfied simultaneously with equations 40–42.
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 441

Plausible ranges of the values of the dimensionless parameters in equations 40–46 for the 2004–5 eruption of Mount St. 

Helens are listed in table 4. The tabulated values imply that 

some of the terms in equations 40–46 will have little influence on numerical results. For example, relevant values of the 

denominator G(1+Ŝ) on the right-hand side of equation 41 are 

undoubtedly very much greater than 1, implying that * p will 

remain close to its static equilibrium value * p =1. This inference, together with the inference that X<<1 (table 4), implies 

that the argument of the exponential function in equation 45 will 

remain close to zero, despite the large probable value of G(1+Ŝ). 

Therefore, R*=1 is typically a good approximation of equation 45, and this approximation implies that the term Z(R*

−1)

in equations 41 and 42 will have only subtle effects. Similarly, 

the term Y in equation 42 will have subtle effects because Y<<1 

(table 4). As a basis for prioritizing investigations, then, it is 

reasonable to assume initially that Y=Z=0.

With these simplifications in mind, equation 40 can be differentiated and combined with equation 41 to obtain a secondorder equation, which can also be obtained through normalization of the original oscillator equation 18. For the case in which 

R*

=1 is a good approximation, the complete system of normalized nonlinear governing equations thereby reduces to

( ) 2 *

* *

2 * *

1 1 1 (2 ) , and

*

* *

d u du Kt K GS u K G

dt dt V V

+ + + + =− (47)

 *

* ( 1). dV X u

dt = − (48) 

Note that S rather than ˆ

S appears in equation 47, and that the 

entire set of dimensionless parameters has collapsed to K, G, 

X, and S in equations 47 and 48. Moreover, the values of K and 

X are typically much smaller than 1 (table 4), and the effects of 

terms containing K and X are therefore apt to be modest. On this 

basis, computational investigations aimed at illuminating the 

physics represented by the SPASM model can focus principally 

on the effects of G and S, secondarily on the effects of K, and 

lastly on the effects of X, Y, and Z.

The strategy of focusing primarily on effects of G and S

reduces the need to explore a large, multidimensional parameter space numerically, and it parallels development of the linearized theory, in which D emerged as the key control parameter. Indeed, D defined in the linearized theory applies also to 

the nonlinear model, provided that D is viewed as a numerical 

index rather than a constant damping factor. For this purpose it 

is useful to define the index D as the value applicable when the 

slip rate equals the steady equilibrium rate (u = u0

):

 (49) 

j

kl 2

2

1

1 2 1

D 2K + GS

c

t t 0 0 0

u u

u

m0 0

u=u

0

( ) =

+ +

=

ref ref

This equation is simply an algebraically expanded version of 

the definition given in equation 31, and I employ this definition of D to index numerical results.

It is also noteworthy that for cases in which K=0 and 

0 / 1 ref u u >> , equation 49 reduces to the simplified form

, (50)

which is commonly a satisfactory approximation. In the 

sparest distillation of the nonlinear SPASM model, then, 

numerical results can be expected to depend primarily on D

as defined in equation 50, which in turn depends only on c, 

λ, and 0 if the equilibrium extrusion rate 0 u and oscillation 

time scale 0t are fixed.

Computational Method

Numerical solutions were obtained by using a standard 

fourth-order Runge-Kutta method described by Press and 

others (1986). To implement the Runge-Kutta algorithm, 

a double-precision FORTRAN program was written and 

executed on a personal computer with a 2.26-GHz processor. 

Constant time steps were used to generate all solutions and 

were typically 0.0001 to 0.01 s. Although some exploratory 

computations required hours of CPU time, no computations 

reported in this paper required more than several minutes of 

CPU time when using this constant-time-step approach.

The accuracy of numerical solutions was checked against 

exact analytical solutions for simple linear cases with constant 

values of D and K (table 2 and equation 33). For nonlinear 

cases, some aspects of numerical solutions were checked 

analytically by exploiting the fact that the governing equations 

14–16 yield nearly exact solutions for p and V for the special 

case in which u = 0. These solutions assume that R = R0, and 

they have the form

, and V = V Wt i + (51)

0

1 2

ln 1 , ( ) i

i

Q R B W

p p t

  W V

−  

=+ +   +  

 (52)

where

[ ] 1

0

1 2

W R B Q Q B. 

 

= − +−

+ 

(53)

Here V i and i p are the values of V and p at the beginning of 

any “stick” episode with u=0. The duration of stick episodes, 

Tstick , may be calculated by solving equation 52 for t while 

setting p equal to the static limiting equilibrium pressure p0 

necessary to trigger any slip episode,

0 12

0

( )( ) exp 1 . i i

stick

V p p W T 

W Q R B

  − +     =    −    −  

 (54)
442 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Parameter Units Value(s) Comments on value(s)

Specified physical parameters

A m2 30,000 Calculated using Q/u0 and values tabulated here

B m3

/s 0–10 Cannot differ greatly from Q

c none -0.01–0.01 Inferred from results of testing by Moore and others, (this volume, chap. 20)

C kg/s | C | < 5×1013 Calculated using equation 23 and values tabulated here

g m/s2 9.8 Typical value at Earth’s surface

m0 kg 5×109

–7×1010 Inferred from plug heights 100–1,000 m and A and ρr tabulated here

Q m3

/s 1–2 Inferred from photogrammetric measurements of dome growth

R0 none -0.5–0.5 Calculated from typical values of ρ0 and ρr

 tabulated here

t

0 s 0.4–150 Calculated from equation 25 and other values tabulated here

u0 m/s 2×10-5–7×10-5 Inferred from measured linear extrusion rate

uref m/s 7×10-8 –7×10-5 Smaller than u0 if friction rate-dependence is nonlinear

V0 m3 3×106

–3×108 Inferred from 8 km conduit height and conduit radii 10–100 m

α1 Pa-1 10-8 –10-6 Typical values for silicic magma with 1–50 vol percent bubble content

α2 Pa-1 ≤ 10-9 Typical values for fractured rock

κ kg/s | κ | < 4,000 Exceptions may occur during dome-collapse events

λ none 0.1–1 Inferred from plug geometry and plausible effective stress states

μ0 none 0.4–0.5 Inferred from results of testing by Moore and others, (this volume, chap. 20)

ρ0 kg/m3 1,200–2,400 Typical values for silicic magma with 1–50 vol percent bubble content

ρr kg/m3 1,600–2,400 Inferred from measurements on dome-rock specimens

Derived dimensionless parameters

D none -107–107 Influence of D can be very significant

G none 6×104–2×107 Influence of G is very significant

K none | K | ≤ 10-4 Influence of K is subtle except in event of abrupt dome collapse

S none -7–7 Influence of S is significant

ˆ

S none -0.5–0.5 Influence of ˆ

S is significant if multiplied by G

X none 4×10-4 –8×10-10 Influence of X is subtle except perhaps where multiplied by G

Y none | Y | ≤ 10-4 Influence of Y is subtle

Z none | Z | < 0.5 Influence of Z may be significant, contingent on value of R*

Table 4. Plausible values of SPASM model parameters applicable to the quasi-steady dome-building eruption of 

Mount St. Helens, 2004–2005.

[Derived dimensionless parameters determine model behavior and are formed from combinations of physical parameters.]

This equation is useful for checking computational results, and 

it also has a significant physical implication: the duration of 

stick episodes increases exponentially with 0 i p − p . From this 

result it may be inferred that plug displacements during slip 

events also increase exponentially with 0 i p − p , because the 

average extrusion velocity is fixed (as given by equation 20), 

and slip-event magnitude must therefore increase in proportion 

to stick duration. Although strict validity of equations 52 and 

54 rests on the assumption that R = R0 , computations that do 

not employ this assumption show that, nonetheless, equations 

52 and 54 generally provide good predictions.

Computational Results

Only a small number of computational solutions are 

presented here, but hundreds of additional solutions were 

computed and examined. Results chosen for presentation 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 443

Figure 9. Start-up behavior of solution computed for D 

= −0.01 with linear rate weakening. Initial condition is

u = Q / A , 0 p = p , V =V0. Oscillations of slip velocity (u), 

magma pressure (p) and conduit volume (V) are sinusoidal 

and match analytical predictions until u = 0 occurs and stickslip behavior begins.

highlight a range of physical effects that appear particularly 

important for understanding the origin of drumbeat earthquakes at Mount St. Helens. To a lesser degree, results were 

also chosen to illustrate the spectrum of behaviors possible 

within the framework of the SPASM model.

Behavior with D Close to Zero

The most basic nonlinear features of solutions 

appear even when departures from linear behavior are 

slight. In particular, stick-slip cycles develop in any case 

in which D<0, provided that a jump in F is imposed at 

u=0. Figure 9 illustrates this behavior for a case in which 

D = −0.01, B=Q, and F depends linearly on u (that is, 

0 0 sgn( ) [1 / ] F u m g c u u =  + ref ). The small value of D

results from use of an unusually small value of λ, λ=0.01; 

physically, this value implies that frictional resistance and the 

damping it produces are small. The computation also used the 

initial conditions u=Q/A, V =V0 , and 0 p = p , which imposed 

a slightly perturbed initial magma pressure (because the static 

equilibrium pressure 0 p slightly exceeds the steady equilibrium pressure) and used the parameter values K=0, Y=0, Z=0, 

X=5×10-6, T=10 s, 2,000 kg/m3, m0=3.6×1010 kg, 

μ0=0.5, c=−1.71×10-5, and uref=0.1(Q/A)=6.66710-6 m/s. The 

fact that the value of uref is significantly smaller than that of 

the typical extrusion rate (Q / A ) implies that effects of rate 

dependence in the friction rule are important.

Under the conditions described above, the computed 

extrusion behavior is initially identical to that predicted by the 

linear analytical theory, and u, p, and V each exhibit exponentially growing sinusoidal oscillations until u=0 occurs (at t ≈

27 s in fig. 9). At that time F momentarily changes sign and 

thereby halts motion of the plug as it starts to descend. This 

event heralds the end of exponential oscillation growth and the 

onset of repetitive stick-slip cycles.

In both the sinusoidal and stick-slip cycles shown in 

figure 9, magma pressure oscillates 1⁄4 cycle out of phase 

with slip velocity, and conduit volume oscillates 1⁄2 cycle out 

of phase with magma pressure. Although it may seem contradictory that conduit volume decreases as magma pressure 

increases, this out-of phase response results from conditions 

at the base of the extruding plug, where solidification and 

accretion occur continuously at the volumetric rate B, even 

as plug velocity diminishes to less than the steady-state value 

u = Q / A =6.667×10-5 m/s. Indeed, as a rough approximation, 

conduit volume declines whenever the plug extrusion rate is 

less than the basal accretion rate. The conduit volume also 

responds elastically to pressure changes, but this effect is typically overshadowed by volume changes associated with plug 

motion and basal accretion.

Key elements of the solution presented in figure 9 are 

recast in a phase-plane diagram in figure 10, which shows 

how pressure deviations from the static equilibrium pressure 

p0

 vary in concert with velocity deviations from the steady 

equilibrium velocity Q/A. As portrayed in the phase plane, 

the initial condition (u = Q/A, 0 p = p , marked I.C. in fig. 

10) is unstable. This instability leads to an outwardly diverging clockwise spiral representing simultaneous oscillations 

of u and p that grow with time. (Note that arrows in phaseplane diagrams throughout this paper point in the direction 

of advancing time.) When the spiral becomes large enough 

to encounter the condition u = 0, divergence ceases and the 

dynamics become locked in stick-slip limit cycles that repeat 

endlessly thereafter. If the initial condition is located elsewhere inside the stick-slip limit cycle of figure 10, behavior 

nonetheless diverges smoothly until locking in the same limitcycle state. If friction exhibits rate-strengthening rather than 

11,818.88

11,818.84

11,818.80

11,818.76

11,818.72

11,818.68

632,008

632,004

632,000

631,996
444 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

11,818.88

11,818.84

11,818.80

11,818.76

11,818.72

632,004

632,002

632,000

631,998

631,996

rate-weakening behavior (that is, D>0), then behavior converges smoothly to a steady fixed-point equilibrium instead. 

An analogous fixed-point equilibrium is exhibited by eruption 

models that assume forces are always balanced (for example, 

Mastin and others, this volume, chap. 22).

Modification of the behavior depicted in figures 9 and 10 

by nonlinearity in the rate-weakening friction rule is illustrated 

in figures 11 and 12. Initial conditions and parameter values 

used to generate figures 11 and 12 were the same as those 

used to generate figures 9 and 10, but the nonlinear friction 

rule 1

0 sgn( ) [1 sinh / ] ref F u mg c u u − = + was employed. 

Comparison of figure 11 with figure 9 demonstrates that the 

most conspicuous effects of the nonlinearity are to delay 

the onset of stick-slip behavior and shorten the duration of 

individual stick events. These effects are unsurprising, because 

the nonlinearity represented by the sinh-1 function increasingly 

suppresses rate weakening as the slip rate increases.

Comparison of the phase-plane diagrams shown in 

figures 10 and 12 demonstrates that the dynamical effects of 

linear and nonlinear rate weakening also differ in other ways. 

In both figures 10 and 12 the feature of greatest interest is 

the outer loop representing stick-slip limit cycles. In the limit 

cycles shown in both figures, the maximum positive magma 

pressure deviates by a factor of only 4×10-6 from the static 

equilibrium pressure (~107 Pa), and the maximum slip velocity 

is only slightly more than double the equilibrium slip velocity. The maximum velocity is somewhat larger in the case 

with linear rate weakening (fig. 10), however, because linear 

rate weakening enables a larger dynamic overshoot of slip 

Figure 10. Phase-plane representation of simultaneous 

evolution of normalized slip velocity and normalized magma 

pressure computed for D = −0.01 with linear rate weakening. 

Initial condition (I.C.) is: u = Q / A , 0 p = p , V =V0 . Arrows 

point in direction of advancing time.

in response to increasing magma pressure. (Here, “dynamic 

overshoot” means that inertia carries the moving plug upward 

past an equilibrium point in which forces are balanced.) This 

larger overshoot produces a commensurately larger decline in 

magma pressure in response to slip, and the magma pressure 

deviation at the onset of stick-slip limit cycles is about three 

times larger in the case with linear rate weakening than with 

nonlinear weakening (~6×10-6 versus 2×10-6). This difference 

in pressure deviation constitutes the single most important difference between the linear case (fig. 10) and the nonlinear case 

(fig. 12), because the duration of stick periods (Tstick ) increases 

Figure 11. Start-up behavior of oscillatory solutions 

computed for D = −0.01 with nonlinear rate weakening. Initial 

condition is u = Q / A , 0 p = p , V =V0 . Oscillations of slip 

velocity (u), magma pressure (p), and conduit volume (V) are 

sinusoidal until u = 0 occurs and stick-slip behavior begins.
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 445

exponentially with this pressure deviation, as indicated by 

equation 54. Therefore, nonlinearity in the friction rule has a 

significant effect on predictions of the periods and amplitudes 

of stick-slip oscillations, and the nonlinear rule was employed 

to generate all results presented subsequently in this paper.

Effects of the Damping Index D

Some of the most important findings of this study are 

summarized in figure 13, which illustrates computational 

results obtained by employing the nonlinear friction rule and 

various values of D. Computations that generated these results 

used the same parameter values used to generate figures 9–12, 

except that here = 0.2 was used, and varying values of c

were used to generate D values ranging from −0.2 to −4. (The 

value = 0.2 constitutes a “best-guess” value applicable to 

Mount St. Helens’ plug geometry and state of effective stress, 

and c values ranging from 1.7×10-5 to 3.4×10-4 were used to 

simulate subtle rate weakening similar to that observed experimentally by Moore and others, this volume, chap. 20).

Displacement time series shown in figure 13A were computed for a family of stick-slip cycles with various values of D, 

and in figure 13B the same results are depicted as limit cycles 

in the velocity-pressure phase plane. Initial conditions used to 

generate figure 13 assumed a static, limiting equilibrium state 

(u=0, 0 p = p , V =V0 ) rather than the state with 0 u = u used 

to generate figures 9–12. Therefore, no divergent oscillations 

precede the development of stick-slip limit cycles.

Figure 12. Phase-plane representation of simultaneous 

evolution of normalized slip velocity and normalized magma 

pressure computed for D = −0.01 with nonlinear rate 

weakening. Initial condition (I.C.) is u = Q / A , 0 p = p , V =V0 . 

Arrows point in direction of advancing time.

Figure 13. Stick-slip cycles computed for various values of parameter D with nonlinear rate weakening. All computations 

employed baseline parameter values Q=2 m3

/s, B=Q, R0 = 0, K=0, Y=0, Z=0, X=5×10-6, T=10 s, and employed m0=3.6×1010 kg, λ=0.2, 

μ0

=0.5, and uref=0.1(Q/A)=6.667×10-6 m/s to determine F. Varying values of c were used to obtain varying values of D. A, Time series 

representation of stick-slip displacements. B, Phase-plane representation of velocity-pressure limit cycles. Arrows point in 

direction of advancing time.
446 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The basic dynamics portrayed in figure 13B are simple. 

When basal magma influx produces pressure exceeding the 

static equilibrium value ( 0 p ), it triggers slip at a rate that may 

slightly or greatly surpass the steady equilibrium rate (Q/A), 

depending on the value of D. When a combination of plug 

inertia and diminishing magma pressure no longer suffices to 

overcome the effects of gravity and friction, slip terminates 

and stick begins. Magma pressure then rebuilds until it triggers 

another slip event. Figure 13A shows that periods of stickslip cycles with D= −0.2 differ little from the T=10 s period 

predicted by linear theory (that is, equation 25), but periods 

increase as D values range further from 0. For D values sufficiently far from 0, periods increase almost in direct proportion to the magnitude of D, and the amplitudes of slip events 

increase accordingly. Maximum slip speeds and associated 

pressure deviations during stick-slip cycles also increase in 

proportion to the magnitude of D (fig. 13B), bolstering the 

inference that D values encapsulate most of the important 

controls on system dynamics.

Stick-slip cycles computed with D= −2 closely resemble 

those thought to be responsible for generating drumbeat 

earthquakes at Mount St. Helens. With D = −2 , individual slip 

events entail about 5 mm of displacement in about 5 s and 

maximum slip rates of ~1.7 mm/s. Attendant fluctuations in 

magma pressure are <0.02 percent of p0 (fig. 13B), equivalent 

to only ~2.4 kPa or <0.2 m of static magma pressure head. 

This result implies that a remarkably delicate shift in the balance of forces distinguishes periods of slip from those with 

no slip. Multiplied by A = 30,000 m2, the ~2.4 kPa pressure 

change also serves as a proxy for the force drop responsible 

for generating seismicity (~7×107 N).

Details of Baseline Case with D=−2

Deeper exploration of the dynamics computed with 

D= −2 provides further insight to physical phenomena that 

may be responsible for drumbeat seismicity at Mount St. Helens. Figure 14 illustrates details of repetitive earthquake cycles 

computed with D=−2. The histories of slip velocity, magma 

pressure, and conduit volume shown in figure 14 illustrate 

abrupt decreases in magma pressure and increases in conduit 

volume during slip events and also illustrate gradual changes 

of these quantities between slip events. As noted above, 

magma solidification at the base of the plug causes the volume 

of the fluid-filled conduit to decline between slip events, 

despite the fact that magma pressure rises.

Perhaps the most intriguing result illustrated in figure 

14 involves the history of shear force along the plug margins. The shear force is large (~3.5×1010 N) because it must 

overcome the effects of both gravity and friction to move the 

massive plug upward, but the force drop during each slip event 

is comparatively small (~7×107 N) (fig. 14E). (This drop in 

shear force is closely related to the stress drop that occurs in 

conjunction with tectonic earthquakes, but force drop is, in 

fact, a more fundamental quantity. The force drop represents 

the product of the stress drop and the area of the slip surface—

a product that appears directly in earthquake energy budgets 

(for example, Scholz, 2002, p. 184). Shear force might be 

concentrated in a relatively small patch of gouge bounding the 

extruding plug at Mount St. Helens, or it might be distributed 

evenly within the gouge; from the standpoint of force drop, 

this distinction makes no difference.) Standard estimation 

methods indicate that such a 7×107

 N force drop, accompanied 

by 5 mm of slip (figs. 14A, E), implies about 2×105 J of seismic energy radiation (Scholz, 2002, p. 185), whereas the work 

done against friction during the slip events depicted in figure 

14 is the total shear force times displacement, which yields 

an estimate of 2×108 J. These results are consistent with prior 

findings that only a small fraction of the work done during 

fault slip produces seismic radiation (McGarr, 1999).

The temporal pattern of the drop in shear force accompanying slip events also has significant implications. As slip accelerates, the shear force declines smoothly owing to the effects of 

rate-weakening friction (fig. 14E). Similarly, as slip decelerates, 

the force smoothly rises as friction gradually increases. When slip 

stops, however, the shear force drops abruptly because it suddenly 

returns to a static equilibrium value imposed by the plug weight 

and magma-pressure force, which has declined during slip. Before 

this reequilibration, the shear force is out of equilibrium with 

these static forces because rate-weakening friction allows slip to 

dynamically overshoot the equilibrium point. The net effect is that 

almost the full force drop (~7×107 N) occurs abruptly at the end of 

the slip cycle, and it thereby provides an impulse capable of radiating high-frequency seismic energy.

Sensitivity of Behavior to Variations of 

Parameters Within D

Values of D encapsulate the effects of most of the 

important parameters affecting stick-slip dynamics, but it is 

nevertheless necessary to examine whether variations in values 

of these individual parameters have significant effects. Results 

presented in figures 15 and 16 show that the behavior computed for D= −2 is quite insensitive to variations in c, λ, and 

0 / ref u u . For example, if values of c and λ range over orders of 

magnitude, while D is held constant, computed stick-slip limit 

cycles differ only slightly in a phase-plane diagram depicting 

coevolution of the normalized pressure deviation and normalized slip velocity (fig. 15A). Differences exist exclusively in 

the magnitude of normalized pressure deviations, reflecting 

the fact that smaller λ values imply that less magma pressure is needed to satisfy limiting equilibrium and trigger slip. 

Moreover, if the pressure deviations are “denormalized,” as 

shown in figure 15B, such differences disappear entirely, and 

results for all values of c and λ collapse onto a single curve 

in the phase plane. This finding demonstrates that values of c

and λ individually are unimportant in the system’s dynamical 

behavior, provided that D is constant and the nonlinearity of 

the friction rule is unchanged.

Results shown in figure 16 illustrate the effects of 

changing the nonlinearity of the friction rule by allowing 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 447

12,937

12,936

12,935

12,934

12,933

632,160

632,120

632,080

632,040

632,000

631,960

Figure 14. Time series depiction of concurrent changes in plug displacement, slip velocity, magma pressure, and conduit volume 

during earthquake cycles computed for baseline case with D = −2. Parameter values are same as those used to generate figure 

13, and c = −1.71×10-4 is used to obtain D = −2. Panels A–D illustrate behavior during seven consecutive slip events, and panels 

E–H show details on an expanded time scale. Panel E is distinct from other panels because it shows force drops accompanying 

slip events.

the value of 0 / ref u u to range over four orders of magnitude 

while holding D= −2. To facilitate comparison with figure 13, 

figure 16 depicts results as both displacement time series and 

stick-slip limit cycles. These results show that reducing values 

of ref u (that is, increasing values of 0 / ref u u ) produces stickslip cycles with increased interevent periods, increased slip 

displacements, increased slip velocities, and increased deviations of magma pressure from its equilibrium value. However, 

effects of 0 / ref u u ranging over four orders of magnitude (fig. 

16) are similar to the effects of D ranging from about −2 to −4 

(fig. 13), reinforcing the view that D values encapsulate most 

of the dynamical controls on system behavior. It is, however, 

unsurprising that effects of 0 / ref u u are not captured entirely by 

values of D, as the effect of 0 / ref u u is inherently nonlinear. For 

values of 0 / ref u u smaller than 1, the periods of stick-slip oscillations lengthen, but this behavior is not pursued here because 

it has little relevance to plug extrusion at Mount St. Helens.

Effects of Plug Mass Change (Nonzero K)

Continuous changes in plug mass are represented by 

nonzero values of K (and of κ, its dimensional equivalent). 

If the condition B=Q is assumed, nonzero values of K imply 

that the rate of mass loss due to surface erosion (rE ) does 

not balance the rate of mass gain due to basal accretion (B), 

and the simplest example of such an imbalance occurs when 

the erosion rate is zero and the plug mass increases at the rate

 = B . Figure 17 illustrates time-series behavior computed 

for this case (that is, κ=4,000 kg/s; K=1.78×10-7) with D= 

−2 and 0 / 10 ref u u = . Comparison of figure 17 with figure 14 
448 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

-1,000

-1,500

-2,000

-2,500

-

Figure 15. Sensitivity of stick-slip behavior with D = −2 

to variations in parameters c and λ in the nonlinear rateweakening friction rule. Other parameter values are same 

as those used to generate figure 14. Only phase-plane 

representations are shown, because time series are identical 

in all cases. Arrows point in direction of advancing time. A, 

Stick-slip cycles represented in terms of normalized pressure 

deviations. B, The same stick-slip cycles represented in terms 

of physical pressure deviations.

shows that stick-slip cycles computed with K>0 have slightly 

larger periods and amplitudes than those computed with K=0, 

an unsurprising finding in view of analytical results indicating 

that positive K values will cause growth of oscillation periods 

(see equation 33). However, in contrast to analytical predictions, periods and amplitudes of stick-slip cycles computed 

with K>0 do not change with time.

The linear analytical model yields spurious predictions 

of the effects of positive K values because it neglects explicit 

coupling between magma-pressure change and conduit-volume change, which is included in the nonlinear computational 

model. Figure 17 shows that magma pressure rises slightly in 

each successive stick-slip cycle, because increased pressure 

is required to drive uplift of the increasingly massive plug. 

This increasing pressure is accompanied by slightly decreasing conduit volume (fig. 17), because volume change due to 

compression of the highly compliant magma exceeds volume 

change due to conduit-wall deflection. Moreover, the percent 

decrease in conduit volume at the end of each stick-slip cycle 

is precisely the same as the percent increase in plug mass 

during the same cycle. Therefore, the effects of changes in 

conduit volume counterbalance the effects of changes in plug 

mass, such that the net effect of these changes on the oscillation period is zero. (As shown by equation 25, the time scale 

for the natural oscillation period, 0t , depends on the product 

of plug mass and conduit volume.) Numerical results show 

that the counterbalancing effect of mass changes and volume 

changes occurs for all positive K values—within reason. As 

suggested by analytical results, however, very large K or κ

values (satisfying the criterion  > −(2 / 3)C ) might stabilize 

extrusion and diminish oscillations. Such large κ values are 

physically unreasonable, because C scales with the large 

quantity 0 / mg uref (equation 23B), and mass accretion at a 

commensurately large rate is not plausible.

The magma pressure build-up depicted in figure 17 is 

further illustrated by its phase-plane representation in figure 

18. As shown in the phase plane, each stick-slip limit cycle 

is like the preceding cycle, except that each successive slip 

event begins at a successively larger magma pressure. In this 

example, use of the maximum plausible rate of mass increase 

(that is,  = B = 4,000 kg/s) causes magma pressure to 

double after about 115,000 stick-slip cycles, or about 113 

days. To within 1 percent, this result agrees with the pressure 

doubling time calculated by applying the analytical formula 

derived for pseudosteady-state extrusion (that is, equation 

21) for the case in which κ = 4,000 kg/s. The chief implication of these results is that if the plug mass increases significantly, magma pressure can increase significantly without 

increasing the extrusion rate. Conversely, changes in extrusion rate need not be linked to changes in magma pressure if 

the plug mass changes.

Computational results obtained with K<0 differ in 

important ways from results obtained with K>0. Figure 19 

depicts results computed using a negative K value with the 

same magnitude as the positive K value used to generate 

figures 17 and 18 (that is, K= −1.78×10-7), and this value 

implies that the mass erosion rate, rE , is twice the rate of 

basal mass accretion, B . Comparison of figures 17 and 19 

shows that the negative K value produces changes in magma 

pressure and conduit volume with signs opposite to those for 

K>0 and also produces more frequent slip events (seven as 

opposed to six in 500 s). Slip events occur more frequently in 

the presence of diminishing plug mass, because less magma 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 449

1,000

1,000

-

Figure 16. Sensitivity of stick-slip behavior with D = −2 to variations in 0 / ref u u in the nonlinear rate-weakening friction rule. Other 

parameter values are the same as those used to generate figure 14. A, Time-series representation of stick-slip displacements. B, 

Phase-plane representation of velocity-pressure limit cycles corresponding to A. Arrows point in direction of advancing time.

pressure is necessary to trigger slip in each successive stickslip cycle. Importantly, the increased frequency of slip events 

is accompanied by an increase in the mean extrusion rate 

(which occurs with a constant magma supply rate, Q=2 m3/s), 

and this increase demonstrates the potential for runaway 

eruptive behavior. Moreover, decreasing plug mass eventually leads to a singularity in which fluid magma reaches the 

surface and the SPASM equations no longer apply (that is, 

the case with C<0 and κ<0 identified in table 3).

Runaway behavior is more evident in computational 

results obtained using negative K values with magnitudes 

much larger than the value −1.78×10-7 used to generate figure 

19. Unlike large positive K values, large negative K values 

are physically plausible, because the plug mass can decrease 

rapidly if large-scale spalling or avalanching erodes the plug 

surface. Figure 20 depicts displacement time series computed 

with negative K values having magnitudes 100, 1,000, and 

10,000 times greater than −1.78×10-7. (The value −1.78×10-3

corresponds to mass loss at a rate of 4×107 kg/s or a volumetric rate of about 20,000 m3

/s—a rate high enough to remove 

the entire plug in less than one hour.) The key point illustrated 

by figure 20 is that, in the presence of significant mass loss, 

time-averaged displacement rates can be much larger than 

the steady equilibrium rate 0 u QA = / = 6.667×10-5 m/s that 

prevails with K=0, and these high rates would likely lead to a 

change in eruptive style.

The unstable growth of extrusion rate that occurs with K= 

−1.78×10-3 is represented in the phase-plane diagram depicted 

in figure 21. The diagram shows that deviations from the equilibrium magma pressure and equilibrium slip rate are exceedingly large during each stick-slip cycle and that maximum 

slip rates increase during each successive slip cycle—while 

magma pressure successively declines. In such scenarios, then, 

extrusion occurs faster and faster until the plug is removed and 

liquid magma reaches the surface. The accompanying decline 

in magma pressure could also result in increased vesiculation 

and explosive potential, although this process is not represented by the SPASM model.

Effects of X

The dimensionless parameter X defined in equation 44 

mediates the interaction between extrusion rate and conduit 

volume change, as shown by equations 42 and 48, and it also 

affects magma pressure change through its influence on * R

(that is, equation 45). The physical meaning of X can be clarified by writing its definition in a simplified, approximate form 

that is valid if1 >> 2 , which is almost certainly the case at 

Mount St. Helens:

1/2

1 0

0

0

. m X u

V

  ≈    

 (55)

This definition implies that, for systems in which 0 u

and 1 are constant, variations in X can be viewed as scaled 

variations in 0 0 m /V . Indeed, because values of 0 0 m /V can be 

changed while holding 0t and D constant, this strategy is used 
450 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

12,937

12,936

12,935

12,934

12,933

632,200

632,150

632,100

632,050

632,000

631,950

Figure 17. Time-series behavior of solutions computed with 

D= −2, u0

/uref =10, and increasing plug mass specified by 

K=1.78×10-7 (that is, κ =4,000 kg/s). Other parameter values are 

same as those used to generate figure 14.

-

Figure 18. Phase-plane representation of solution depicted 

in figure 17. Numbers shown in red denote time sequence of 

successive limit cycles. Arrows point in direction of advancing time.

to assess the effects of variations in X computationally. The 

condition K=0 is also assumed in this assessment.

Computations in which X is increased or decreased 

by one order of magnitude from its baseline value (5×10-6) 

show that stick-slip time series and phase-plane diagrams 

are identical for all values of X. However, figure 22 shows 

that some important details of the earthquake cycle change 

when X is increased by an order of magnitude (to 5×10-5) 

by simultaneously increasing m0 and decreasing V0 by one 

order of magnitude each from their baseline values, while 

retaining D= −2. Comparison of figure 22 with figure 14 (the 

baseline case) shows that, although the timing and magnitude 

of slip events is unchanged when X is increased by one order 

of magnitude, the accompanying magma pressure, pressure 

change, shear force, and force drop are each increased by 

one order of magnitude, whereas the conduit volume change 

is reduced by somewhat less than an order of magnitude. 

These effects are all logical consequences of the increased 

plug mass and reduced conduit volume that are imposed by 

increasing the value of X.

The most important fact illustrated by figure 22 is 

that the predicted magma pressure at the base of the plug 

(~1.3×108 Pa) is roughly an order of magnitude larger than 

the expected lithostatic pressure near the plug base (that is, 

at depths ~500 m). Such a large magma pressure is unrealistic, as it would probably cause hydraulic fracturing and a 

marked change in eruption style. This result implies that a 

value X≈5×10-5 is too large to be realistic. Similarly, if X is 

reduced to 5×10-7 while holding t

0 and D constant, computed 

basal magma pressures are an order of magnitude smaller 

than the likely lithostatic pressure—a result that is also 

unrealistic. These findings imply that an X value similar 

to the baseline value (5×10-6) is probably appropriate for 

Mount St. Helens. Therefore, the computed drop in shear 

force accompanying slip in the baseline case (fig. 14) is also 

probably appropriate.

Effects of Y

The dimensionless parameter Y defined in equation 

44 represents the scaled difference between the magma-
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 451

1,000

Figure 21. Phase-plane representation of stick-slip limit cycles 

occurring with D = −2, u0/uref =10, and rapid decrease in plug 

mass (K = −1.78×10-3; κ = −4×107

 kg/s). Other parameter values 

are same as those used to generate figure 14. Arrows point in 

direction of advancing time.

Figure 20. Comparison of displacement time series computed 

with D = −2, 0 ref u u/ =10, and differing rates of decreasing plug 

mass specified by differing values of K. Other parameter values 

are same as those used to generate figure 14.

12,937

12,936

12,935

12,934

12,933

632,200

632,150

632,100

632,050

632,000

631,950

Figure 19. Time-series behavior of solutions computed with 

D= −2, u0/uref =10, and decreasing plug mass specified by K = 

−1.78×10-7 (that is, κ = −4,000 kg/s). Other parameter values are 

same as those used to generate figure 14.

influx rate Q and magma-solidification rate B. All results 

presented thus far assume that B=Q, which is necessary 

for the magma-plug system to long remain close to equilibrium. Indeed, if BQ for a sustained period, a transition 

in eruptive behavior is inevitable. If B>Q, for example, 

the solidification front would propagate downward and the 

plug mass would increase unless mass loss due to surface 

erosion balances mass gain due to basal accretion. If such 

a balance were sustained, the plug would appear to sink 

progressively, even as extrusion continued, and friction 

on the plug margins would progressively increase. On the 

other hand, if B<Q, the solidification front would migrate 
452 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

12,937

12,936

12,935

12,934

12,933

632,000

631,600

631,200

630,800

Figure 23. Time series behavior of solution computed with D = −2, 

u0/uref =10, and excess basal accretion specified by B = 2Q = 4 m3/s 

and K = 1.78×10-7 (that is, κ = 4,000 kg/s). Here Y = −5×10−6. Other 

parameter values are same as those used to generate figure 14.

129,370

129,360

129,350

129,340

129,330

63,360

63,320

63,280

63,240

63,200

63,160

Figure 22. Details of an earthquake cycle computed with D = −2, 

u0

/uref =10, and X = 5×10-5 (one order of magnitude larger than the X 

value used to compute the baseline results shown in figure 14). Other 

parameter values are same as those used to generate figure 14.

upward. In the extreme case of B=0, any changes in plug 

mass would result exclusively from surface erosion, and the 

plug would eventually be pushed out the ground as liquid 

magma reached the surface. This scenario represents singular behavior analogous to that occurring with a negative 

mass-change rate (K<0). Below, I focus only on short-time 

behavior for cases with B>Q and B=0.

Behavior computed for a case with B=2Q=4 m3/s and 

K=1.78×10-7 (that is, κ=4,000 kg/s) is shown in figure 23; this 

case is exactly like that illustrated in figure 17 except that, 

here, growth of plug mass occurs as a result of basal accretion in excess of Q. Thus, figure 23 depicts the response to 

growth of mass exclusively below the surface (with downward 

migration of the solidification front), whereas figure 17 depicts 

the response to growth exclusively above the surface (with no 

downward migration of the solidification front). The results 

shown in figures 23 and 17 are in most respects identical, except 

that the conduit volume decreases much more quickly in the 

case with basal accretion in excess of Q (fig. 23)—an obvious 

consequence of downward migration of the solidification front. 

Increasing pressurization of the conduit fluid accompanies 

this migration, and in the long run this combination of effects 

is unsustainable because magma pressure eventually would 

become large enough to fracture surrounding rock.
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 453

12,937

12,936

12,935

12,934

12,933

633,200

632,800

632,400

632,000

Figure 24. Time series behavior of solution computed with 

D = −2, u0

/uref =10, and zero basal accretion (B = 0). Here 

Y = 5×10−6. Other parameter values are same as those used to 

generate figure 14.

Figure 25. Displacement time series computed with varying 

values of R0 and Z (with accompanying variation of D and X). 

Other parameter values are same as those used to generate 

figure 14.

The limiting case with small B (that is, B=0) exhibits behavior almost exactly like that computed for the baseline case with 

B=Q and K=0. Indeed, with B=0, graphs of the behavior for early 

times are indistinguishable from those shown in figures 13 and 14, 

except that conduit volume increases with each successive stickslip cycle (fig. 24). This increase is an obvious consequence of 

upward migration of the base of the plug, and although this migration affects plug dynamics negligibly in the short term, persistence 

of this migration eventually must lead to the singularity noted 

above (that is, magma reaching the surface).

Effects of Z

The dimensionless parameter Z defined in equation 44 

may be interpreted as a scaled version of the mass-density 

ratio difference, 0 0 1( / ) R = −  r , which plays its most important role in determining the mean (that is, time-averaged) 

extrusion rate, 0 0 u Q RB A = ( )/ − given by equation 19. The 

mean extrusion rate is affected by Z because R0 determines 

the change in density that occurs as magma ascending at rate 

Q solidifies at rate B—such that an increase in density during 

solidification reduces the volumetric extrusion rate of the solid 

plug. Values of R0 plausibly range from about −0.5 to 0.5, and 

for the case with B=Q, these values yield 0 u values ranging 

from 0.5(Q/A) to 1.5(Q/A). Moreover, with B=Q, the definition of Z also reduces to 0 0 Z = R R /(1 ) − , so that Z depends 

exclusively on R0. Examples of displacement time-series 

solutions computed for this case, with various values of Z, are 

shown in figure 25.

Interpretation of figure 25 is complicated by the fact 

that with A and B held constant, values of Z cannot be varied 

independently of values of D and X (because Z, X, and D all 

depend on 0 u , which in turn depends on R0 ). Therefore, the 

figure shows results of computations in which Z, X, and D

vary simultaneously. Nevertheless, the effect of values of Z on 

the average extrusion rate is clearly evident in figure 25; large 

departures from the baseline value Z=0 yield similarly large 

departures from the average extrusion rate observed with Z=0 

(that is, 0 u Q= / A). The period of stick-slip cycles remains 

nearly proportional to the magnitude of D (just as in the baseline case illustrated in fig. 13), despite variations in Z. Unlike 
454 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 27. Phase-plane representation of extrusion behavior 

computed with an initial excess magma pressure (0.005 percent 

of p0

) and D = 0.2. Parameter values are same as those used to 

generate figure 14, except that c = 1.71×10-5 is used to obtain D = 

0.2. Arrows point in direction of advancing time.

Figure 26. Phase-plane representation of behavior computed 

with four disequilibrium initial conditions (that is, initial excess 

magma pressures 0.005 percent, 0.01 percent, 0.015 percent, and 

0.02 percent of p0

), for each of two negative values of D. Parameter 

values are same as those used to generate figure 14, except that c

= −1.71×10-5 is used to obtain D = −0.2. Arrows point in direction of 

advancing time. A, D = −0.2. B, D = −2.

the baseline case, however, the time series depicted in figure 

25 each involve slip events with about the same magnitude of 

displacement (~5 mm). This behavior indicates that, while D

largely controls the system’s dynamics, the magnitude of slip 

events is additionally regulated by Z.

Effects of Initial Conditions

The preceding results were computed using initial 

conditions that assume a state of mechanical equilibrium, but 

volcanic eruptions presumably begin with an initial disequilibrium state, such as that due to magma pressure in excess of 

the static limiting equilibrium pressure ( 0 p > p ). This section 

summarizes results of computations that used a disequilibrium 

initial condition of this type.

All computations with 0 p > p predict that an initial pulse 

of rapid motion occurs until plug momentum is depleted, 

magma pressure relaxes, and static equilibrium is restored, and 

this behavior occurs irrespective of the sign or value of D (figs. 

26, 27). Pressure then rebuilds until it triggers a second stage 

of motion. For D<0 this stage consists of endlessly repetitive 

stick-slip limit cycles (closed loops in fig. 26) exactly like 

those produced with equilibrium initial conditions (for example, fig. 13), whereas for D>0 the second stage converges to a 

fixed-point equilibrium representing a state with dynamically 

balanced forces (fig. 27). It is noteworthy, however, that cases 

with rate-strengthening friction (D>0) and rate-weakening 

friction (D<0) exhibit similar initial pulses if the initial pressure disequilibrium is the same (for example, compare results 

for D=±0.2 in figs. 26A and 27).

A key point illustrated in figure 26 is that the maximum 

speed of the initial extrusion pulse and the magnitude of the 

associated pressure deviation increase almost linearly with 

increasing initial magma overpressure. Moreover, this trend 

is insensitive to the value of D, and linear extrapolation can 

therefore be used to infer the maximum extrusion speed (and 

pressure deviation) associated with any initial overpressure. 

With D= −2, for example, each increase of 0.005 percent in 

initial overpressure increases the maximum speed of the initial 

movement pulse max u by about max u QA /( / ) 14 = (fig. 26B). 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 455

In conjunction with the baseline value Q/A=6.67×10-5 m/s 

for Mount St. Helens, this result implies that a 0.005 percent 

magma overpressure (relative to the static limiting equilibrium 

pressure 0 p ) would produce a maximum extrusion velocity 

of about 3 mm/s, and that an overpressure of 5 percent would 

produce a maximum extrusion velocity of ~1 m/s.

Magnitudes of displacements that occur during the initial 

movement pulses shown in figures 26 and 27 can be estimated 

from the maximum pressure deviation 0 p − p , because the 

displacement during slip is proportional to the duration of the 

subsequent “stick time” (as shown by equation 54). By inference from equation 54, then, the displacement during the initial 

slip event increases like 1 2 0 ( )( )i p p e  + − . On this basis, extrapolation from figure 26B shows that initial magma overpressures of 

even a few percent would lead to movement pulses of at least 

several meters, provided that D= −2.

Discussion

The SPASM model was developed to explain the relation 

between nearly steady solid-state extrusion and nearly periodic 

drumbeat earthquakes observed during the 2004–5 eruption 

of Mount St. Helens. Although this relation could result from 

a complicated interaction of numerous physical and chemical processes, the SPASM model aims for a parsimonious 

mechanical explanation that employs a minimum of postulates 

and variables. Some of the parsimony of the SPASM model 

derives from its central, simplifying assumption: that the 

volumetric flux of magma into the base of the eruption conduit 

(Q) is constant. This assumption is consistent with a top-down 

perspective of eruption dynamics, in which no time-dependent 

changes in a deep magmatic system are invoked to explain 

phenomena observed at Earth’s surface.

From a mathematical standpoint, the assumption of constant Q enables the SPASM model to be reduced to a nonlinear 

system of three first-order ordinary differential equations. The 

relative simplicity of these equations facilitates both analytical 

and numerical studies of model properties. Indeed, the mathematical behavior of the SPASM model can be understood 

completely, and understanding of this simple model provides 

a steppingstone toward understanding behavior of Mount St. 

Helens during solid-state extrusion in 2004–5.

Analytical study of the SPASM model demonstrates that 

steady, solid-state extrusion is an equilibrium condition that 

can be satisfied exactly if the magma influx rate, Q, equals the 

rate of magma solidification at the base of an extruding plug, 

B, which in turn is balanced by erosion of the plug surface, 

E. In addition, pseudosteady states can exist in which the 

extrusion rate remains constant and changes in plug mass are 

accompanied by compensating changes in magma pressure. 

Such pseudosteady states cannot persist indefinitely, however, 

because changes in plug mass and magma pressure cannot 

continue without eventually causing a change in eruption 

style. Therefore, long-term steady-state extrusion at Mount St. 

Helens in 2004–5 probably involved a near-equilibrium state 

in which the plug mass effectively remained almost constant 

and B ≈ Q applied.

Analysis of the SPASM model shows that, even when a 

near-steady eruptive state persists, extrusion velocities have an 

inherent tendency to exhibit short-term oscillations about the 

long-term equilibrium rate. These oscillations are an inevitable 

consequence of the interaction of plug momentum, a variable upward force due to magma compression and pressure, 

and a downward force due to the plug weight. The oscillations are damped mostly by plug-boundary friction, and if 

friction exhibits rate-weakening behavior like that observed 

in experiments with fault gouge obtained from the surface of 

the Mount St. Helens plug, damping is negative and oscillation amplitudes grow unstably. Oscillation growth is necessarily arrested, however, because friction has the potential to 

reverse its direction of action when the plug extrusion velocity 

declines to zero. Growing oscillations are thereby transformed 

to repetitive stick-slip cycles, and these cycles continue indefinitely (that is, until a change in a parameter such as Q, B, or 

E occurs). According to the SPASM model, these repetitive 

stick-slip cycles are responsible for generating nearly periodic 

“drumbeat” earthquakes observed during solid-state extrusion 

at Mount St. Helens in 2004–5.

Properties of stick-slip cycles predicted by the SPASM 

model are controlled by a variety of factors, but both analytical 

and numerical results support the conclusion that the natural 

oscillation time scale 0t and dimensionless damping D exert 

the most important controls. The oscillation time scale is fixed 

by conduit and plug properties that affect elastic strain in the 

system, whereas nonlinearly rate-dependent friction causes 

damping to vary as a function of the extrusion velocity u. 

Nevertheless, computational results show that most effects 

of variable damping are encapsulated by D evaluated at the 

steady equilibrium extrusion rate, 0 u = u . These results show 

that the amplitudes and periods of stick-slip cycles increase 

almost in proportion to the magnitude of D, provided that 0t

remains constant.

Computations using diverse values of D show that D= 

−2 produces stick-slip cycles most similar to those inferred 

to generate drumbeat earthquakes at Mount St. Helens 

throughout much of the 2004–5 eruption, and the dynamics of these cycles are relatively insensitive to variations in 

values of the physical parameters that constitute D. With 

D= −2 , computed interevent periods are about 80 s, the slip 

distance per event is about 5 mm, maximum slip speeds are 

about 2 mm/s, and reduction of magma pressure during slip 

is about 2.4 kPa. This reduction in magma pressure is strikingly small in comparison to the ambient, roughly lithostatic 

magma pressure (~1.3×104

 kPa) inferred to exist at the base 

of the plug, about 500 m beneath the ground surface. The 

small size of magma-pressure fluctuations indicates that the 

system deviates little from mechanical equilibrium, even during slip events. Moreover, changes in basal magma pressure 

during slip events can be multiplied by the inferred basal 

area of the plug (~30,000 m2

) to provide a proxy for the force 

drop available to generate seismicity.
456 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A more refined assessment of the force drop during slip 

events is provided by SPASM output that shows how the shear 

force due to plug-boundary friction (F) evolves during slip. 

With D= −2, the computed force drop is about 7×107 N, and 

much of the force drop occurs abruptly, despite the fact that 

slip events have durations of ~5 s. This nearly instantaneous 

force drop occurs just as the slip velocity, u, decreases from a 

finite value to zero, and its abruptness results from the interplay of three phenomena: (1) owing to rate-weakening friction 

during slip, the upward-moving plug gains momentum that 

causes it to overshoot an equilibrium state in which forces are 

dynamically balanced; (2) as the upward-moving plug decelerates, a potential for downward motion exists, because basal 

magma pressure has relaxed to a value smaller than its static 

equilibrium value; and (3) incipient downward motion of the 

plug is arrested immediately, however, because friction along 

the plug margin reverses direction so as to oppose motion. 

Friction adjusts just enough to stop motion and balance forces, 

of course, because friction can only oppose motion, not drive 

it. The abrupt adjustment of the friction force as plug motion 

ceases produces a sudden force drop that can radiate highfrequency seismic energy. 

The mechanics that produce the stick-slip behavior and 

force drops predicted by the SPASM model are robust because 

they derive from basic physical principles, but the weakest 

link in the SPASM formulation involves the nature of friction 

and its rate dependence. Frictional properties of granulated 

solids such as fault gouge are poorly understood at a fundamental level, although some degree and type of rate weakening is generally observed in a variety of both idealized and 

geological granular media (for example, Nasuno and others, 

1997; Marone, 1998). Rate-weakening behavior has also been 

measured in tests of fault gouge collected from the surface of 

the extruding plug at Mount St. Helens (Moore and others, this 

volume, chap. 20). In the context of the SPASM model, the 

most crucial point revealed by these tests is that some degree 

of rate-weakening occurs as slip velocities increase from zero 

to a steady equilibrium value 0 u . This weakening suffices to 

generate stick-slip behavior, regardless of subtler nuances of 

friction. If rate-strengthening friction develops at higher slip 

rates, for example, it will help arrest slip events but not prevent 

them. Therefore, the occurrence of stick-slip cycles appears 

almost inevitable.

Occurrence of stick-slip cycles large enough and abrupt 

enough to generate drumbeat earthquakes is also contingent 

on elastic properties of the magma-plug-conduit system. At 

least one elastic element in the system must be soft enough 

to strain significantly in response to driving forces that are 

smaller than those required to shear the fault gouge irreversibly (compare Rice and Ruina, 1983). (The shallow depths 

of drumbeat earthquakes implies that these driving forces are 

probably smaller than those causing fault slip at hypocentral 

depths typical of nonvolcanic earthquakes (that is, >1 km), 

because confining stress and frictional resistance are relatively small at depths <1 km.) At Mount St. Helens, fluid 

magma underlying the extruding plug provides an exceptionally soft elastic element, because its estimated 12 percent (by 

volume) exsolved gas content makes it orders of magnitude 

more compressible than solid rock. In essence, then, the 

magma serves as a spongy spring that compresses significantly as it delivers the force to shear the plug-bounding 

gouge. Strain energy stored during magma compression is 

released in abrupt slip events that would be smaller and more 

frequent if the magma were stiffer. Therefore, the earthquake 

cycle described by the SPASM model differs from a typical tectonic earthquake cycle in two important ways: (1) in 

SPASM, strain energy is stored principally in a compressed 

fluid, not in a solid deformed in shear; repeated, nondestructive compression of “soft” fluid enables seismogenic 

plug slip to occur repeatedly at shallow depths; and (2) in 

SPASM, after slip has ceased, the plug is reloaded by forces 

due to gravity and magma influx. This reloading occurs 

very rapidly in comparison to reloading by tectonic strain 

accumulation, enabling drumbeat earthquakes to occur much 

more frequently than tectonic earthquakes.

Gradual evolution of the magnitude and periodicity of 

drumbeat earthquakes observed during the 2004–5 eruption 

of Mount St. Helens prompts questions about the cause. Can 

evolution of drumbeats be a harbinger of changes in eruption rate or style? The SPASM model provides insight to this 

issue, but it provides no unequivocal answers. According to 

the model, amplitudes and periods of slip events can change 

in response to changes in values of any of the parameters 

that constitute D, even when magma ascent and solidification rates are constant. Because D encompasses the effects of 

gouge frictional properties and the effective stress state, as 

well as the natural oscillation time scale 0t , evolution of any 

of a number of phenomena could be responsible for changing the character of drumbeats. Nonetheless, because field 

data can to some degree constrain changes in quantities such 

as plug mass or conduit volume, it is tempting to ascribe 

changes in drumbeats to unobserved changes in gouge properties or the state of effective stress. Computational results 

from SPASM indicate that rather subtle changes in these phenomena can alter drumbeat properties significantly. Indeed, it 

is easy to imagine that the source of drumbeats could migrate 

around the periphery of the extruding plug as scattered 

patches of particularly strong gouge form, fail, and reform. 

A model more elaborate than SPASM would be required to 

analyze the details of such behavior.

The SPASM model does provide a clear picture of potential changes in eruption style that can occur if conditions far 

from equilibrium develop. One type of disequilibrium develops if the rate of magma accretion at the base of the extruding 

plug differs significantly from the rate of plug extrusion. If 

the accretion rate is less than the extrusion rate, liquid magma 

eventually will reach the ground surface. Magma pressure will 

simultaneously decline, and this decline could enhance magma 

vesiculation and explosive potential. On the other hand, if the 

plug accretion rate exceeds the extrusion rate, the solidification front will propagate downward and magma pressure will 

increase, assuming that deep magma influx remains constant. 
21. Dynamics of Seismogenic Volcanic Extrusion Resisted by a Solid Surface Plug, Mount St. Helens, 2004–2005 457

Increasing pressure could eventually trigger an exceptionally large slip event or fracture surrounding rock, causing a 

transition in eruption style—or it could suffice to stop magma 

influx, thereby halting the eruption.

Effects of magma pressure exceeding the static limiting 

equilibrium pressure can be treated as disequilibrium initial 

conditions in the SPASM model. If excess pressure exists 

but is insufficient to fracture rock, a rapid pulse of plug 

motion occurs until static equilibrium is restored. If magma 

pressure exceeds the equilibrium pressure by even a few 

percent, SPASM predicts that this pulse can involve velocities of meters per second and displacements of many meters. 

Such a pulse could even eject the plug from the volcanic 

vent and instigate the type of transition in eruption style 

described above.

The fact that no rapid pulses of extrusion or transitions in 

eruption style occurred during the 2004–5 eruption of Mount 

St. Helens implies that the magma-plug system never deviated much from equilibrium, even during the eruption onset. 

This behavior is, of course, very different from that during the 

cataclysmic eruption of Mount St. Helens on May 18, 1980. 

In that case explosive activity was caused by rapid depressurization of a shallow magma body triggered by a massive 

landslide. Without the landslide trigger, the 1980 eruptions of 

Mount St. Helens might have been relatively quiescent, much 

like the eruption of 2004–5.

Conclusion

The central conclusion of this study is that stick-slip 

oscillations are almost inevitable during an eruption in 

which steady ascent of compressible magma drives upward 

extrusion of a solidified plug with margins that exhibit rateweakening friction. Whether such oscillations are large and 

abrupt enough to generate repetitive earthquakes like those 

observed at Mount St. Helens depends on a host of factors, 

nearly all of which are encompassed within two quantities derived in this paper: the natural oscillation time scale 1/2

0 0 1 20 tm V A = [ ( )] /  + and the frictional damping index, 

which can be approximated as ( ) 000 D c gt u ≈ (1/ 2) /   . 

Large values of 0t favor the occurrence of relatively large, 

infrequent slip events, because they imply that large elastic 

strains can be accommodated during magma compression. 

Large negative values of D have similar effects because 

they imply that effects of rate-weakening plug friction are 

significant, and rate weakening is responsible for “dynamic 

overshoot” during slip events. As a consequence, negative D

values far from 0 cause the period between successive slip 

effects to exceed the period expected on the basis of 0t alone.

Computations using D= −2 predict the occurrence of 

stick-slip cycles with interevent periods of ~80 s, slip displacements of ~5 mm, and force drops of ~7×107

 N, and these 

properties appear consistent with those of events inferred to 

produce drumbeat earthquakes during the 2004–5 eruption of 

Mount St. Helens. Although individual modeled slip events 

last about 5 s, most of the accompanying force drop occurs in 

a fraction of a second, consistent with requirements for radiation of high-frequency seismic energy.

Persistence of nearly periodic drumbeat earthquakes also 

requires that slip events are driven by nearly steady forcing. At 

Mount St. Helens this forcing was provided by nearly steady 

ascent of magma. Magma solidification at the base of the 

extruding plug apparently occurred at a rate nearly equal to the 

rate of magma ascent, enabling the system to remain close to 

equilibrium. Indeed, model results show that a near-balance 

between ascent rate and solidification rate is essential for 

maintaining persistent drumbeats.

The presence of a near-surface body of compressible 

magma that serves as a driving element may be necessary 

to generate repetitive, seismogenic stick-slip events at the 

shallow focal depths (<1 km) observed during the 2004–5 

eruption of Mount St. Helens. The strength of gouge at such 

shallow depths is relatively small (owing to relatively small 

normal stresses), and the gouge can therefore shear irreversibly before much elastic strain accumulates in a stiff adjacent 

body such as solid rock. Therefore, in the absence of a soft, 

near-surface magma body, stick-slip oscillations could still 

occur, but they would be reduced in size and period, perhaps 

to a degree that would make them aseismic. Moreover, nearsurface fluid magma can undergo repeated elastic compression and decompression without accumulation of irreversible 

damage that would likely accompany similarly repetitive 

strain in solid rock.

Lack of large movement pulses during the 2004–5 

eruption of Mount St. Helens reinforces the view that the 

magma-plug system remained close to equilibrium. Indeed, 

model results indicate that magma pressure exceeding the 

static equilibrium pressure by even a few percent was probably never present. An implication of this finding is that the 

dynamic equilibrium state exhibited during the eruption differs 

little from the static equilibrium state before the eruption 

onset. Therefore, the eruption trigger was likely very subtle, 

perhaps involving nothing more than weakening of the conduit 

cap rock by percolating water derived from late summer rains 

and glacier melt.

Acknowledgments

Virtually every member of the staffs of the Cascades 

Volcano Observatory and Pacific Northwest Seismic Network 

helped make this work possible through their contributions 

to monitoring the 2004–5 eruption of Mount St. Helens. I am 

indebted to all of them for their extraordinary efforts, and particularly to Seth Moran for fielding my many questions about 

volcano seismology. I am also indebted to Nico Gray of the 

Department of Mathematics, University of Manchester, U.K., 

who showed me how to obtain the analytical results summa-
458 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

rized in equations 33–37 and appendix 1, and to Peter Moore 

and Neal Iverson, Department of Geological and Atmospheric 

Sciences, Iowa State University, who performed laboratory 

tests of the frictional properties of the Mount St. Helens 

gouge. Roger Denlinger and Joseph Walder provided insightful reviews that helped improve the manuscript.

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460 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Solution of Equation 27 by Transformation to Bessell’s Equation

Conversion of the homogenous part of equation 27 into Bessel’s equation is accomplished by using a simultaneous change 

of independent and dependent variables, given by

2 2

* 1 1 . 2

K D

K K z * z t w u K 2

−

     = −=            

 (A1)

Substitution of equation group A1 into equation 27 converts it into Bessel’s differential equation

2

2 2 2

2 ( ) 0, d w dw z z z w 

dz dz

+ +− =  (A2)

where

2 . K D K

 − = (A3)

The general solution of Bessel’s equation is

1 2 w cJ z cY z ( ) ( ), =  +  (A4)

where 1 c and 2 c are arbitrary constants and J and Y are Bessel functions of the first and second kind, of order  (Abramowitz 

and Stegun, 1964). Equation 33 is obtained by transforming equation A4 back to the original variables *

t and * u and adding a 

particular solution of the inhomogeneous version of equation 27, * u KG = − .
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, Reston, VA 20192; now at Southern Methodist 

University, Office of Research and Graduate Studies, P.O. Box 750240, 

Dallas, Texas 75275

Chapter 22

Constraints on the Size, Overpressure, and Volatile 

Content of the Mount St. Helens Magma System from 

Geodetic and Dome-Growth Measurements During 

the 2004–2006+ Eruption

By Larry G. Mastin1, Evelyn Roeloffs1, Nick M. Beeler1, and James E. Quick2

Abstract

During the ongoing eruption at Mount St. Helens, Washington, lava has extruded continuously at a rate that decreased 

from ~7–9 m3/s in October 2004 to 1–2 m3/s by December 

2005. The volume loss in the magma reservoir estimated from 

the geodetic data, 1.6–2.7×107 m3, is only a few tens of percent 

of the 7.5×107 m3 volume that had erupted by the end of 2005.

In this paper we use geodetic models to constrain the size 

and depth of the magma reservoir. We also ask whether the relations between extruded volume and geodetic deflation volume 

are consistent with drainage of a reservoir of compressible 

magma within a linearly elastic host rock. Finally, we compare the time histories of extrusion and geodetic deflation with 

idealized models of such a reservoir. Critical parameters include 

erupted volume Ve, dome density e, reservoir volume VC , initial 

reservoir overpressure 0

ex p , pressure drop during the eruption ∆p, 

reservoir compressibility C≡ (1/VC)(dVC / dp), magma density 

M, and magma compressibility M ≡(1/M)(dM /dp). Seismic 

velocity and reservoir geometry suggest C ≈ 2×10–11 Pa–1, but 

mechanical considerations suggest C =7–15×10–11 Pa–1.

The geodetic data are best fit with an ellipsoidal source 

whose top is 5±1 km deep and whose base is ~10–20+ km 

deep. In the absence of recharge, the decrease in magma-reservoir 

volume dVC is theoretically related to the erupted volume Ve by 

Ve /dVC=(M /e)(1+ M /C). For C=~7–15×10–11 Pa–1 and M ≈e, 

estimates of Ve and dVC suggest that M =1.4–3.0×10–10 Pa–1, corresponding to a magmatic gas content in the reservoir of vg=0 to 

1.8 percent by volume.

If we assume that effusion rate is linearly related to reservoir pressure and that the recharge rate into the reservoir is 

constant, the effusion rate should decrease exponentially with 

time to a value that equals the recharge rate. Best-fit curves of 

this form suggest recharge rates of 1.2–1.3 m3/s over the first 

500 days of the eruption. The best-fit constants include the 

product 0 ( ) ex V p C C M   + , making it possible to constrain 

reservoir volume using values of C and M constrained from 

ratios of erupted volume to geodetic deflation volume. If, on 

the other hand, we assume a logarithmic pressure-effusion 

rate relation and a constant recharge rate, the dome volumetime curve should follow a modified logarithmic relation, 

with the total erupted volume at a given time proportional to 

V p C C M ∆ + ( )   . Using C =7–15×10–11 Pa–1, results from log 

and exponential curves suggest a reservoir volume of at least 

several cubic kilometers if ∆p or p0

ex is less than ~30 MPa. 

Similar results are obtained from numerical calculations that 

consider temporal changes in (1) magma compressibility, (2) 

the weight of the lava dome suppressing effusion, and (3) 

recharge rate. These results are consistent with the notion 

that the reservoir volume is at least a few times larger than 

the largest Holocene eruption of Mount St. Helens (4 km3

dense-rock-equivalent + volume for the 3.4-ka Yn eruption).

Both the exponential and logarithmic models predict a 

history of reservoir decompression that imperfectly matches 

displacement data at GPS station JRO1. Neither model, for 

example, predicts the rapid radially inward movement at 

JRO1 during the first month of the eruption. Such movement, followed by long-term linear deflation, suggests that 

erupted magma has been replaced in increasing proportions 

by recharge, but that the recharge rate remains somewhat less 

than the current (early 2006) effusion rate.
462 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 1. Geodetic stations and results. A, Map showing geodetic stations used to model the source of deflation at 

Mount St. Helens. B, Displacement radial to the Mount St. Helens crater versus time at geodetic stations. Outward radial 

displacements are positive. Error bars represent one standard deviation above and below the data point. Numbers 

beneath each station label indicate the map distance of each station from the crater center (46.2002° N, 122.1911° W).

-20

0

20

40

60

80

100

120

140

2004.25 2004.5 2004.75 2005 2005.25 2005.5 2005.75 2006 2006.25

RADIAL DISPLACEMENT, IN MILLIMETERS

P702

16.3 km

TSTU

4.8 km

JRO1

8.6 km

P698

3.7 km

P696

3.1 km

2.9 km

P687

16.2 km

B

*

*

*

*

*

*

* *

*

*

*

*

122°20' 122°10' 122°0'

46°10'

46°20'

P702

P687

JRO1

TSTU

TGAU

P693

TWRI

P698

LVCY

P696

P695

TWIW

A

DATE, IN YEARS

Mount St. Helens GPS

relative to fixed North America

0 10 KILOMETERS

TWIW

TGAU

2.1 km

AREA

OF MAP

Washington
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 463

Introduction

The current Mount St. Helens eruption extruded about 

85 percent as much lava in 14 months (October 2004 to 

December 2005) as was extruded in 6 years to form the 

1980–86 lava dome. At the time of writing (spring 2006), 

there is no obvious indication of waning growth; indeed, 

compared with other dome-building eruptions worldwide, the 

current eruption of Mount St. Helens only slightly exceeds 

the median duration (540 days) and volume (3.5×107 m3) 

among historical dome-building eruptions (Newhall and Melson, 1983, and updates, C. Newhall, written commun., 2005). 

This eruption hardly approaches the duration (>50 years) 

of the dome building at Bezymianny Volcano, Kamchatka, 

whose growth followed a sector collapse and lateral blast in 

1956 that resembled the 1980 Mount St. Helens sequence.

Dome-building eruptions, like eruptions of mafic lava 

flows, range in duration from days to decades and in volume 

from cubic meters to cubic kilometers. Eruption volume 

and duration are most strongly controlled by reservoir size, 

exsolved gas content, the amount of recharge into the magma 

system, and the composition and volatile content of recharging 

magma (Wadge, 1981; Newhall and Melson, 1983; Huppert 

and Woods, 2002; Woods and Huppert, 2003). In this study we 

use geodetic and growth-history data to constrain the volume, 

overpressure, history of recharge, and exsolved volatile content of the Mount St. Helens magma system.

Key Observations

Over several years prior to September 2004, campaignstyle geodetic surveys and measurements at the only continuous global positioning system (CGPS) site at Mount St. 

Helens (JRO1, fig. 1A) showed no discernable inflation signal 

(Dzurisin and others, this volume, chap. 14; Lisowski and 

others, this volume, chap. 15). Deflation at JRO1 began with 

the onset of seismicity on September 23, 2004; uplift on the 

crater floor was first noticed on about September 26 (Dzurisin and others, this volume, chap. 14) and the first material 

at magmatic temperature on October 11. Eight new CGPS 

instruments installed between mid-October and early November 2004 (fig. 1A) have recorded more or less radially inward 

movement toward the crater (fig. 1), suggesting that the source 

of deflation (the magma reservoir) is roughly spherical or 

elliptical in shape rather than dike-shaped (Lisowski and others, this volume, chap. 15).

Throughout the eruption, lava has extruded as spines or 

lobes of more or less solid, nearly holocrystalline lava, their 

surfaces covered by unconsolidated fault gouge on the order of 

1 m thick (fig. 2). The gouge thickness, considering relations 

between gouge thickness and displacement on tectonic faults 

(Robertson, 1983; Power and others, 1988), and petrologic 

information (Cashman and others, this volume, chap. 19) suggest that the faulting may extend from perhaps tens of meters 

to several hundred meters into the subsurface.

Gouge

~50 m

B C

A

B

C

Figure 2. Photos of the growing lava dome of Mount St. Helens. 

A, View of spine 4 from the northwest on February 22, 2005. 

USGS photo by S.P. Schilling. B, Close-up of striated fault gouge 

covering the exterior of spine 3 as it emerged from the ground 

on November 11, 2004. Approximate location of that spine on 

February 22, 2005, shown by box B on panel A. USGS photo by 

J.S. Pallister. C, Close-up of fresh rockfall scar (~100 m long 

and 50 m high), revealing gouge thickness in cross section. 

Approximate location shown by box C on panel A. USGS photo 

taken February 22, 2005, by S.P. Schilling.
464 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Geodetic Constraints on Reservoir Size and 

Pressure Drop

The amplitude of the geodetic signal scales with the 

product R2∆p/G, where R is the horizontal radius of the ellipsoid, ∆p is the pressure drop, and G is the host-rock shear 

modulus. Best-fit values of this product are listed in table 1. 

Using formulas for ellipsoid volume VC = (2/3)R2h (where h

is ellipsoid height) and elastic volume change of the ellipsoid 

∆VC = 3VC∆p/(4G) (McTigue, 1987; Tait and others, 1989; 

Tiampo and others, 2000), we find that ∆VC = (h/2)[R2∆p/G]. 

For the first time window, excluding sources at 4–7 km and 

10–14 km depth that clearly do not fit the data, estimates of 

volume shrinkage of the magma body are 2.1–3.5×106 m3. 

By comparison, the hot-rock volume (∆Ve) erupted during this time (Schilling and others, this volume, chap. 8; 

and fig. 3) was about 2.7×107 m3—eight to twelve times 

the volume shrinkage of the reservoir. For the second time 

window, ∆VC = ~3–8×106 m3

 (poorly constrained), whereas 

the change in dome volume was ∆Ve= 1.8×107 m3. For the 

entire eruption through late 2005, Lisowski and others (this 

volume, chap. 15) and Poland and Lu (this volume, chap. 

18) estimate ∆VC ~1.6–2.7×107 m3 from geodetic and InSAR 

data, whereas hot-rock erupted volume by mid-December 

2005 was about 7.3×107 m3 (Schilling and others, this volume, chap. 8). Differences in density of the erupted versus 

unerupted magma (estimated later) are not great enough to 

account for these discrepancies.

The reservoir volume and pressure drop can be constrained if the shear modulus G can be estimated. On the basis 

of estimated seismic P-wave velocity P=6.7±0.2 km/s at 

8–15 km depth (Musumeci and others, 2002, fig. 5), host-rock 

density R =2,700±200 kg/m3 (Williams and others, 1987), 

an assumed Poisson’s ratio ν of 0.25±0.03, and the formula 

G=RP

2

(1–2)/(2(1–)) (for example, Rubin, 1990), we 

obtain G=40±4 GPa. Using this value and ∆VC =2.3×107 m3 

in the formula ∆VC =3VC ∆ p/(4G), we obtain VC ∆ p=1.2×1018

Pa·m3. If we further assume that the pressure drop since the 

start of the eruption is less than a few tens of megapascals, the 

reservoir volume would have to exceed about 41 km3

—significantly larger than previously hypothesized at Mount St. Helens 

(>10 km3 by Scandone and Malone, 1985; 5–7 km3 by Barker 

and Malone, 1991). Several factors considered later and in 

appendix 1 suggest that a lower value of G, and hence smaller 

reservoir volume, is more appropriate.

An Idealized Magma Reservoir

We idealize the magma system (fig. 5) as an ellipsoidal 

magma body several kilometers deep within linearly elastic 

host rock, connected to the surface and to a source of recharge 

through relatively narrow conduits. The reservoir contains 

magma of density  and has a total mass  VC. As long as 

the assumption of linear elastic host rock holds and the geometry of the reservoir does not change, the relation between 

A series of digital elevation models (DEMs) based on 

1:12,000-scale aerial photographs records the growth of 

the lava dome (Schilling and others, this volume, chap. 8). 

These DEMs show that in early October 2004 the rate of 

uplift of cold rock and glacial ice was on the order of 8–9 

m3

/s, with nearly 1.1×107 m3 uplifted by the time lava first 

became visible at the surface on October 11. Lava (“hot 

rock”) extrusion rates were initially ~6–7 m3/s in November 2004 but declined to less than 2 m3

/s in March 2005. 

Since March 2005 (about 150 days into the eruption) the 

extrusion rate has continued to decline gradually, such that 

the volume-time curve can be nearly fit with a straight line 

(fig. 3). Overall, the history of lava-dome volume versus 

time (fig. 3) has defined a remarkably regular monotonic, 

concave-downward curve that was noted in early 2005 and 

could be fit using simple exponential (fig. 3C) and logarithmic (fig. 3A) curves. In this study we endeavor to find the 

physical basis for these curve forms.

Geodetic Source Models

Using data from eight continuous GPS stations (fig. 1), 

we estimate the depth, location, and size of the source of 

deflation (the magma reservoir) by comparing inward displacements with those predicted for a vertical prolate ellipsoid embedded in a homogeneous, isotropic, linearly elastic 

half-space, using the equations of Bonaccorso and Davis 

(1999) (table 1). We evaluate the displacements during two 

time windows: (1) November 4, 2004, to February 5, 2005, 

and (2) February 5, 2005, to July 14, 2005. The start and end 

dates of these windows were chosen to coincide with DEM 

acquisition dates. We use only two time windows because 

data quality is insufficient to allow subdivision into shorter 

time windows. The second time window ends in summer 

2005 because displacements since that time have been too 

small to be accurately modeled.

Measured displacements were adjusted for regional 

plate movement using the rates 3.461 mm/yr east, 5.91 mm/

yr north, 1.46 mm/yr down; and for seasonal changes using 

sinusoidal adjustments having east, north, and z (vertical) 

amplitudes of 1.9149, 1.667, and 1.5289 mm and phase 

angles relative to January 1, 2004, of 51.86, 132.89, and 

32.62 degrees, respectively (Lisowski and others, this volume, chap. 15). For the first time window, best-fit models 

place the top of the reservoir at 3–6 km below the mean 

altitude of the geodetic stations (which is ~1,300 m above 

sea level). In plan view the best-fit models lie 1.3–1.6 km 

east and 5–320 m south of the crater center (figs. 4, 5; table 

1). Placing the top deeper than about 6 km tends to underestimate the radial displacement at stations proximal to the 

crater and overestimate both radial and vertical displacement 

components in the distal stations. The depth to the bottom 

of the reservoir is not well constrained but likely lies somewhere below 10 km (fig. 4A). Data from the second time 

period provide significantly poorer constraints (fig. 4B).
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 465

0 100 200 300 400 500

Ve = 3.57x107ln(1+0.0151t )

Log growth without recharge Log growth with recharge

Exponential growth without recharge Exponential growth with recharge

DAYS SINCE OCTOBER 1, 2004 DAYS SINCE OCTOBER 1, 2004

Total volume

Hot-rock volume

Prediction on 12/11/04

2/1/05

3/10/05

6/15/05

12/15/05

0 100 200 300 400 500

0 100 200 300 400 500

DAYS SINCE OCTOBER 1, 2004 DAYS SINCE OCTOBER 1, 2004

0 100 200 300 400 500

2004 2005

a = 1.8x107 m3

b = 8.5x104 m3/d

d = 7.7x105

 m3/d

0

20

40

60

80

100

0

20

40

60

80

100

0

20

40

60

80

100

0

20

40

60

80

100

2004 2005

VOLUME, IN MILLION CUBIC METERS VOLUME, IN MILLION CUBIC METERS

VOLUME, IN MILLION CUBIC METERS VOLUME, IN MILLION CUBIC METERS

2004 2005

a

Ch. 22, Mastin, FIGURE 3

A B

C D

Ve = a ln[1–(d/b)(1–e(bt /a) V )] e = a ln(1+bt )

Ve = a (1–e–bt )

Ve = 8.29x107(1–e–0.0044t )

Ve = a (1–e–bt )+ct

Ve = 2.77x10 7(1–e–0.0178t )+1.09x105t

2006

2006 2006

Model Time period Depth to 

top (m)

Depth to 

bottom (m)

Distance 

east of crater 

center (m)

Distance 

north of crater 

center (m)

Scale factor 

pR 2

/G

VCp

Pa•m3

×1017

VC

106 m3

1 11/4/04–2/5/05 3,419 20,000 1,541 -68 -135 -2.18 2.83

2 11/4/04–2/5/05 3,654 15,000 1,511 -85 -155 -1.72 3.69

3 11/4/04–2/5/05 4,294 10,000 1,448 -100 -236 -1.32 4.68

4 11/4/04–2/5/05 4,000 7,000 1,492 -5 -286 -0.84 1.80

5 11/4/04–2/5/05 6,000 10,000 1,390 -321 -459 -1.79 3.85

6 11/4/04–2/5/05 10,000 14,500 1,600 -500 -885 -3.89 8.34

7 2/5/05–7/14/05 3,500 15,500 1,400 -100 -113 -1.32 2.83

8 2/5/05–7/14/05 7,191 20,000 100 -3000 -286 -3.58 7.67

Figure 3. Lava dome 

volume (red dots) 

and total surfacedeformation volume 

(green dots) versus time 

in days since October 1, 

2004. Error bars are ±5 

percent of the volume. 

Lines in each plot are 

fitted to a subset of data 

as of a certain date 

and then extrapolated 

onward as predictions. 

(Line of 12/11/04 is mostly 

extrapolation whereas 

line of 12/15/05 is mostly 

fitted.) Solid blue line fits 

volume measurements 

until about December 11, 

2004; dashed blue line, 

February 1, 2005; dotted 

blue line, March 10, 

2005; dot-dash blue line, 

June 15, 2005; solid red 

line, December 15, 2005. 

Plots A, B, C, and D show 

these best-fit curves 

using equations of 

different forms, shown in 

upper left corner of each 

plot. Terms a, b, c, and 

d in these equations are 

fitting coefficients; Ve

 is 

total erupted volume, and 

t is time since October 

1, 2004.

Table 1. Parameters in geodetic source models plotted in figure 4.

[Parameters in plain type are specified by the user; those italicized are obtained by optimizing the fit between the model and the data.]
466 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

-10

-8

-6

-4

-2

0

2

4

6

8

TWIW

P696

P698

TSTU

JRO1

P687

P702

RADIAL DISPLACEMENT, IN MILLIMETERS

VERTICAL DISPLACEMENT, IN MILLIMETERS

RADIAL DISPLACEMENT, IN MILLIMETERS

VERTICAL DISPLACEMENT, IN MILLIMETERS

Geodetic data

3.4–20 km depth

3.6–15 km

4.2–10 km

4.0–7.0 km

6.0–10.0 km

10–14 km

-20

-15

-10

-5

0

5

10

-20

-15

-10

-5

0

5

10

TGAU

TGAU

TWIW P698

TSTU

JRO1

P687

P702

DISTANCE FROM SOURCE, IN KILOMETERS DISTANCE FROM SOURCE, IN KILOMETERS

DISTANCE FROM SOURCE, IN KILOMETERS DISTANCE FROM SOURCE, IN KILOMETERS

-10

-5

0

5

10

TWIW

P696

P698

TSTU JRO1

P687

P702

3.5–15.5 km

7.2–20 km

0 5 10 15 20 25 0 5 10 15 20 25

0 5 10 15 20 25 0 5 10 15 20 25

TWIW

P696

P698 TSTU

JRO1

P687 P702

3.5–15.5 km

7.2–20 km

A

B

November 4, 2004, to February 5, 2005

February 5, 2005, to July 14, 2005

Geodetic data Geodetic data

TGAU

P696

TGAU

Figure 4 Measured and theoretical geodetic displacements for 252 days of eruption from November 4, 2004, to July 

14, 2005. Data points, from continuous global positioning system (CGPS) receivers (fig 1A), show radial and vertical 

displacement; error bars are given as ±5 mm for radial displacement, ±10 mm for vertical displacement. Displacements 

are positive for outward radial and upward vertical directions. Lines represent theoretical displacements resulting from 

deflation of a vertical prolate ellipsoid in an elastic half-space. Depths in the explanation indicate, respectively, the 

depth to top and bottom of the ellipsoid below the mean altitude of the geodetic stations (1,300 m above sea level). A, 

Radial and vertical displacement between November 4, 2004, and February 5, 2005. B, Radial and vertical displacement 

between February 5 and July 14, 2005.
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 467

2

0

-5

-10

-15

-20

-20

-18

-16

-14

-12

-10

-8

-6

-4

-2

0

N S

Mean altitude of:

geodetic stations (top)

seismic stations (bottom)

Qi

Frictional

plug

H1

H

H2

3456 7

0 20 40 60

SHEAR MODULUS G,

IN GIGAPASCALS

G

Model

1

2

3

4

5

6

7

8

No vertical exaggeration

H3

ALTITUDE (SEA LEVEL DATUM), IN KILOMETERS

ALTITUDE, IN KILOMETERS, ABOVE MEAN ALTITUDE OF SEISMIC STATIONS

Qo

0 1 km

υP

SEISMIC VELOCITY υP,

IN KILOMETERS PER SECOND

Figure 5. Left, Idealized north-south cross section through Mount St. Helens showing approximate location 

of the magma body (shaded), as judged from the geodetic modeling and some of the parameters used in this 

analysis. Blue and black dashed ellipses represent locations of geodetic models 1 through 8 listed in table 1; 

width of each ellipse represents its radius taken from value of R2

Δp/G in table 1, using Δp=10 MPa and G= 35 

GPa. Right, Profile of seismic velocity (from Musumeci and others, 2002) and shear modulus G (estimated 

from seismic data and from density data of Williams and others, 1987). Shear modulus was estimated from the 

formula G= RP

2

(1–2)/(2(1–)), where Poisson’s ratio, , is taken as 0.25 and  is taken as 2,150 kg/m3 above 

the seismic datum and 2,700 kg/m3

 below it (Williams and others, 1987, fig. 8). Cross section and seismic 

profile have same vertical scales, but latter is set to datum of mean altitude of seismic stations.
468 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

reservoir pressure p and volume VC is linear with a proportionality given by the reservoir compressibility C:

V

V

j (1)

For a sphere or prolate ellipsoid, C =3/(4G) (McTigue, 1987; 

Tiampo and others, 2000). We also assume that the magma has 

a finite compressibility (M) given by:

1 M

M

M p





∂ ≡ ∂ . (2)

Finally, we assume that the change in reservoir mass 

(dMC) equals the mass added by recharge (dMi=idVi) minus 

the mass erupted (dMe ≈edVe), where i, e, dVi, and dVe are 

the densities and volumes of injected and erupted magma, 

respectively. In mathematical terms,

( ) M C M C C M

e e i i

d V dV V d

dV dV

= +

= − +



 

  . (3)

Adding terms for C and M into equation 3 and rearranging 

leads to:

1

i i e e

C

M

M

C

dV dV dV  







− =     +

 

. (4)

In the absence of recharge we have:

1 e M M

C e C

dV

dV 





  

= − +    

. (5)

Equation 5 carries the important implication that the 

erupted volume should not equal the volume shrinkage in the 

magma body except in the limiting case where the densities 

of erupted and unerupted magma are equal and the magma 

is incompressible (M→0) (this was also pointed out by 

Johnson and others, 2000). If compressibility and density do 

not vary greatly with time during an eruption, equation 5 can 

be used to give the ratio of erupted volume Ve to the volume 

change of the magma reservoir ∆VC. In spherical or ellipsoidal reservoirs, magma compressibility is generally thought 

to greatly exceed the reservoir compressibility (for example, 

Huppert and Woods, 2002); hence erupted volume should 

greatly exceed ∆VC.

In the absence of recharge, what value of dVe / dVC might 

one expect at Mount St. Helens? The answer requires careful 

estimation of M, e, M, and C, which we provide in the following several paragraphs.

The density of unerupted, volatile-saturated rhyolitic melt 

at ~200–250 MPa pressure is about 2,200 kg/m3 (estimated 

using the method of Ghiorso and Sack in the program Conflow 

of Mastin, 2002). Combined with roughly 45 volume percent 

plagioclase crystals (Pallister and others, this volume, chap. 

30) having a density of 2,600 kg/m3

, the bulk density of the 

magma M would be about 2,380 kg/m3. Density measurements 

of most dome rock samples are about 2,300–2,500 kg/m3 (K. 

Russell, written commun., 2006), although pores and voids 

could reduce the bulk density of the dome, e, by perhaps 

10–20 percent below that of the dome rock. Within the uncertainties, we estimate the ratio M /e to be about 1.0 to 1.2. We 

use a reservoir compressibility of roughly 2×10–11 Pa–1 based 

on the formula C =3/(4G) for an ellipsoidal reservoir and the 

earlier estimate of G=40±4 GPa.

Magma Compressibility

Magma compressibility M depends on gas volume 

fraction, solubility, crystallinity, and rate of loading. When 

changes in pressure are much more rapid than rates of gas 

exsolution (for example, the time scale of seismic-wave 

disturbances), the crystal, melt, and gas phases can be 

regarded as inert, and the bulk compressibility is simply the 

sum of the compressibilities of the crystal, melt, and gas 

phases (x, m, g) multiplied by their respective volume fractions (vx

, vm, vg) (for example, Mastin, 2002):

M m m x x g g  = + + v v v    . (6)

If, on the other hand, pressure changes occur over months 

or years, as in the current eruption, gas exsolution must be 

considered. Previous investigators (Tait and others, 1989; 

Huppert and Woods, 2002; Woods and Huppert, 2003) used a 

simple Henry’s solubility law for H2O and found an abrupt discontinuity in compressibility at the saturation pressure (~240 

MPa in fig. 6C). Huppert and Woods (2002) and Woods and 

Huppert (2003) suggested that this discontinuity could have a 

dramatic, rejuvenating effect on the course of an effusive eruption once the magma reservoir reaches the saturation pressure.

The Mount St. Helens magma contains both H2

O and 

CO2, and gas in such a two-component system should exsolve 

more gradually and over a wider range of pressures than it 

would if only H2O were present. We estimate exsolved volatile 

content and magma compressibility using petrologic constraints from other studies. Phase equilibrium experiments 

(Rutherford and Devine, this volume, chap. 31) suggest that 

the currently erupting magma last equilibrated at a temperature 

of ~850°C, a pressure of ~120 MPa, and a source depth near 

5 km. The crystallinity at this depth was 40 to 55 percent (Pallister and others, this volume, chap. 30). The center of deflation, however, is substantially deeper than 5 km, perhaps equal 

to that of the May 18, 1980, magma at around 8–9 km depth 

and 220 MPa pressure (Rutherford and Devine, 1988). Following Gerlach and others (this volume, chap. 26), we assume that 

present-day magma properties at the source resemble those 

in 1980, with a temperature of about 900°C, pressure of ~220 

MPa, 30 percent crystals (Cashman and Taggart, 1983), and 
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 469

0

0.02

0.04

0.06

GAS VOLUME FRACTION

wt% H2O = 1.3x10-6 p0.559 (p in MPa) 30% crystals

T = 900°C

6.29 wt % H2O

1,900 ppm CO2, 5.2% H2O

1,000 ppm CO2, 5.3% H2O

2,100

2,200

2,300

2,400

DENSITY, IN KILOGRAMS

PER CUBIC METER

150 200 250 300 350 400

150 200 250 300 350 400

150 200 250 300 350 400

10-10

10-9

PRESSURE, IN MEGAPASCALS

COMPRESSIBILITY, IN RECIPROCAL PASCALS

Without exsolution

With H2O exsolution

Without gas

A

B

C H2O saturation

pressure

1,900 ppm CO2, 5.2% H2O

1,000 ppm CO2, 5.3% H2O

Figure 6. Characteristic features of magma having composition given in text, as function of pressure. 

Significance of the various lines is explained in text. A, Gas fraction by volume. B, Density. C, 

Compressibility.

a dissolved water concentration in the melt of about 5 weight 

percent (Blundy and Cashman, 2001).

Assuming that the CO2 emitted into the atmosphere during this eruption originated from a mass of magma equal to 

that of the lava dome, Gerlach and others (this volume, chap. 

26) estimate a preeruptive CO2 concentration in the magma of 

about 1,100 ppm, or 1,900 ppm normalized to the melt alone 

(assuming 30 percent crystallinity). At 220 MPa pressure and 

900°C temperature, with 5 weight percent dissolved H2O, 

such a melt would contain roughly 350 ppm dissolved CO2

; 

the remaining CO2 and H2O would be exsolved in bubbles 

composing roughly 1.2 volume percent of the magma (Gerlach 

and others, this volume, chap. 26). The total water content 

normalized to the melt plus fluid phases would be about 5.2 

weight percent. 

Using the solubility code VolatileCalc 1.1 (Newman and 

Lowenstern, 2002), we calculate mass fractions of exsolved 

H2

O and CO2 over pressures ranging from 150 to 400 MPa, 

and we combine these results with ideal gas relations to obtain 

the volume fraction of exsolved gas (fig. 6A). We assume a melt 

density m =2,200 kg/m3 and crystal density x=2,600 kg/m3; 

the melt density corresponds to a water-saturated melt of 1980 
470 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

composition (sample SH-084 of Rutherford and others, 1985), 

calculated by the method of Ghiorso and Sack (1995) using 

the program Conflow (Mastin, 2002). From these volume fractions and phase densities we calculate the bulk density over the 

pressure range 150–400 MPa (fig. 6B) and then numerically 

calculate (1/M)( ∂M /∂p) to obtain M (fig. 6C). We use a melt 

compressibility of 2.0×10–10 Pa–1, estimated using the method 

of Ghiorso and Sack (1995) for water-saturated rhyolite at 

p=220 MPa, and a crystal compressibility of 2×10–11 Pa–1

obtained for albite at T=900°C and p=220 MPa from the program MELTS using the method of Elkins and Grove (1990).

Our calculations suggest that, over the pressure range of 

150–400 MPa, the magma may have a bubble volume fraction of 0.0008 to 0.048 (fig. 6A). For comparison, we plot 

results for CO2=1,000 ppm and H2O=5.3 weight percent of 

the melt+fluid phases and also for a single-component (H2O) 

volatile system using the Henry’s law solubility illustrated 

in figure 6A, with the H2

O content (6.29 weight percent) set 

so that the volume fraction of gas at 220 MPa equals that of 

the two-component system. Both the two-component and the 

Henry’s solubility laws show nearly an order-of-magnitude 

variation in compressibility over this pressure range, but the 

two-component systems show little or no discontinuity in M

at the saturation pressure. At p=220 MPa, the two-component 

system (1,900 ppm CO2, 5.2 weight percent H2O) gives 

M =2.8×10–10 Pa–1. We use this number as a starting point in 

our calculations. The relation between M and volume fraction of gas for these two-component magma compositions is 

illustrated in figure 7.

Expected Ratio of Erupted Volume to Volume 

Change in the Reservoir

Using M /e=1, C =2×10–11 Pa–1, and M =2.8×10–10 Pa–1

(~1.2 percent bubbles), then dVe / dVC predicted by equation 5 

should be about 16, which is somewhat higher than the ranges 

of 7.7–12.8 and 2–6 calculated for the first and second time 

windows, respectively. A temporal increase in M is suggested 

by the difference in apparent values of dVe / dVC between the 

first and second time windows, though uncertainties in these 

ratios are too great to make such an inference with confidence. 

For the overall eruption, the erupted volume by mid-December 

2005, 7.3×107

 m3, was only about three times the reasonably 

well-constrained reservoir shrinkage dVC of ~2.3×107 m3 estimated by Lisowski and others (this volume, chap. 15). A ratio 

this low suggests a magma compressibility of ~4×10–11 Pa–1, 

which is about one-fourth that of even a bubble-free magma 

of this crystallinity (fig. 7). If the magma reservoir is in fact 

bubble free, C must be ~8×10–11 Pa–1 in order to have dVe

 / 

dVC≈3; if it contains roughly 1 volume percent bubbles at the 

source depth, as inferred by Gerlach and others (this volume, 

chap. 26), C must be ~1.4×10–10 Pa–1. In either case, M is 

only a few to several times greater than C, which we take to 

indicate that the volume fraction of bubbles in the reservoir is 

zero or very small, consistent with the findings of Gerlach and 

10-4 10-3 10-2 10-1

10-10

10-9

COMPRESSIBILITY, IN RECIPROCAL PASCALS

GAS VOLUME FRACTION

1,900 ppm CO2, 5.2% H2O

1,000 ppm CO2, 5.3% H2O

30% crystals T = 900°C

others (this volume, chap. 26). If the reservoir was partially 

recharged during the eruption (a factor we consider later), a 

still lower ratio of M /C would be implied, suggesting still 

lower bubble content in the reservoir.

Idealized Models for History of 

Extrusion and Deflation

 Previous studies (Wadge, 1981; Stasiuk and others, 1993; 

Huppert and Woods, 2002) point out that drainage of an elastic 

magma reservoir through a Newtonian conduit results in an 

exponentially decreasing extrusion rate. When combined with 

constant recharge rate into the magma system, the resulting 

curve has the form of an exponentially decreasing rate superimposed on a constant rate (fig. 3D). If these processes are 

responsible for determining the dome-growth curve at Mount 

St. Helens, then some information on the magma reservoir and 

conduit properties should be embedded in the coefficients to 

these equations. Information on the magma-reservoir volume 

and driving pressure are of particular importance.

We know that processes besides Newtonian flow and 

elastic relaxation may affect the growth history. The highly 

crystalline magma at Mount St. Helens, for example, probably has a non-Newtonian rheology. The extrusion of lava as 

a nearly solid plug bounded by frictional faults may also limit 

the growth rate. Different constitutive laws may predict difFigure 7. Compressibility versus volume fraction gas for a rhyolitic 

melt containing 30 percent crystals at 900°C temperature, using 

two different gas compositions.
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 471

ferent growth curves for the lava dome. Some curves fit to the 

hot-rock data (fig. 3) that have simple analytical forms are: 

logarithmic (fig. 3A) V a b t c e = + − ln 1( ) ( ) , (7)

exponential (fig. 3C) ( ) ( ) 1 e b t c V a e

− − = − , (8)

modified log (fig. 3B)

( )

ln 1 1

b t c

a e

d V a e

b

−    = − −        

, (9)

and exponential plus linear (fig. 3D) 

( ) ( ) 1 ( ) b t c V a e d t c e

− − = − + − . (10)

The terms a, b, c, and d are fitting coefficients; their bestfit values are listed in table 2. (Parameter c, which represents 

the day of the eruption start, is used as a variable in these fits, 

even though its value is roughly known, making the number 

of truly unknown parameters equal to three.) Curve forms that 

fit the data best are equations 9 and 10, which, as shown later, 

assume a constant rate of recharge. The fitting errors are lower 

using these forms than using equations 7 and 8 (table 2). More 

importantly, however, curves of equations 9 and 10 have done 

a better job predicting future growth, as the best-fit coefficients for equations 9 and 10 have changed relatively little 

since early 2005 (table 2).

Exponential Growth Curve

Exponential curve forms of equations 8 and 10 are 

derived from two main assumptions. The first is that the 

magma-reservoir pressure p is linearly related to the mass of 

magma in the reservoir:

0 C( ) e i p p M M = − − , (11)

where Me is the mass that leaves the reservoir (assumed to 

equal the erupted mass), Mi is the mass that enters the reservoir as recharge, p0

 is initial reservoir pressure, and C is a 

constant that represents the change in pressure with reservoir 

mass, ∂p/∂MC. By substituting equations 2 and 1 into equation 

3, and rearranging, we find that C=[(C+M)mVC]

–1.

The second assumption is that the mass effusion rate Me

is linearly related to magma reservoir pressure (p):

M p e = − A B, (12)

where A and B are constants. This equation describes, among 

other possibilities, Newtonian (Poiseuille) flow (fig. 8A); 

Newtonian flow capped by a frictional plug (fig. 8C); flow of 

a solid mass through the conduit separated from the conduit 

walls by a Newtonian fluid (a “greased plug”; fig. 8B); and 

Date

Days 

since 

10/1/04

Hot-rock 

volume 

m3×106

Rate 

m3/s

Exponential best-fit parameters Logarithmic best-fit parameters

With recharge Without recharge With recharge Without recharge

a b c d Σ(y -y)

i

2 a b c Σ(y -y)

i

2 a b c d Σ(y -y)

i

2 a b c Σ(y -y)

i

2

m3×106 s–1×10–7 s×105 m3/s

m3×1011 m3×106 s–1×10–7 s×105 m3×1011 m3×106 m3/s s×105 m3/s

m3×1011 m3×106 s–1×10–7 s×105 m3×1011

10/1/04 0 010/13/04 12 5.411/4/04 34 11.8 5.9211/29/04 59 21.3 4.412/11/04 71 25.5 4.05 14.0 2.94 9.56 2.79 0.31 53.8 1.23 9.49 0.75 8.0 3.81 9.70 8.28 0.21 39.5 1.74 9.55 0.4

1/3/05 94 30.5 2.52 86.7 1.01 9.61 -1.98 1.76 46.7 1.49 9.66 2.27 29.7 1.47×10-7 9.71 7.55 2.42 29.5 2.58 9.75 4.8

2/1/05 123 35.1 1.84 94.3 0.96 9.59 -2.26 1.74 45.0 1.58 9.74 2.94 24.2 2.20×10-7 9.88 8.34 3.31 24.5 3.40 9.94 13.52/21/05 143 39.2 2.37 28.1 2.19 9.76 1.15 10.16 48.2 1.41 9.49 13.24 24.6 1.40×10-7 10.00 8.36 11.77 25.3 3.24 9.90 14.43/10/05 160 41.9 1.84 23.6 2.55 9.81 1.48 11.81 50.7 1.29 9.22 24.36 25.8 1.82×10-7 10.04 8.11 15.17 25.9 3.12 9.86 15.54/19/05 200 47.5 1.62 22.9 2.63 9.82 1.52 11.87 55.4 1.10 8.43 57.96 22.2 0.52 9.87 8.28 17.58 26.9 2.92 9.74 18.7

6/15/05 257 53.9 1.3 25.7 2.32 9.72 1.35 14.62 61.1 0.92 7.16 107.49 21.9 0.56 9.87 8.31 17.62 27.8 2.75 9.61 21.8

7/14/05 286 57.1 1.28 26.4 2.25 9.67 1.31 15.41 64.5 0.83 6.31 143.93 21.3 0.63 9.88 8.38 17.93 28.6 2.61 9.47 21.8

8/10/05 313 61.7 1.97 24.9 2.42 9.79 1.39 23.82 69.3 0.73 4.96 253.01 18.1 0.96 9.98 8.92 39.02 30.8 2.29 9.02 77.1

10/24/05 388 70.0 1.28 25.0 2.41 9.77 1.38 23.86 78.3 0.57 1.65 485.74 16.2 1.13 10.06 9.44 51.27 34.4 1.87 7.98 179.4

12/15/05 440 73.0 0.67 27.7 2.06 9.40 1.26 71.01 82.9 0.51 -0.16 551.65 18.0 0.98 9.96 8.89 61.32 35.7 1.75 7.54 195.1

Table 2. Fitting coefficients to curves in figure 3.[Columns labeled 

Σ(y –i y)

2 give the sum of the squares of errors between best-fit predictions and data.]
472 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Newtonian

flow

“Greased

plug” flow

Without frictional cap With frictional cap

∆r

R

σn τ

H1

H

H2

C

A B

D

R

H

velocity

profile

greased-plug flow capped by a frictional plug (fig. 8D). We 

also assume that the linear relations in equations 11 and 12 do 

not change with time.

If one further assumes that the rate of mass recharge 

( ) M Q i i = is constant, equations 11 and 12 can be combined 

and integrated (appendix 2) to give the erupted mass as a function of time. Noting that the erupted volume (Ve

) is equal to 

Me /e, we obtain:

( ) ( ) 1 A B 0 1

AC

i ACt i

e

e e

p Q Q V e t

 

− − − = − + . (13)

This equation has the same form as equation 10 with the following coefficients:

( ) A B 0

AC

i

e

p Q

a



− − = , (14)

b = AC, and (15)

i

e

Q d

 = . (16)

Differentiating equation 13 with respect to time, we 

find that this curve has an initial volumetric extrusion rate t 0 Ve

= = ab+d = (Ap0

–B)/e but asymptotically approaches a 

linear trend having the slope t Ve

→∞ = d = Qi

/e. Best-fit values 

of a, b, and d for the growth curve (table 2) suggest that t 0 Ve

= & = ~7.0 m3

/s (605,000 m3/day) and t Ve

→∞ = ~1.26 m3

/s 

(109,000 m3/day). The latter value (the recharge rate) is 

several times greater than the long-term magma-supply rate 

of 0.2 m3/s at Mount St. Helens, estimated by assuming that 

most of the volume of the edifice (~25 km3) was erupted in 

the past 4,000 years (Iverson and others, 2006). Finally, the 

volume constant a, roughly 2.8×107 m3, is the y-intercept 

of the long-term growth line in fig. 3D (the black dashed 

line). The y-intercept represents the volume of magma that 

has erupted and has not been replaced in the reservoir by 

recharge. The remaining volume, more than 4.5×107

 m3, 

represents recharge. 

The physical significance of terms in a and b can be further refined, depending on the type of flow in the conduit. For 

Poiseuille flow, the mass flow rate is (Mironer, 1979, p. 194):

( )

4

8

e e dM R p gH dt H

 

 = − , (17)

where H is the conduit length,  is the average viscosity, and 

 is the average magma density between the magma reservoir and the Earth’s surface. This equation assumes that the 

pressure at the top of the conduit is negligible (an assumption 

we will evaluate later). For Poiseuille flow, the constants in 

equation 12 are A=qeR4/(8H) and B=qeR4  g/(8). Inserting these expressions into equation 13, we can recast a in the 

following form:

( ) ( ) 0

0 0

t t

M e e

t C M C

e e

V V

a V p gH V   



 = →∞

=

  − =   + −

 

. (18)

Equation (18) can be generalized by noting that ( ) 0 p gH −  is 

the initial pressure at the base of the conduit in excess of that 

required to initiate upward flow. Denoting this term as 0

ex p , we 

can rewrite the equations as:

( ) 0

0 0

t t

M e e ex

t C M C

e e

V V

a V p V

 



= →∞

=

  − =   +

 

& &

& , (19)

( )

0

0

t

e e

ex

M C M C

V b AC

V p



  

=

= = + . (20)

It can be shown (appendix 3) that these equations also apply to 

greased-plug flow (fig. 8B) and to flow of either a Newtonian 

Figure 8. Types of conduit flow that are considered in deriving 

theoretical lava-dome growth curves. See text for full explanation.
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 473

fluid or a greased plug capped by a frictional plug (figs. 8C, 

8D; appendices 4, 5), so long as the plug’s geometry and coefficient of friction do not change with time. If a frictional cap is 

present, 0

ex p represents the pressure in the magma reservoir in 

excess of both the pressure due to the weight of the overlying 

magma ( ) gH and the pressure required to overcome frictional resistance of the plug.

Some important relations fall out of the equations above. 

In the absence of recharge, the final erupted volume is:

( ) 0

M ex

C M C

e

a V p   



= + . (21)

Using C =8×10–11 Pa–1, M /e≈1, VC =10 km3, 0

ex p =~30 MPa, 

and M =1.6×10–10 Pa–1 appropriate for a bubble-free magma, 

the volume a is 7.2×107 m3—roughly equal to the 7.3×107

m3 that has erupted by the end of 2005. The time required 

for extrusion of 90 percent of this volume is about 3.5/b, or 

~416 days. Using a magma compressibility consistent with 1 

percent bubbles (M =2.8×10–10 Pa–1) and C =1.4×10–10 Pa–1 (to 

maintain dVe / dVC≈3 following equation 5), we obtain a theoretical extrusion volume a=1.26×108

 m3 and duration of 2.0 

years. These results suggest that the eruption could continue 

beyond early 2006 without recharge, but only if the exsolved 

volatile fraction in the reservoir is significantly greater than ~1 

percent or if the reservoir is significantly larger than ~10 km3, 

or both.

For the case of recharge, taking t Ve

→∞ =1.26 m3

/s and 

a=2.8×107 m3 (table 2), the product 0 ( ) ex V p C M C   + obtained 

from equation 19 is about 3.4×107 Pa·m3. Constraining the 

reservoir volume and initial overpressure requires some additional constraint on magma compressibility, which is considered later.

Friction, Faulting, and the Growth Curve

A log-based formula having the form of equations 7 or 

9 would be predicted if the growth of the lava dome were 

controlled by frictional resistance of a solid mass in the upper 

conduit, with the coefficient of friction increasing with the rate 

of slip.

Although the coefficient of friction is commonly taken as 

a constant with a value of ~0.6–1.0 (Byerlee, 1978), it actually 

varies slightly with sliding rate and with time between sliding 

events (for example, Scholz, 1998). When  increases with 

displacement rate, acceleration is dampened out and stable sliding (or fault creep) results. When  decreases with displacement 

rate, sliding can accelerate unstably, leading to earthquakes. In 

general, rate-strengthening friction is favored when the shearzone temperature is near the brittle-ductile transition (Chester, 

1994), when a thick gouge layer is present (Byerlee and Summers, 1976), and in near-surface conditions when normal stress 

on the fault plane is low (Marone and Scholz, 1988). These 

conditions all exist in the shallow conduit at Mount St. Helens.

On the other hand, experimental studies of the Mount St. 

Helens fault gouge at 25°C (Moore and others, this volume, 

chap. 20) suggest rate-weakening behavior when displacement 

rates are less than about 1×10–4 m/s and rate-strengthening 

behavior at rates above 5×10–4 m/s. Assuming a 100-m-diameter conduit near the surface, the range of observed volumetric 

extrusion rates (~1–7 m3/s) translates into displacement rates 

of 1–9×10–4 m/s, crossing over the transition between these 

behavior types. We consider it likely that both rate-strengthening and rate-weakening sliding exist at shallow depth at any 

given time. Fault patches of rate-weakening gouge will slip 

abruptly to create small drumbeat earthquakes whereas other 

parts of the fault surface creep stably under rate-strengthening 

conditions. If rate-weakening behavior controls conduit flow, 

the appropriate friction coefficient to use in this model would 

be a value averaged over many stick-slip cycles. If that average 

doesn’t change with time, the long-term growth curve will be 

exponential. But if rate-strengthening behavior controls conduit flow, we need to consider the stress-strain rate relations of 

rate-strengthening fault creep.

The Logarithmic Curve

When the coefficient of friction is rate-dependent, the 

shear stress that resists slip on a fault plane is commonly 

expressed as (for example, Scholz, 1998):

1 ln o n

o

A 



  = +  . (22)

Here n is normal stress at the wall and A1 is the rate dependence of fault strength. The constant o is an arbitrary reference, the strength of the wall interface when the slip rate is o .

In order to derive a growth curve, we assume that the 

frictional plug of mass Mp and length H2 occupies a cylindrical conduit of radius R (fig. 8C). The frictional force along the 

plug margin is 2 2RH n , where n is the mean normal stress 

on the plug margin. The mass flow rate Me is then related 

to o by 2 M R e e =    , and pressure at the plug base (p) is 

related to mass flow rate as:

ln e

o

o

M

p p M= +  , (23)

where 1 2 2 / A H R n   = and 2

0 0 2 p H R Mg R = + 2 / /   are 

constants.

Rearranging this equation yields:

( ) 0 exp M o

e

p p M e

a

  − =    

. (24)

Combining equation 24 with equation 11 (assuming recharge 

Mi=0) and integrating leads to (appendix 6):

0 ln 1 e

e

tM V D

D



 

  = +    

, (25)
474 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

where D =VM (C + M) and VM is the volume of the magma 

reservoir plus conduit. This equation has the same form as 

equation 7, with a V = +    M C M ( ) and 0 / t

e b V a = = .

If recharge into the magma reservoir is included, the 

equation has the form (appendix 6):

0 ln 1 1

i

e

Q t

D

e

i

M V D e

Q

  

   = − −          

. (26)

This equation has the form of equation 10, with a volume 

constant a=D=1.47×107 m3 (by regression through the most 

recent data set, table 2), a recharge rate b=Qi /e=1.0 m3/s, 

and an initial extrusion rate d = 0

0 / t M V e e  = = = 8.9 m3

/s 

(table 3). The recharge rate is about 30 percent less than the 

1.26 m3/s obtained from the exponential curve. By rearranging equation 23, substituting M V e e e =  , 0

0

t M Ve e  = = , and 

a=D=VM(C+M), we can obtain a formula for the product 

of volume of the magma system and pressure drop from the 

beginning of the eruption until the time of the last data point 

used in this paper (December 15, 2005):

( ) ( ) 0 ln / t last V p M C M e e   a V V = ∆ + = . (27)

The parameters p and last Ve are the pressure change at the 

base of the frictional plug and the extrusion rate at the end 

of this time period; the variable a is the numerical value 

of the fitting coefficient. From the first derivative of equation 9, last Ve =1.20 m3

/s as of December 15, 2005 (table 3), 

giving V p M C M ∆ + ( )   =3.0×107 m3. It should be noted 

that this term contains slightly different parameters from 

0 ( ) ex V p C C M   + derived for the exponential curve: VC represents reservoir volume, whereas VM represents volume of the 

reservoir plus conduit below the frictional plug, and p represents pressure drop at the base of the plug, whereas p0

ex gives 

the initial overpressure in the magma reservoir. Nevertheless, 

the values of these terms should be roughly comparable, and 

they are: VMp(C + M)=3.0×107

 m3 from the log fit (equation 9) versus 0 ( ) ex V p C C M  + =3.4×107

 m3 from the exponential fit (equation 10). The fact that these values differ by 

only 10 to 15 percent suggests that inferences about magmareservoir size and overpressure do not depend strongly on the 

assumptions regarding factors that control conduit flow.

Additional Constraints from the Geodetic 

Time Series

On the basis of their fit to the dome-growth data, neither 

the logarithmic (equation 9) nor the exponential (equation 10) 

model can be confidently eliminated. Each, however, predicts 

a history of reservoir deflation that can be compared with 

geodetic data. For the case of exponential dome growth, differentiating equation 13 with time under conditions of constant 

recharge and substituting in equations 4, 14, and 15 gives the 

following for reservoir deflation with time:

( ) 1

1

bt

e

C

M M

C

a e

V 

 



− − ∆ = −     +

 

. (28)

The reservoir deflates with the same time constant as the 

dome-growth curve, implying, for the best-fit value of b with 

recharge through December 15, 2005 (2.06×10–7 s–1), that 90 

percent of the geodetic deflation should have occurred after 

about 200 days, by mid-April 2005, and that by late summer 

2005 the deflation should have essentially stopped. This is 

inconsistent with geodetic data, which show a nearly linear 

rate of inward displacement through at least the end of 2005.

The log curve can theoretically provide a better match to 

the geodetic data, but the predictions at some point become 

physically unrealistic. For the case of zero recharge, for 

VCp0

ex(M+C

) exponential Log with recharge Log without recharge

With recharge

m3

×106

Without recharge

m3×106

t 0= Ve last Ve VMp(M+C

) t 0= Ve

last Ve VMp (M+C

) Date

m3

/s m3/s m3×106 m3/s m3/s m3×106

12/11/2004 24 54 8.3 3.94 5.9 6.9 3.4 28

1/3/2005 67 47 7.5 2.52 33 7.6 2.5 33

2/1/2005 71 45 8.3 1.82 37 8.3 1.8 37

2/21/2005 33 48 8.4 1.64 40 8.2 1.7 40

3/10/2005 29 51 8.1 1.54 43 8.1 1.5 43

4/19/2005 29 55 8.3 1.40 39 7.8 1.3 48

6/15/2005 31 61 8.3 1.20 42 7.6 1.1 54

7/14/2005 32 64 8.4 1.14 42 7.5 1.0 57

8/10/2005 31 69 8.9 1.23 36 7.1 1.0 61

10/24/2005 31 78 9.4 1.24 33 6.4 0.9 68

12/15/2005 34 83 8.9 1.11 38 6.2 0.8 72

Table 3. Calculations of VC p0

ex(M+C

) or VMp(M+C) obtained from exponential or logarithmic best-fit solutions.
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 475

example, combining equations 5 and 25 and substituting 

a V = +    C C M ( ) and 0 / t

e b V a = = , the volume shrinkage of 

the magma system should follow the curve:

ln 1( )

1

e

C

M

M

C

a bt

V 







+

∆ = −     +

 

. (29)

In other words, the volume shrinkage of the magma reservoir 

with time should look like a negative mirror image of the dome 

growth curve, adjusted by the constant e /(M(1+M /C)). This 

curve leads to the physically unrealistic result that deflation 

continues indefinitely, even to negative reservoir volumes, at 

the same time that the lava dome keeps growing. This implication is an outcome of the logarithmic relation between stress 

and displacement rate (equation 22), which adequately fits 

experimental data on rate-dependent friction within the range 

of shear stresses applied during experiments but cannot be 

realistically extrapolated outside that range.

A More Realistic Model

We are therefore left with the result that neither the exponential curve nor the logarithmic curve can adequately fit both 

the lava-dome growth curve and the geodetic deflation history. 

What additional processes might account for the dome growth 

and deflation histories? Some possibilities include:

•	 The effect of the dome’s weight in suppressing further 

extrusion.—Digital elevation models indicate that the 

dome rapidly grew to more than 200 m height in the first 

two months of the eruption, potentially adding several 

megapascals of increased pressure to the vent at the base 

of the dome. Our records on dome-height variations 

with time (fig. 9) can be used to constrain this effect.

•	 Changes in magma compressibility with time.—A 

decrease in reservoir pressure of 30 MPa can increase 

magma compressibility several tens of percent (fig. 

6C), increasing the ability of the magma reservoir to 

maintain a long-term eruption with time.

•	 A nonconstant rate of recharge.—A more realistic 

model would have recharge into the reservoir increasing as reservoir pressure decreases.

These effects require a numerical solution to account for 

changing values with time. We solve the problem using differential equations described below.

To account for the relation between mass eruption rate 

and reservoir pressure, we assume again that the effusion rate 

is linearly related to the reservoir overpressure. For Poiseuille or greased-plug flow, the overpressure pex is simply

p gH −  . If a frictional cap is present, the overpressure is

( ) ex p p gH F = − +  , where F is the strength of the frictional 

cap (assumed constant). The growth of the lava dome changes 

the distance H from the reservoir to the free surface, and in 

order to account for this, we divide this term into two parts; 

0  gH , where H0

 is the distance from the top of the reservoir to 

the vent at the base of the lava dome, and e 3  gH , where H3 is 

the height of the lava dome. The relation between pressure and 

effusion rate is then:

( 0 3 ) 2

0 3

A e e dM p gH gH F

dt H H

 − + +    =   +  

. (30)

The constant A2 has the value eR4/(8) for Poiseuille flow 

and eR3r/(8) for greased-plug flow (appendix 4). This 

Figure 9. Maximum height of the Mount St. Helens lava dome above the 2003 crater floor (the Crater Glacier surface, 

approximately 2,115 m above sea level) during the course of the 2004–6 eruption.

0 50 100 150 200 250 300 350 400 450

-100

-50

0

50

100

150

200

250

300

DAYS SINCE OCTOBER 1, 2004

DOME HEIGHT, IN METERS
476 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

equation can be simplified by noting that, at t=0, pex =p0

ex and 

dMe / dt = 0

/ t Ve e  = :

0 0 3

0

ex

e t e

e ex

dV p p gH V

dt p

= + ∆ −  = . (31)

The initial extrusion rate t 0 Ve

= , the dome height H3

, and the 

density e are constrained from measurements; the pressure 

change in the magma reservoir p is calculated by integration (below); and the initial overpressure 0

ex p is an adjustable parameter whose value is likely less than a few tens of 

megapascals.

The rate of pressure change in the reservoir is obtained 

by differentiating equation 11 with time and substituting 

C=[(C+M)mVC]

–1:

( )

i e

e

M C C M

dM dV

dp dt dt

dt V



  

−

= + . (32)

In this calculation, the terms C and e are considered known; 

dVe / dt is obtained from equation 31, and the recharge rate 

dMi / dt is calculated from a separate differential equation 

(below). The magma compressibility M at the beginning of 

the eruption is an adjustable parameter; during the course of 

the eruption, it increases with decreasing pressure at a rate that 

equals the average slope of the curve of M versus p in figure 

6C (for CO2=1,900 ppm, H2O=5.2 weight percent). In each 

calculation, we use the magma density M shown in figure 6B

at the given compressibility. Thus the magma density changes 

with pressure, though the changes are minor.

In accounting for recharge, we assume that the rate of 

input into the magma reservoir at the onset of the eruption was 

negligible but increased as reservoir pressure was depleted. 

The simplest such relation is linear, using an adjustable proportionality constant Qlt:

0

i

lt ex

dM p Q dt p

∆ = – . (33)

The parameter p is the reservoir-pressure change since the 

start of the eruption (negative p implies a pressure decrease). 

Like Qi in equation 13, Qlt in equation 33 largely controls the 

long-term extrusion rate. The linear assumption implies laminar flow of magma into the reservoir from some deeper source 

whose pressure remains constant. A more realistic model 

would consider a finite source whose pressure decreased 

over time, but we have no constraints on the rate of pressure 

decrease and hence ignore it under the assumption that the 

deeper reservoir is much larger than the shallow one that feeds 

the eruption.

Equations 31, 32, and 33 can be simultaneously integrated 

to yield both a dome-growth curve and deflation history. The 

calculations involve four adjustable constants: VC, 0

ex p , an initial 

value of M, and Qlt. Our solution takes C to be a known quantity, although its value is known only approximately. The above 

estimates of Ve /ΔVC suggest C to be at least a few times greater 

than our initial estimate of 2×10–11 Pa–1, but best-fit values of 

0 ( ) ex V p C C M   + in table 3 constrain ( )   C M + to be less 

than about 5×10–10 Pa–1 for magma-reservoir volumes greater 

than about 5 km3 and initial overpressures exceeding about 10 

MPa. With these constraints, we run the model using two possible values of 7×10–11 and 1.5×10–10 Pa–1 for C. In theory, the 

values of Qlt and M can be uniquely determined, as they are 

the only factors that significantly affect the long-term eruption 

rate and the ratio Ve /VC, respectively. The values of VC and 

p0

ex are interdependent and nonunique, but ranges of possible 

combinations can be identified. 

In order to compare the deflation history to geodetic 

measurements, we convert the history of pressure change 

p into a history of magma-chamber shrinkage VC , using 

VC =(VC /C)p , and then convert VC into a theoretical displacement at JRO1 using one of the geodetic models in table 

1 (fig. 4). For the first time period (fig. 4A), model 5 in table 

1 (source depth 6–10 km) matches the JRO1 radial displacements best and predicts a radial displacement of 5.32 mm for 

a volume loss VC of 3.85×106

 m3, or 1.4×10–6 mm displacement per cubic meter volume loss.

Results

Numerical model results are compared with measurement histories of lava-dome volume and radial displacement 

at JRO1 in figures 10A and 10B, respectively. The solid black 

line in fig. 10A gives model results that match the hot-rock 

data for VC =17 km3, p0

ex =17 MPa, and Qlt /e

 = 1.7 m3/s. The 

dashed black line gives analytical results using the exponential 

curve of equation 10 with a=2.77×107 m3, b=2.06×10–7 s–1

and d=1.26 m3/s. These theoretical curves cannot be easily compared with the JRO1 data because deflation at JRO1 

began around September 23, 2004, 20 days before the first 

lava appeared and three days before the first visible surface 

deformation (Dzurisin and others, this volume, chap. 14). The 

deflation between September 26 and October 11 was probably associated with extrusion of cold rock ahead of the rising 

magma. Deflation before September 26 may have been associated with intrusion at shallow depth or gas escape, neither of 

which can be easily quantified.

In order to simultaneously fit both curves, we add the 

volume of cold rock extruded before October 11 to the cumulative hot-rock volume (green data points, fig. 10A) and use 

September 27 as the start date (a date determined by a best-fit 

exponential curve through these new data). We also start with 

3 mm of deflation at t=0, the approximate amount of deflation 

measured at JRO1 on September 27.

The blue dashed and solid lines in figure 10A represent 

best-fit analytical (using equation 10) and numerical curves 

through the modified dataset, respectively. Best-fit coefficients 

of the analytical curve give a=3.61×107

 m3, b=2.2×10–7 s–1, 

and d=1.28 m3/s. Using C =7×10–11 Pa–1, M =1.2×10–10 Pa–1

(adjusted to match the deflation curve) and equation 28 to 

calculate magma-reservoir deflation with time, the calculated 

displacements (dashed blue line, fig. 10B) roughly match the 

measurements during the first few months of the eruption and 
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 477

Figure 10. Comparison of theoretical and measured lava-dome volume and magma-reservoir deflation with time at 

Mount St. Helens. A, Erupted volume from digital elevation models. Phrase “HR+def before 10/11/04” refers to data 

points that represent hot-rock volume of the dome plus volume of uplifted cold crater-floor material that appears 

before the beginning of lava extrusion on October 11, 2004. Phrase “Analytic, HR+def before 10/11/04” refers to an 

analytical best-fit line through these data. Details are explained in text. B, Radial displacement measured at JRO1 

continuous GPS station (fig 1A). Negative displacements are radially inward. Red vertical dashed lines in figure 10B

refer to the start and end dates of time windows used in geodetic analysis (fig. 4 and table 1).

during the following summer (~250–400 days into the eruption), 

but they do not predict continued deflation that one would infer 

from a best-fit line through the JRO1 data after about 150 days.

A numerical solution (solid blue lines) provides a slightly 

better fit through the data using VC =17 km3, p0

ex=18 MPa, 

Qlt /e=1.7 m3/s, and M ranging from an initial value of 

1.85×10–10 Pa–1 to a final value of 1.98×10–10 Pa–1 (~0.40–0.48 

percent bubbles). In this solution, the total volume of erupted 

magma that has not been replaced by recharge is about 

3.9×107 m3. This amount is 1.3×107 m3 less than predicted 

by the analytical solution (fig. 10A). The lower total recharge 

implies more geodetic deflation; hence the numerical curve 

can be fit to the geodetic data using a slightly higher average M than required by the analytical solution. Over the time 

window of the simulation, the magma-reservoir pressure drops 

by about 11 MPa, so that the excess pressure ( )1 p gH − 

by mid-December 2005 is about 7 MPa, three megapascals 

greater than the pressure e gH3 at the vent, owing to the weight 

of the overlying dome (taking e=2300 kg/m3 and H3=167 m 

on December 15, 2005).

0

20

40

60

80

100

Recharge after 500 days = 38.7×106 m3

Numerical recharge

Analytical recharge

Hot-rock volume

HR+ def. before 10/11/04

Analytic hot-rock

Analytic, HR+ def before 10/11/04

Numerical, VC = 17 km3, p0 = 18.0 MPa

Numerical, VC = 17 km3

, p0 = 17.0 MPa

0 50 100 150 200 250 300 350 400 450 500

0 50 100 150 200 250 300 350 400 450 500

-40

-30

-20

-10

0

10

pex =0

∆p=–11.2 MPa

�C=7.0x10–11 Pa–1

DAYS SINCE OCTOBER 1, 2004

JRO1 RADIAL DISPLACEMENT, IN MILLIMETERS

JRO1, seasonally adjusted

A

B

ERUPTED VOLUME, 

Ve, IN MILLION CUBIC METERS

Analytic, �M=1.2x10–10 Pa–1

�M=1.9x10–10 Pa–1

�M=1.9x10–10 Pa–1

ex

VC = 17 km3, p0 = 17.0 MPa

VC = 17 km3

, p0 = 18.0 MPa

ex

ex

ex
478 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Other combinations of VC, p0

ex, Qlt, and M that produce reasonable fits by numerical solution are listed in table 

4. All combinations require Qlt /e=1.7±0.1 m3/s. Runs that 

use C =7×10–11 Pa–1 require average values of M around 

1.6–1.9×10–10 Pa–1 (0.18–0.43 volume percent bubbles) and a 

magma-reservoir volume exceeding ~9 km3 for p0

ex<~30 MPa. 

Runs that use C =15×10–11 Pa–1 require M =4.0–4.7×10–10

Pa–1 (1.0–1.8 volume percent bubbles) and a magma-reservoir 

volume exceeding ~5 km3 for p0

ex less than about 30 MPa. 

Using a still larger reservoir compressibility (C =3×10–10 Pa–1), 

a magma reservoir larger than 2 km3 (for p0

ex less than about 

30 MPa) can still fit the curves using M =8.9–9.4×10–10 Pa–1

(2.8–3.8 percent bubbles). The results that involve C =7 to 

15×10–11 Pa–1 correspond to exsolved fluid contents in the 

source region that match most closely with gas emission data 

(Gerlach and others, this volume, chap. 26).

The Quandary of Continued Deflation

Like the exponential function, the numerical solution predicts that deflation should have nearly ended several months 

after the eruption began, which does not agree with the geodetic data. Factors that might keep both the extrusion rate and 

the deflation rate more or less constant include (1) decreasing 

magma viscosity or friction coefficient with time; (2) increasing conduit diameter with time; and (or) (3) a nonlinear relation between extrusion rate and friction coefficient, similar to 

the logarithmic relation.

As of March 2006 there have been no obvious temporal 

changes in petrology or fault-gouge characteristics that might 

reflect changes in viscosity or friction coefficient (Pallister and 

others, this volume, chap. 30). Changes in conduit diameter 

cannot, however, be dismissed (our field observations are 

insufficient), nor can the possibility that conduit enlargement 

alone, in the absence of recharge, is responsible for sustained 

extrusion rates. Figure 11B shows a theoretical deflation curve 

calculated in the absence of recharge by solving equation 5 for 

reservoir volume loss dVC using (M /e

)=1, M =3.3×10–10 Pa–1 

(adjusted to optimize fit), C =7×10–11 Pa–1, and the erupted 

volume dVe

 (=Ve) obtained from the best-fit curve, equation 

10, through modified hot-rock data (fig. 11A). The theoretical curve matches the long-term linear trend quite well but 

underestimates the deflation in the first few months of the 

eruption. The rapid early deflation implies that the volume 

removed from the reservoir per unit erupted volume was 

initially high but then decreased with time, a characteristic 

that could be explained by either increasing recharge or by 

increasing magma compressibility with time. Starting with a 

slightly lower compressibility (3.2×10–10 Pa–1) that increases 

to 4.1×10–10 Pa–1 over the course of the eruption (fig. 11B, red 

dashed line) does not appear to reconcile the difference.

The pressure change p is related to the volume shrinkage dVC by p=dVC /(VCC). For VC =15 km3

, for example, 

the deflation in figure 11B represents a pressure drop of about 

15–16 MPa, requiring an initial overpressure above this value 

to sustain the eruption for the observed time period. The 

conduit radius R that could give the instantaneous growth rate 

in figure 11A with the pressure in figure 11B can be calculated from equation 17 by substituting 0 3

ex

e p p gH + ∆ −  for 

( p gH −  ), t 0 Ve

= for ( ) 4

0 0 / 8 ex   R p H , and e e V for dMe /dt:

1/ 4

0

0 0

0 3

ex

e

t ex

e e

V p R R

V p p gH  =

   =      + ∆ − 

. (34)

Here, R0 is the initial conduit radius. Calculating the change 

in R with time requires values of VC and p0

ex to be assumed 

in advance. For VC =15 km3 and p0

ex =30 MPa, the observed 

extrusion history can be produced by a roughly 15-percent 

decrease in conduit radius in the first few months of the eruption, followed by widening at a slow but accelerating rate (fig. 

11C). As the overpressure approaches zero, the conduit radius 

must approach infinity to keep the extrusion rate constant (for 

example, the line for p0

ex =15 MPa in fig. 11C).

These results suggest that the changes in conduit radius 

required to maintain the observed eruption rate without 

recharge are less than a few tens of percent and thus probably 

too small to be easily detected by observed variations in lithic 

content of the lava or ratios in linear to volumetric extrusion 

rate. The changes shown in figure 11, however, involve a 

seemingly unrealistic shrinkage in radius early in the eruption, 

when effusion rates are high, followed by enlargement at an 

accelerating rate when effusion rates are low. We cannot dismiss conduit-radius changes in the absence of recharge, but we 

C

Pa-1×10-11

VC

km3

p0

ex

MPa

avg M

Pa-1×10-10

Qlt/e

m3/s

p

MPa

Recharge

m3×106

7 6 41 1.6 1.7 -30.5 48.7

7 8 32 1.6 1.7 -23 46.7

7 10 27 1.8 1.7 -18.7 44.4

7 12 24 1.8 1.7 -16.3 42.9

7 14 21 1.9 1.7 -13.7 40.9

7 17 19 1.9 1.7 -12.0 39.0

7 20 16.5 1.9 1.7 -9.9 37.0

7 24 15 1.9 1.7 -8.6 34.9

7 27 14 1.9 1.7 -7.8 33.4

14 4 34 3.0 1.7 24.7 47.1

14 6 25 3.1 1.7 17.1 43.6

14 8 19 4.0 1.7 -12.0 39.2

14 10 16 4.1 1.7 -9.4 36.4

14 12 14.5 4.2 1.7 -8.2 34.2

14 14 13.3 4.2 1.7 -7.1 32.2

14 17 12 4.3 1.7 -6.0 29.6

14 20 11 4.7 1.7 -5.0 26.7

14 24 10.2 4.7 1.7 -4.3 24.3

Table 4. Combinations of VC , p0

ex , Qlt, and average value of M

that yield reasonable fits to the growth curve and geodetic data by 

numerical calculation; also given are the pressure drop p in the 

magma reservoir and the recharge volume calculated after 500 

days of eruption.
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 479

are inclined to consider them less likely than recharge-driven 

flow. Future developments may help distinguish these possibilities: If extrusion is sustained by conduit widening without 

recharge, geodetic deflation will continue and the eruption 

will eventually wane, then stop. If it is sustained by recharge, 

geodetic deflation will soon stabilize, but the eruption may 

continue for years and end gradually as the deeper magma 

source is depleted.

Finally, one explanation for continued deflation may 

lie in the results of the numerical models. Model runs using 

Figure 11. Modeled relation between extruded volume, magma pressure and overpressure, and conduit radius at 

Mount St. Helens. A, Erupted volume, Ve. Data points with error bars represent hot-rock volume of the lava dome 

plus volume of cold rock uplifted before October 11, 2004 (compare fig. 10A). Error bars are ±5 percent as in figures 

2 and 10. B, Station JRO1 radial displacements (dots) compared with the displacements predicted at this point for 

deflation of a 15-km3

 ellipsoidal magma body at 6–10 km depth with C =7×10–11 Pa–1 and (1) a constant M =3.3×10–10

(blue dashed line) or (2) M increasing from 3.2×10–10 to 4.2×10–10 as pressure drops by 15 MPa (dashed red line). 

Negative displacements are radially inward. C, Conduit radius, R, required to maintain extrusion rate illustrated in 

figure 11A, normalized to the initial conduit radius, R0

. Magma pressure is inferred from deflation curve, assuming 

an initial overpressure of 30 MPa (blue line) or 15 MPa (black line). Method of calculating pressure change is 

explained in text.

0 50 100 150 200 250 300 350 400 450 500

A

B

C

∆p=–30 MPa

∆p=–15 MPa

40

20

0

0

–10

–20

–30

–40

–50

5

4

3

2

1

0

10

60

80

100

HR+ def. before 10/11/04

Analytic, HR+ def before 10/11/04

0 50 100 150 200 250 300 350 400 450 500

0 50 100 150 200 250 300 350 400 450 500

DAYS SINCE OCTOBER 1, 2004

JRO1 RADIAL DISPLACEMENT,

IN MILLIMETERS

NORMALIZED CONDUIT

RADIUS, R/R0

JRO1, seasonally adjusted

ERUPTED VOLUME, Ve,

IN MILLION CUBIC METERS

Variable �M = 3.23x10–10–4.12x10–10 Pa–1

Constant �M = 3.3x10–10 Pa–1

VC = 15 km3, p0 = 15 MPa ex

VC = 15 km3, p0 = 30 MPa ex
480 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

a small magma body show a rapid deflation curve at JRO1 

followed by stabilization as the recharge rate approaches the 

eruption rate. Runs that involve larger magma bodies show 

continued long-term deflation but cannot match the rapid 

deflation seen in the first several weeks of the eruption. The 

blue solid line in figure 10B represents an intermediate reservoir size that shows a little deflation still after 500 days of 

eruption but not enough to match the data. The steep deflation at short times and the long-term continued deflation at 

longer times may suggest that decompression initiated within 

a relatively small body of eruptible magma but expanded 

with time to a much larger volume that may include cooler or 

more crystalline, partially solidified material.

Discussion and Conclusions

Within the range of uncertainty, the above results are 

consistent with the view that the magma reservoir at Mount St. 

Helens is several to perhaps a few tens of cubic kilometers in 

size, that the pressure drop in the magma system is several to 

a few tens of megapascals, and that the reservoir contains less 

than a few volume percent exsolved gas at the source depth of 

8–9 km. Numerous uncertainties prevent us from more accurate estimates of the size, overpressure, volatile content, and 

degree of recharge of the Mount St. Helens magma reservoir. 

The greatest limitation is the uncertainty in reservoir compressibility C, which appears to be at least a few times greater 

over the period of this eruption than during the very brief time 

scale of seismic wave disturbance. Our conclusion that C is 

only a few to several times less than magma compressibility 

contrasts with expectations (for example, Huppert and Woods, 

2002) that M would be much greater than C in deep, equant, 

volatile-saturated magma bodies. The similarity between M

and C during this eruption likely reflects both the lower rigidity of the Earth over long time scales and the low exsolved 

volatile content of the magma.

The observation that geodetic deflation volume dVC is 

less than erupted volume Ve is an expected consequence of the 

finite compressibility of magma, and need not imply recharge. 

We expect dVC to be less than Ve during nearly all eruptions. 

Conversely, in a reservoir that is intruded before an eruption, the injection volume Vi

 should be significantly greater 

than the resulting change in reservoir volume dVC. Deep, stiff 

reservoirs containing volatile-rich magmas are likely to show 

the lowest ratios of dVC /Vi, raising the question of whether 

preeruptive magma injection into a volatile-rich reservoir 

might fail to generate a detectable geodetic signal.

Finally, like Huppert and Woods (2002) and Woods and 

Huppert (2003), we find that the compressibility of magma, 

which is related to exsolved volatile content, has an overriding influence on the duration and final volume of effusive eruptions. Our study, however, advances those works 

by showing that, when exsolution of both H2

O and CO2 is 

considered, there is little or no discontinuity in compressibility at the saturation pressure, and hence no dramatic change 

in eruptive style or the rate of decrease in eruptive activity 

when the saturation pressure is reached in a magma reservoir.

The idealizations in this paper are necessary for developing a simple model. In testing our assumptions, we acknowledge that many such idealizations are unrealistic. Our simple 

balloon-and-soda-straw cartoon of a magma reservoir may not 

even approximately resemble the complex of partially molten 

bodies that could make up the real magma system. Given this 

complexity, the question of whether the magma body is being 

recharged may be primarily a question of where one draws 

boundaries. We nevertheless hope that these simple models 

offer some insight.

Acknowledgments

Discussion of the form of the growth curve and its relation to physical processes was initiated by Terry Gerlach and 

Dan Dzurisin, to whom we owe many thanks. Mike Lisowski 

played a crucial role gathering deformation data used in this 

study and was offered coauthorship but politely declined. 

Dan Dzurisin and Peter Cervelli reviewed this manuscript 

and offered several improvements. Insights into various 

aspects of this study were provided by John Pallister, Carl 

Thornber, Richard Iverson, Shaul Hurwitz, Emily Brodsky, 

and Seth Moran.

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Appendix 1. Processes That Could 

Affect Reservoir Compressibility

The static shear modulus of large rock masses (kilometers in size) is generally known to be up to an order of magnitude less than that of laboratory-scale specimens (Pollard 

and Fletcher, 2005, p. 322). The reduction in shear modulus 

with increasing scale is generally attributed to the presence 

of fractures that can open or move (Rubin, 1990). In this 

paper we estimate host-rock shear modulus from the velocities of seismic waves whose wavelength is on the order of a 

kilometer and does not differ greatly from the dimensions of 

the rock mass under stress near the magma reservoir. On the 

other hand, seismic velocities are controlled by stress oscillations that act over a time scale of milliseconds, whereas 

eruption-associated stress changes evolve over a period of 

years. Over the longer time scale, subcritical crack growth, 

poroelasticity, and inelastic creep could deform rock and 

therefore reduce the shear modulus, G. Because the geodetic 

signal is measured at the surface, a low shear modulus at 

shallow depth could perhaps also affect the geodetic signal. 

These effects are considered below.

Crack growth and poroelasticity.—Crack growth may 

greatly reduce G at <1–2 km depth (Rubin, 1990), but at 6–12 

km depth it is unclear whether cracks of any significant size 

exist. Near the hot reservoir, cracks are likely to anneal and 

seal off interstitial fluids (Fournier, 1999). Poroelastic effects 

theoretically have no effect on shear modulus, though they can 

decrease Poisson’s ratio with time and change estimates of G

if such estimates are based on the formula G=RP

2

(1–2)/

(2(1–)) (for example, Rubin, 1990). Poroelastic effects, however, change G by only a few tens of percent at most; at depths 

of 6–12 km, interstitial fluids are likely sparse and these 

effects even smaller.

Elastic inhomogeneity.—Figure 5 shows variations in 

G with depth estimated from the seismic-velocity profile 

and density data (explained in the figure caption). In the 

uppermost 1–2 km the estimated value of G may drop to 

about 11 GPa; however, at depths below 1–2 km the value 

of G remains above about 35 GPa. Because the surface 

displacements are affected by the elastic properties of all 

materials between the magma reservoir and the surface, the 

lower elastic moduli of near-surface materials must affect 

displacements to some degree, but the relation between 

reservoir stress drop and displacement should be primarily 

controlled by rock properties near the magma body. For this 

reason we consider that the effect of less stiff near-surface 

materials on G is likely to be less than about 20 percent.

High-temperature inelastic deformation near the 

magma body.—In long-lived magma systems, such as Long 

Valley in eastern California, viscoelastic creep may reduce 

by about two-thirds the pressure change required for a 

given volume change (Newman and others, 2001; Newman 

and others, 2006). However, the Mount St. Helens magma 

system is relatively young and surrounded by cooler rock 

that is gabbroic in composition (Heliker, 1995) and resistant to creep. Regional heat-flow studies suggest that the 

ambient temperature at 9–10 km depth in this region is 

about 350°C (Blackwell and others, 1990). If we assume 

the magma reservoir has existed at its present temperature 

for about 4,000 to 40,000 years, we can estimate the temperature profile around the magma reservoir by numerically 

integrating the following one-dimensional transient equation for conductive heat flow:

R

R R

T k T r t c r r r 

∂

∂ ∂ ∂

∂ ∂   =     , (35)

where kR and cR are the thermal conductivity and specific heat, 

respectively, of the host rock, and r is the radial distance of a 

given point from the center of the magma body. Holding the 

temperature at the reservoir wall constant at 850°C and using 

kR =2 W/(m•K) (Clauser and Huenges, 1995) and cR =1,300 J/
484 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 12. Response of a 1-km-radius magma body through time, 

as a function of distance from reservoir walls. A, Temperature 

over increasing time periods since emplacement of the magma 

body. Temperatures were calculated by integrating equation 35 

with time. B, Circumferential () and radial (rr) normal stresses 

versus radial distance from the reservoir wall, calculated 

assuming elastic deformation at short time periods following a 

rapid pressure change in the magma body. C, Viscous relaxation 

time versus distance for hot rock. The solid blue curve shows 

relaxation time versus distance assuming the host rock has a 

uniform temperature of 350°C. Red curves give relaxation time 

assuming the host-rock temperature varies with distance from the 

reservoir wall as illustrated by corresponding curves in panel A.

(kg•K), we obtain the temperature profiles from a 1-km-radius 

magma body illustrated in figure 12A.

If decompression is rapid, the host rock will deform 

elastically, then relax with time as viscous creep reduces 

wall stress. The equations for the radial (rr) and normal () 

stresses near a cylindrical body in an infinite linear elastic 

medium under plane-strain conditions are (Jaeger and Cook, 

1979, p. 251):

2 2

1 2 2 1 rr

R R

p r r       = − +         , (36)

2 2

1 2 2 1 R R

p r r 

    = + −        

  , (37)

where 1 is the far-field normal stress (assumed equal in all 

directions perpendicular to the cylinder axis), p is the inter400

600

800

TEMPERATURE,

IN °CELSIUS

t = 0 yr

100 yr

4,000 yr

10,000 yr

40,000 yr

200

220

240

260

Remote normal stress

Chamber pressure

1.5 years

1 day

NORMAL STRESS,

IN MEGAPASCALS

σrr

σθθ

1,000 1,500 2,000 2,500 3,000

10 0

10 10

RELAXATION TIME,

IN YEARS

DISTANCE FROM RESERVOIR WALL, IN METERS

A

B

C

nal pressure in the reservoir, and R is the cylinder radius. For 

1=240 MPa and p=210 MPa, values of rr and  are plotted 

in figure 12B. Note that at the reservoir wall, the difference 

–rr is equal to twice (1–p).

The dominant form of stress relaxation is likely to be 

dislocation creep (Turcotte and Schubert, 2002), which involves 

a power-law dependence between normal strain rate ( or rr) 

and normal-stress difference (−rr) (Turcotte and Schubert, 

2002, eq. 7-187):

( ) /

1

a g n E R T

rr C rr e       − = − = − , (38)

where C1, n, and Ea are fitting parameters, Rg is the gas 

constant, and T is temperature (in Kelvin). We use C1=520 

MPa–n/s, n=3, and Ea=356 kJ/mol, which are appropriate for 

diabase (Turcotte and Schubert, 2002, table 7-4). The powerlaw dependence implies that viscosity is not constant at a 

given temperature; however, we can estimate a rough average 

viscosity from the relation:

( )

( )

/

1

1

E RT a

rr

eff n

rr rr

e

C











  

−

−

− ≈ = − . (39)

Using this viscosity and Young’s modulus 

( )( ) ( ) 2

P E = − + −     1 2 1 / 1 =1×1011 Pa, (where =0.25, 

tP=6.6 km/s and q=2,700 kg/m3), the viscous relaxation time  is:

eff

E = 

  . (40)

Rocks having eff<~5×1018 Pa·s will relax in less time 

than the 1.5-year duration of the eruption to date (early 

2006). For a magma body that has existed for about 4,000–

40,000 years, rocks within ~110–250 m of the reservoir wall 

will relax within this time period. If these rocks are considered part of the mechanical magma reservoir, its effective 

volume would be about 20 percent to 50 percent greater than 

the volume of magma alone. By comparison, no host rock 

was hot enough to relax during the 24-hour period following 

the Mount St. Helens eruptions of May and June 1980 (fig. 

12C). Thus the aseismic body identified by Scandone and 

Malone (1985) could be as much as a few hundred meters 

smaller in diameter than the mechanical magma body that is 

deforming during the current eruption. 

Appendix 2. Derivation of the 

Exponential Equation

Derivation of the exponential relation involves differentiating equation 11:

e

i

dp dM C Q dt dt

  = − −     (41)

and substituting equation 12 into this equation to give:
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 485

i

dp BC CQ ACp dt

= + − . (42)

Reversing the denominator in the left-hand term with the righthand side of the equation and integrating gives:

( ) 2

1 ln BC CQ ACp C t i AC

− + − + = , (43)

where C2 is the constant of integration. We find a value for C2

by noting that, at t=0, p=p0; hence

BC CQ ACp i  + − 

BC CQ ACp + −i

   0 

1 − ln = AC

( ) 0

1 0

1 ln 1 AC p p t

BC CQ ACp

 −  −   +

+ = AC  − 

. (44)

Further rearrangement leads to:

0 0 ( ) 1 B Qi ACt p p p e

A

  + − = − − −     . (45)

This equation can be substituted into equation 12 to give:

0 ( )0 ( ) 1 e ACt

i

dM Ap B Q Ap e B dt

− = + + − − −

( )0

ACt Q B Q Ap e i i

− = − + − , (46)

which can be integrated to give:

0

2

i ACt

e

B Q Ap M Qt e C

AC

+ − − = + + , (47)

where C2 is a constant of integration, which can be evaluated 

by noting that, at t=0, Me=0. After evaluating C2, we get:

( ) ( ) 0 1 i ACt

e i

Ap B Q M e Q t AC− − − = − + . (48)

This equation can be expressed in eruptive volume (Ve) by 

dividing by lava density qe

:

( ) ( ) 1 0 1 i ACt i

e

e

Ap B Q Q V e t

 AC e

− − − = − + . (49)

Appendix 3. The Exponential Equation 

for “Greased Plug” Flow

The assumption of Newtonian flow implies that the flow 

profile in the conduit is parabolic. Given the high crystal content of the magma, a more realistic scenario may be that the 

magma in the center of the conduit moves upward as a solid 

plug and that shear is concentrated along the conduit margins 

(fig. 8B). If the material in the shear acts in a Newtonian manner, a force balance leads to the following equation:

2 2 2 0 e H dM R p g R H R r dt − − = ∆  



 , (50)

where r is the thickness of the shear zone and q is the magma 

density in the conduit. The first term is the upward force at 

the base of the conduit, the second is the weight of the magma 

plug, and the third is the viscous force resisting upward flow. 

The equation can be rearranged as:

( ) 3

2

e dM R r p gH dt H

 



∆ = − . (51)

In this case, the terms A and B in equation 12 have the value 

A=qR3r/(8H) and B=R3rq2g/(8). Substitution of 

these terms into equations 14 and 15 leads to expressions for 

a and b that are identical to equations 18 and 20; hence the 

constraints on the product VC(p0–qmgH) are exactly the same 

for greased-plug flow as for Newtonian flow.

Appendix 4. Exponential Equation 

for Newtonian Flow Capped by 

a Frictional Plug with Constant 

Frictional Properties

The presence of fault gouge along the dome surface at 

Mount St. Helens suggests that, over some distance near the 

surface, magma moves upward as a solid plug with frictional 

sliding along its margins (fig. 8D). Assuming that flow below 

this plug is Newtonian, the equation for mass flux is:

( ) ( )

4

1 1

1 8

e dM R p p gH dt H

 

 = − − , (52)

where H1 is the length of conduit over which flow is Newtonian and p1

 is the pressure at the base of the frictional plug 

(fig. 8C). That pressure is a function of both plug weight and 

friction. We assume that the shear stress along the plug margin 

must exceed the normal stress (n

) times a coefficient of friction (), which is assumed to be constant:

n   ≥ . (53)

If the plug is cylindrical, vertical, of the same radius (R) as the 

conduit below, and of length H2, the pressure (p1) at the base 

of an upward-moving plug must exceed the sum of the plug 

weight and the frictional resistance: 

2

1 2

2 nH p gH R

 ≥ +  . (54)

In this case, n represents the mean normal stress on the conduit walls over the length of the plug. In solid rock, the horizontal normal stress could vary greatly even at shallow depth, 

depending on the state of gas pressure and on geometric factors. On the other hand, at Mount St. Helens, the crater floor is 

composed primarily of unconsolidated fallback from the 1980 
486 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

eruptions. If we take this material to be cohesionless, faults 

of favorable orientation will form when the ratio of effective 

normal to shear stress on any potential fault plane exceeds that 

allowed by the coefficient of friction of the material (h). Thus 

the ratio of most compressive (1) to least compressive (3) 

normal principal stress at any depth is limited to (Brace and 

Kohlstedt, 1980; Zoback and Healy, 1984):

( ) 2 1/ 2 1 2

3

1 p

h h

p

p

p



 



−   ≤ + + −   , (55)

where pp is the pore pressure in the host rock. For h=0.6 and 

pp ranging from 0 to 3, 1 /3 ranges from 1 to ~3. For this 

reason, we consider the normal stress on the conduit wall to 

be one-third to three times the vertical stress, and express the 

normal stress as the vertical stress times a constant  of order 

1. If the frictional plug extends from the surface to a depth 

H2, the mean normal stress on the plug wall is 2 = / 2 n  gH . 

We also assume that  ≅ 0.5; these simplifications allow us to 

rewrite equation 54 as:

2

1 2 1 H p gH R     ≈ +     . (56)

Inserting this value into equation 52 and noting that 

H1+H2=H, we have:

4

2

1

1

8

e

e

dV R H p g H dt H R

  

 

   =  − +     

. (57)

Hence A=eR4/(8H1) and B=R4e

2

g(H+H2 /R)/(8H1). 

These terms lead to the following values of a and b:

( ) ( ) 0

0 2

0

/ / e t e t

M e M

e t

C

dV dt dV dt

a

dV K

dt H V p g H R





 

= →∞

=

− =           − +     

≈

2

0

0.8

C

M

H V p g H R

K

      − +      , (58)

0 2

0

e e M

M t

C

dV K b AC

dt H V p g H R





  =

  = =         − +     

. (59)

Appendix 5. Exponential Equation 

for “Greased Plug” Flow Capped 

by a Frictional Plug with Constant 

Frictional Properties

As with the case above (appendix 4), this case involves 

modifying the greased plug equation to include a term for the 

pressure at the base of the frictional plug:

( ) ( ) 3

1 1

1 2

e dM R r p p gH dt H







∆ = − − . (60)

Inserting the expression for p1 in equation 56, we have:

3

2

1 2

e dM R r H p g H dt H R

qp q c



∆    =  − +     

. (61)

Hence A=R3Dr/(2H1) and B=R3Dr2g(H+H2 /R)/(2H1).

Appendix 6. Derivation of Logarithmic 

Growth Curve

We envision a one-dimensional system consisting of a 

magma-filled reservoir and conduit system applying a pressure 

p to the base of an extruding solid rock plug of mass Mp and 

displacement rate  against gravity and frictional resistance 

to slip between the plug and the conduit wall. The rate of 

extrusion is controlled entirely by frictional resistance. That 

is, in this end-member model the magma below the plug has 

negligible viscosity.

Plug force balance.—Consider a “quasi-static” force balance for motion of a cylindrical plug (force resulting from acceleration is assumed to be negligible, M d dt  ≈ 0). This quasistatic assumption is well justified by results of simulations in 

which inertia is considered; these “dynamic” simulations are not 

discussed in this appendix or elsewhere in this paper. The plug 

mass is assumed to be constant ( 0 p dM dt ≈ ), resulting from a 

balance between the rate of surface erosion of the plug and an 

equivalent subsurface accretion rate (for example, Iverson and 

others, 2006). The force balance per unit cross-sectional area in 

the conduit is:

2

2

M gp 2 H

p R R





= + , (62)

where p is the fluid pressure of the magma applied to the 

base of the plug, R is plug radius, H2 is plug height, g is the 

acceleration due to gravity, and  is the shear resistance of the 

interface between the plug and the conduit wall. Because the 

mass is assumed constant, R and H2 are also constant.

Faults have a well-known second-order dependence of 

shear strength on slip rate  (Dieterich, 1979; Ruina, 1983) 

and related, somewhat complicated dependencies on accumulated slip and time of contact (“state” effects in rate and state 

friction) (for example, Linker and Dieterich, 1992). However, 

fault strength can be assumed to be purely slip-rate dependent when subject to sustained sliding if the ratio of asperity 

contact size to slip rate is small relative to the duration of 

sustained slip. Daily extrusion rates at Mount St. Helens from 

October 2004 to October 2005, converted to boundary slip 

rates, are in the range 70 to 7,000 m/s, assuming that the plug 

has radius in the range of 25 to 75 m. Taking asperity contact 

size to be no more than 20 m, as laboratory data on rock 
22. Constraints on the Size, Overpressure, and Volatile Content of the Magma System of Mount St. Helens 487

friction suggest (for example, Dieterich and Kilgore, 1996), 

its ratio to slip rate is 0.29 to 0.0029 seconds, meaning that 

time-dependent and slip-dependent changes and friction can 

be ignored for sustained slip durations longer than 0.5 to 1.0 

seconds. Data on extruded volume are collected over intervals 

of a few weeks to a month. Therefore, we ignore complicated 

“state” effects on fault shear strength, and represent it by a 

simple slip-rate-dependent relation,

1 ln o n

o

A 

  



= + . (63)

Here n is normal stress at the wall and A1 is the rate dependence of fault strength. The constant o

 is an arbitrary reference, the strength of the wall interface when the interface slip 

rate is o .

For shear of thick fault-gouge layers (for example, 

Byerlee and Summers, 1976; Marone and others, 1990), for 

shear near the brittle-ductile transition (for example, Blanpied 

and others, 1998), and for near-surface faulting (Marone and 

Scholz, 1988), fault strength increases with slip rate (velocity strengthening, rate strengthening). Because all of these 

conditions are present at the plug wall at Mount St. Helens, we 

expect that fault slip is predominantly rate strengthening, and 

thus A1 in equation 63 is a small positive constant, typically 

between 0.001 and 0.03 (Marone and Scholz, 1988; Blanpied 

and others, 1998).

The volume of extruded material Ve is the product of the 

conduit cross-sectional area and the slip at the wall, and the 

extruded mass Me is proportional to the extruded volume, so 

equation 63 is equivalently

1 ln e

o n

o

M A

M

   = + , (64)

where notation for the mass rate of extrusion e e dM dt M=

is used. Combining equations 62 and 64 leads to a relation 

between magma pressure and the rate of plug extrusion

ln e

o

o

M

p p M= +  , (65)

where 1 2 2 / A RHn  =  and 2

0 0 2 p H R Mg R = + 2 / /   are 

constants.

Magma mass balance.—During plug extrusion, the 

volume and driving pressure of the magma will change. To 

characterize these changes we consider, in turn, mass and 

volume balances for the magma. We define the magma mass 

M V M M M =  as the mass of magma in the magma reservoir and 

in the conduit below the frictional plug. The rate of change of 

mass can be expressed as

M M M

M M

dM dV d V

dt dt dt

 = +  , (66)

where M and VM are magma density and volume, respectively. 

The mass change rate is also equivalent to the difference 

between the rate of mass input to the system Qi

 and the rate 

out of the system Qo, or

M

i o

dM Q Q dt = − . (67)

Because magma is not being extruded at the surface, Qo represents the magma-volume loss due to magma freezing onto the 

plug (Iverson and others, 2006). Combining and rearranging 

so that fluid volume is the dependent variable yields

M M M i o

M M M

dV V d Q Q

dt dt



  

= − + − . (68)

Reference to the magma density and its time derivative can be 

replaced by the pressure dependence through expanding the 

density derivative in equation 68:

M M d dp

dt p dt

 ∂

∂ = (69)

and using the definition of the elastic compressibility of the 

magma ( )( ) 1/ / M M M  ≡  ∂ ∂  p . Making these substitutions 

into equation 68 leads to

1 M o i

M M M M

dp dV Q Q

dt V dt   

  = −  + −   

. (70)

Magma volume balance.—The volume of the magma 

system increases as the solid plug is extruded at a rate dVe /dt, 

and decreases as magma freezes to the plug at the rate −Qo /qM. 

We also allow the walls of the magma system (reservoir and 

conduit) to respond elastically to changes in magma pressure 

using a representative reservoir and conduit compressibility c so that the rate of elastic change of magma volume is

c M V dp dt. The combined rate of magma volume change is 

then 

M e o

c M

M

dV dV Q dp V

dt dt dt





= − + . (71)

Combining equations 70 and 71 yields the relation between 

the mass rate of extrusion and the rate of change of the driving 

pressure

( )

e e i M

M C M

dp M Q

dt V

 

 

− = − + , (72)

where the extruded mass is M V e e e =  .

Solutions

Solutions for extrusion rate with time can be found by 

taking the time derivative of equation 65

e

e

dp dM

dt M dt

 = & (73)

and equating to equation 71, resulting in the single differential 

equation

( )

2

e 1 e e i

M c M e M

dM M M Q

dt V    

  = −  −  +  

. (74)
488 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

If the magma volume is large relative to the extruded volume, then VM can be treated as a constant and equation 74 is 

separable. In the solution that follows we assume negligible 

density contrast between the magma and plug, e =  M , as 

justified elsewhere in this paper.

No Recharge.—When Qi=0 the extrusion rate is:

0

0 / e

e

D M M

tM D



  = + , (75)

where M0 is the extrusion rate at t=0, and D V= + M C M ( )   . 

The cumulative mass of extruded material goes as

0

e e ln 1

e

tM M D

D  

 

  = +    

. (76)

This is the logarithmic form that well characterizes stress 

relaxation due to fault slip in some laboratory experiments 

(Reinen and others, 1994) and during earthquake afterslip (for 

example, Marone and others, 1991; Schaff and others, 1999).

With recharge.—If Qi

>0 the extrusion rate is

0

1 1

i

e

i

e Q t

i D

Q M

Q e

M

 

− =  

− −    

. (77)

Note that equation 77 is for Qi>0 and does not easily reduce to 

equation 75 for Qi=0.

Cumulative extruded mass goes as































 = −− D

tQ

i

ee e

i

e

Q

M DM   11ln 0 . (78)

This expression can be converted to erupted volume Ve by 

dividing by e.
Heading 1 489

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490 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 491

Crisis Management 

Mount St. Helens reawakened during the age of electronic media and 

an unprecedented public appetite for real-time information. The volcano’s 

legacy of a large explosive eruption in the 1980s—still within the memory 

of many—added an element of interest and drama to the volcanic unrest of 

2004. The effort to satisfy that interest and meet the information needs of 

both media and the public, while simultaneously providing prompt hazard 

information to Federal, State, and local government officials, forms a story of 

its own, told here in two chapters on crisis management.

The emergency-response discipline has, over the years, recognized the 

value of Joint Operations and Information Centers as useful management 

tools. These centers are implemented on a temporary basis in response to 

crisis situations—be it wildfire, severe weather, or volcano related—and are 

purposely terminated as a crisis winds down. Such centers operated for about 

two weeks early in the 2004 eruption to help coordinate interagency operations and maintain a consistent, centralized source of current information 

about volcanic activity, hazards, and area closures.

The case is made herein for the value of advance planning and teamwork during the actual crisis. Successes during the eruption that began in 

2004 stem from the depth to which partner agencies with responsibilities 

for Mount St. Helens have built their collaborative networks during the past 

quarter century. State and county emergency response agencies, the Gifford 

Pinchot National Forest, and the U.S. Geological Survey have been working together in the business of volcano-hazard awareness and risk mitigation since the 1980s. The framework of cooperation and understanding 

among these groups has been longstanding, originated through rank-and-file 

employees and passed unbroken to successive managers.
492 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

View southwest on July 26, 2005, with spine 5 on the right and a partially 

distintegrated spine 4 in center. USGS photo by S.P. Schilling.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 23

1

 Gifford Pinchot National Forest, Mount St. Helens National Volcanic 

Monument, 42218 NE Yale Bridge Rd, Amboy, WA 98601

2

 Gifford Pinchot National Forest, 10600 NE 51st Circle, Vancover, WA 98682

Managing Public and Media Response to a Reawakening 

Volcano: Lessons from the 2004 Eruptive Activity of 

Mount St. Helens

By Peter M. Frenzen1 and Michael T. Matarrese2

Abstract

Volcanic eruptions and other infrequent, large-scale 

natural disturbances pose challenges and opportunities for 

public-land managers. In the days and weeks preceding an 

eruption, there can be considerable uncertainty surrounding 

the magnitude and areal extent of eruptive effects. At the same 

time, public and media interest in viewing developing events 

is high and concern for public safety on the part of local land 

managers and public safety officials is elevated. Land managers and collaborating Federal, State, and local officials must 

decide whether evacuations or restrictions to public access are 

necessary, the appropriate level of advance preparation, and 

how best to coordinate between overlapping jurisdictions. In 

the absence of a formal Federal or State emergency declaration, there is generally no identified source of supplemental 

funding for emergency-response preparation or managing 

extraordinary public and media response to developing events. 

In this chapter, we examine responses to escalating events that 

preceded the 2004 Mount St. Helens eruption and changes in 

public perception during the extended period of the largely 

nonexplosive, dome-building eruption that followed. Lessons 

learned include the importance of maintaining up-to-date 

emergency-response plans, cultivating close working relationships with collaborating agencies, and utilizing an organized 

response framework that incorporates clearly defined roles and 

responsibilities and effective communication strategies.

Introduction

Mount St. Helens has undergone major changes in volcanic activity and land-management direction since 1979. Public 

perception has run the gamut from “tranquil, snow-covered 

mountain” to “notorious killer volcano” and, most recently, to 

“celebrated volcanic attraction and research laboratory.” These 

transformations, together with the recent return to eruptive 

activity in 2004, provide useful insight into how infrequent 

events such as eruptions can influence people’s perception of 

natural hazards. A person’s perception of “normal volcanic 

behavior” changes as a function of degree of personal experience with a volcano and time since the last eruption. Managers 

need to be prepared to address a wide array of public perceptions and responses as they seek to provide for public access, 

education, and visitor safety in these dynamic landscapes. Our 

experience at Mount St. Helens suggests that, in the days and 

weeks leading up to a potentially explosive eruption, it is the 

management of people and their responses to perceived events 

that poses the greatest challenge.

Events Shape Human Responses

Since the 1980 eruption, management and education 

programs at Mount St. Helens have been developed largely 

in response to catastrophic eruptive events and the prevailing 

lens of public perception. Between 1980 and 1986, memory of 

the catastrophic eruption was intense, and agency efforts were 

largely centered on emergency response, restoration of damaged 

resources, and creation of the congressionally designated Mount 

St. Helens National Volcanic Monument (U.S. Department of 

Agriculture Forest Service, 1984). Response to volcanic activity 

was managed by an Emergency Coordination Center (ECC) at 

the Gifford Pinchot National Forest headquarters in Vancouver, 

Washington. After explosive eruptions ended in 1980, geologists 
494 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

focused their efforts on monitoring an intermittently growing 

lava dome and developing increasingly sophisticated methods 

of eruption forecasting. In the early 1990s, several years after 

the dome-building eruption ended, the general consensus among 

geologists and emergency managers was that eruptive activity 

that began in 1980 had run its course. The ECC was discontinued, and an emergency response plan was developed that 

formalized calldown procedures and the role of collaborating 

State and local authorities in the event of renewed activity (U.S. 

Department of Agriculture Forest Service, 1992). During the 

next decade, in the absence of volcanic activity, emphasis on 

volcano emergency planning was reduced and efforts focused 

on restoration of damaged roads and construction of visitor 

facilities in and around the Monument (U.S. Department of 

Agriculture Forest Service, 1984).

Following an 18-year period of quiet, the rapid acceleration of events leading up to the 2004 eruption surprised many 

geologists and emergency managers. Agency managers quickly 

shifted their focus from visitor education and protected-area 

stewardship to management of a fast-paced media event and 

renewal of multiagency working relationships and contingency 

closure zones around the volcano. Public reaction to renewed 

volcanic activity varied with people’s memory of the catastrophic 1980 eruption and familiarity with volcanic processes. 

The level of concern of some residents and emergency responders was heightened by their memory of the largely unpredicted, 

devastating lateral blast and geologists’ uncertainty expressed 

prior to the 1980 eruption about the expected degree of explosiveness and areal extent of the eruption.

Planning is Key to an Effective Response

Following the 1980 eruption, land managers implemented 

revised land-use allocations in the Monument, on the basis of 

existing hazards assessments and recent experience. Federal 

acquisition of private and leased lands effectively created an 

uninhabited 8-km (5-mi) buffer around the volcano. Facilities and roads were placed on ridges above the level of valleys 

draining the volcano outside of immediate hazard zones. Visitor Center roofs were designed to support the combined weight 

of projected ash fall and precipitation. This groundwork greatly 

reduced the potential hazard to life and property and simplified 

the situation faced by emergency managers in 2004. Federal 

ownership of adjacent lands also greatly facilitated the October 2, 2004, evacuation and identification of 5-, 8-, and 11-km 

radius (3, 5, and 7 mi) contingency closure zones around the 

volcano. Mount St. Helens offers a compelling example of the 

importance of incorporating volcanic-hazards mapping in landuse planning, road location, and facility design.

The rapid pace of public and media response to events 

leading up to the 2004 eruption provided a vivid reminder 

of the importance of an up-to-date emergency-response plan 

and clearly defined roles for collaborating local, State, and 

Federal responders (table 1). Within hours of the release of 

the initial U.S. Geological Survey (USGS) Information Statement on September 24, 2004, media flocked to the USGS 

Cascades Volcano Observatory (CVO). The Monument’s 

visitor-center staff received hundreds of media phone calls, 

conducted numerous drop-in interviews, and saw increased 

visitation. Live media coverage of the eruption greatly accelerated the pace of events both for scientists monitoring the 

eruption and monument employees. Two days after the initial 

Information Statement, the USGS issued an Alert Level 1: 

Notice of Volcanic Unrest, the lowest of their three alert levels, triggering emergency calldown procedures and initiating 

a series of coordination meetings and consultations between 

Federal, State, and local officials. Prompt notification and 

consultation proved to be important, because in only nine 

days the volcano progressed from no activity (background 

levels of seismicity) to rapid deformation of the crater floor 

and steam and ash eruptions. The pace and intensity of media 

and public response and need for thoughtful coordination 

among Federal, State, and local partners proved to be a challenge for participating agencies. A key lesson learned during 

the 2004 eruption has been the importance of developing a 

shared understanding of interagency roles and responsibilities and of ensuring a timely flow of information at both the 

field and leadership levels.

Emergency Preparedness Requires a Long-Term 

Commitment

Before the onset of renewed volcanic activity in 2004, it 

was difficult for the Gifford Pinchot National Forest staff and 

multiagency partners engaged in the press of daily business to 

find time to update the emergency-response plan and calldown 

list. Fortunately, status of the emergency-response plan was 

monitored as one of the National Forest’s internal performance 

measures, and an updated response plan was completed in 

2003. Planning efforts were largely a paper exercise, however, 

and many years had passed since the participating agencies last 

engaged in a table-top response exercise or a field-implementation drill―the need for which has been an important lesson 

learned during the 2004 eruption. Since that time, participating 

local, State, and Federal agencies have engaged in a table-top 

exercise, and Monument employees have conducted periodic 

readiness reviews and field-implementation drills. The future 

challenge will be maintaining awareness of volcanic hazards 

and a long-term commitment to effective interagency response 

as memory of the 2004 eruption fades.

Use of the Incident Command System

Volcanic eruptions and other large-scale disturbances 

can potentially impact large areas, triggering emergency 

responses by numerous agencies from multiple, often 

overlapping jurisdictions. Coordination of the response to 

the 2004 eruption was greatly facilitated by activation of 

the Incident Command System (ICS; Federal Emergency 

Management Agency, 2004). ICS provided the framework 
23. Managing Public and Media Response to a Reawakening Volcano: Lessons from the 2004 Eruptive Activity 495

for bringing together a regional, type-2 Incident Management Team (IMT–2) composed of Federal, State, and local 

emergency responders (fig. 1).

In 2004, following a USGS issuance of their highest alert 

level (Alert Level 3: Volcanic Alert), operations were directed 

through a Unified Command composed of a lead Incident 

Commander (IC) who shared command responsibility with a 

co-incident commander (Co-IC) from the Washington State 

Emergency Management Division (EMD) and a rotating 

Co-IC representing sheriffs from the four counties surrounding the volcano. Both the pace and efficiency of the response 

effort benefited from the resulting interagency coordination 

and unified voice.

The ICS is a highly organized, flexible structure that was 

developed for responding to fire and other emergency and 

nonemergency incidents on Federal lands. A major strength 

of ICS is the clear delegation of authority from the land-managing agency to the Incident Commander. The local Agency 

Administrator sets broad incident objectives and delegates 

management authority and responsibility for all aspects of the 

incident to the IC. This allows the IC to focus on a safe and 

effective response to the incident while the Agency Administrator focuses on day-to-day operations of the surrounding 

Co-Incident Commander

staffed by four-county

rotating Sheriffs’ Representative

Agency Administrator

(Gifford Pinchot National Forest

Supervisor or designated

representative)

Lead Incident Commander (USDA–FS),

Washington Incident

Management Team No. 4

Co-Incident Commander,

Washington Emergency

Management Division

INFORMATION

 •Joint Information Center

 •Visitor Center

SAFETY

 •Safety Planning

 •Monitor Compliance

AIR OPERATIONS

 •Observation

 •Air Traffic Control

DIVISIONS

AND

GROUPS

Visitor Centers Road Blocks

and Closures

OPERATIONS PLANNING

 •Daily incident

 action plan

 •Daily incident

 resources summary

 •Establish contingency

 closure zones with

 State and local

 emergency managers

LOGISTICS

 •Ordering

 Personnel

 Supplies

 •Communications

 Radios

 Telephones

 Computers

 •Medical

 •Food

FINANCE

 •Purchasing

 •Timekeeping

 •Compensation

 claims

 •Cost accounting

 and control

UNIFIED COMMAND

Figure 1. Organization chart for the short type-2 Incident Management Team that responded to the 2004 eruption of 

Mount St. Helens, Washington. USDA–FS, U.S. Department of Agriculture, Forest Service.

area. In cases where multiple jurisdictions are involved in 

response to a single incident, a Unified Command is established with command responsibility shared between two or 

more ICs, each representing his or her respective agency and 

Agency Administrator.

ICS offers the advantage of a uniform organizational 

structure composed of working groups that are universally recognized throughout the emergency-response community. This 

enables personnel trained in ICS functions to come together 

on a case-by-case basis and operate as an effective team. The 

size of an IMT is based on the size, complexity, or duration 

of the incident. On small (type 3) incidents, the IMT generally is staffed by personnel drawn from the local managing 

agency. As the incident grows in complexity—or if its duration 

exceeds local staffing capability—a larger, regional (type 2) 

IMT takes command. When size or complexity exceeds the 

capacity of a type-2 team, a national (type 1) IMT is brought 

in to manage the incident.

ICS is organized around five principal components or 

groups: Command, Planning, Operations, Logistics, and 

Finance (fig. 1). The Command component has primary 

authority and is responsible for setting overall objectives and 

priorities. Information and Safety are included in the Com-
496 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

mand component because of their overall importance to ICS 

and the management of an incident. The Planning group 

develops action plans and collects and evaluates information 

about the status of the incident and available resources. The 

Operations group develops and conducts tactical operations 

to implement the plan; the Logistics group provides needed 

resources and support; and the Finance group provides cost 

accounting and procurement.

On October 3, 2004 (day 11 of the volcanic crisis), the 

Gifford Pinchot National Forest staff ordered a type-2 Incident Management Team known as a “short team” (fig. 1). 

Short teams are abbreviated versions of an IMT that provide 

an abbreviated command team and utilize local resources to 

fill in needed ICS functions. The National Forest staff utilized 

a Washington IMT that already included Gifford Pinchot 

employees and local collaborating agencies who were familiar with the volcano. The short team carried a full Command 

Staff composed of an Incident Commander, Deputy IC, Safety 

Officer, and Information Officer. Command was supported 

by a General Staff composed of Section Chiefs for Planning 

(responsible for contingency planning with collaborating 

agencies and preparation of daily action plans); Operations 

(responsible for managing area closures, patrols, and traffic management); Logistics (responsible for supporting field 

operations and the Joint Operations Center); Finance (responsible for managing business operations and cost containment); 

and Air Operations (responsible for a fixed-wing observation 

aircraft and managing air space around the volcano).

Factors Affecting Initial Response

The early phases of the 2004 eruption of Mount St. 

Helens did not fit any of the normal agency criteria for an 

incident response, which posed some interesting challenges 

for agency managers and field employees. Public interest was intense, but the eruption did not trigger resources 

for the logistical support and staffing normally associated 

with response to a forest fire or other large-scale event. At 

its public-interest peak, there were 24 satellite trucks at the 

volcano, media calls were arriving from around the world at 

Figure 2. Television and print media gather at Castle Lake Viewpoint, 

Washington, on October 4, 2004. USGS photo by L.G. Mastin. Inset: 

Monument Scientist conducts twice-daily joint press conferences with 

U.S. Geological Survey scientists at the volcano. Inset photo by T.S. 

Warren, copyright © 2004, AP/WIDE WORLD PHOTOS. 
23. Managing Public and Media Response to a Reawakening Volcano: Lessons from the 2004 Eruptive Activity 497

Day Date Volcanic events Management challenges and responses Lessons learned

1 Sept. 23 A swarm of small, shallow earthquakes begins at 0200 (depth less 

than 1 km, Md

 less than 1) with 

200 events recorded by 1700 

PDT. The previous notable earthquake swarm occurred at 9 km 

depth in spring/summer, 1998.

Monument Scientist is notified by 

USGS–CVO, briefs other FS officials. 

CVO suggests earthquake swarm 

may be rock-fracturing from elevated 

ground water due to heavy rainfall.

On-staff science expertise helps transfer information from CVO to FS 

officials. Close working relationship with CVO staff proves invaluable for interpreting uncertain and 

rapidly changing events.

2 Sept. 24 CVO issues Information Statement describing earthquake 

swarm beneath 1980–86 lava 

dome and increased probability 

of small steam explosions in the 

crater. Number of seismic events 

peaks at midday and then starts 

to decline.

Monument posts information on Web 

site and at trailheads to notify climbers and hikers. Media flock to CVO, 

and visitor center (VC) staff responds 

to calls from media across the country. Staff familiarity with Mount St. 

Helens geology was an important 

ingredient in information transfer and 

effective response.

Monument response to fast-paced 

events was assisted by availability of employees well-trained in 

volcanic processes through past 

interpretive training by CVO staff. 

Managers of newly active volcanoes 

may need to bring in additional outside expertise for the short-term and 

(or) provide for needed employee 

training.

3 Sept. 25 Seismicity continues to decline 

through the afternoon and then 

begins to increase.

County sheriffs, emergency responders, 

and adjacent land managers all want to 

receive the latest information. FS and 

CVO initiate regular conference calls 

to brief collaborating State and local 

agencies.

The timely flow of information is 

a key element of an effective 

response. Since 2004, lead agencies have developed streamlined 

calldown procedures to facilitate 

the distribution of information and 

reduce duplication of effort.

4 Sept. 26 Shallow seismicity increases with 

10 larger events (magnitude 2.0 

to 2.8). Alert Level 1: Notice 

of Volcano Unrest is issued, 

the first such alert since October 

1986. Character of some earthquakes suggests involvement of 

pressurized fluids (gas or steam) 

or perhaps magma. Increased 

possibility of small explosions, 

ashfalls above the crater rim, or 

small landslides and lahars.

Volcano is closed to climbing, and trails 

immediately north of the crater are 

closed. Satellite trucks, reporters, and 

volcano visitors continue to arrive, 

and a media center is established on 

a ridge west of VC. As number of 

media and visitors grows, logistics 

becomes increasingly challenging. By 

design, the FS Emergency Coordination Center (ECC) and Forest-Level 

Incident Management Team (IMT-3) 

are not activated until Alert Level 2.

Since 2004, the monument has developed local emergency-response 

procedures to support field operations during the early phases of a 

volcano-driven or public-interestdriven event. The plan recognizes 

the importance of maintaining 

strong working relationships between FS personnel and their local, 

State, and Federal counterparts. History of past collaboration with CVO 

and availability of CVO scientists 

for twice-daily briefings aids media 

response.

5 Sept. 27 Seismicity continues to increase 

slowly, although no events 

greater than magnitude 1.5 in last 

24 hours. CVO crews report new 

crevasses in crater glacier south 

of 1980–86 lava dome. Gas 

flight does not detect magmatic 

gas.

FS officials review response plan and 

meet with collaborators to discuss 

interagency roles and responsibilities. 

The number of agencies and jurisdictions involved poses a challenge 

in the establishment of a Unified 

Command. Maintaining coordination 

and information flow is increasingly 

difficult as key personnel engage in 

interagency coordination and are assigned to VCs and other remote field 

sites.

2004 eruption demonstrated the need 

for maintaining up-to-date emergency-response plans that include 

clearly defined roles and lines of 

supervision for participating agencies. It is essential that co-incident 

commanders have a clear delegation 

of authority from the agencies that 

they represent. Periodic table-top 

exercises and response drills can 

help develop a shared vision of an 

effective multiagency response.

Table 1. Initial management actions and lessons learned by local Ranger District personnel during days 1 through 5 of the 2004 

eruption of Mount St. Helens, Washington. 

[Volcanic events from this volume, Scott and others, chap. 1; Moran and others, chap. 2; Moran and others, chap.6. USGS, U.S. Geological Survey; CVO, Cascades Volcano Observatory; FS, U.S. Department of Agriculture Forest Service; Md

, earthquake coda magnitude.]
498 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

a rate of two per minute, and daily visits to the Monument’s 

Web site exceeded 15 million (fig. 2).

By design, the Volcano Emergency Response Plan and 

Emergency Coordination Center (ECC) were not activated 

until CVO issued a Notice of Volcanic Unrest, three days after 

the earthquake swarm began. Activation of a local Incident 

Management Team (IMT–3) and associated logistical support was triggered when the Alert Level 2 (Volcano Advisory) 

was issued three days later (day 7). A major lesson learned 

in 2004 was the speed with which public and media response 

can outpace the actual progression of eruptive events. The 

task of responding to continuous, live media coverage and the 

many thousands of visitors who flocked to the volcano greatly 

exceeded the actual response needs generated by geologic 

events in what proved to be a remarkably quiet dome-building 

eruption (fig. 3). The Monument’s response plan now includes 

provisions for augmented staffing and logistical support triggered by media and public response events irrespective of 

predetermined volcanic alert levels.

The use of predefined trigger points in emergency-response 

planning is useful because it focuses resources where and when 

they are most needed. In the case of the 2004 eruption, activation of ICS by the volcano alert system (Alert Level 1 on day 

4 and Alert Level 2 on day 7; table 1) effectively compressed 

many critical response tasks into perhaps the busiest and most 

uncertain three-day period of the entire eruptive period. During 

Figure 3. In the days following the evacuation of the Johnston 

Ridge Observatory, Washington, thousands of volcano watchers 

gather along State Route 504. USGS photo by M.P. Poland. Insets: 

Emergency managers were faced with a large “tailgate party” 

along the primary evacuation route northwest of the volcano.

U.S. Forest Service photo by R.M. Petersen.
23. Managing Public and Media Response to a Reawakening Volcano: Lessons from the 2004 Eruptive Activity 499

this three-day period, senior managers from the Gifford Pinchot 

National Forest and the Mount St. Helens National Volcanic 

Monument were engaged in coordination efforts involving 

multiple levels of the U.S. Department of Agriculture’s Forest 

Service organization and those of collaborating local, State, and 

Federal agencies. Preparations were simultaneously underway 

to (1) establish a Joint Operations Center (JOC) managed under 

a multiagency Unified Command to respond to a potential largescale eruption; (2) establish a Joint Information Center (JIC; 

Drieger and others, this volume, chap. 24) to handle steadily 

increasing national and international media coverage; and (3) 

activate a local Incident Management Team (IMT–3) to provide 

much-needed logistical support to field employees at the volcano. Local, State, and Federal managers were faced with considerable uncertainty as to the size of the potential eruption, the 

degree of hazard in adjacent areas, and the extent of a closure 

zone to implement around the volcano (fig. 4). In the absence of 

a State or Federal declared emergency, funding for associated 

personnel and response activities was also uncertain. To complicate matters further, these activities occurred at the end of the 

Federal fiscal year during a period of restricted purchasing and 

fiscal authority.

The combination of escalating earthquake activity and 

uncertainty about the anticipated eruption produced a heightened level of concern among emergency responders. Geologists at CVO did an excellent job of describing the most 

probable eruptive scenarios but were also careful to point out 

a broad range of less likely but potentially more destructive 

outcomes. Given the recent history and memory of the catastrophic eruption of May 18, 1980, geologists were careful to 

Figure 4. Local, State, and Federal officials meet with Gifford Pinchot National Forest and U.S. Geological Survey officials on 

October 11, 2004, to discuss potential volcanic hazards and to define contingency closure zones. Photo courtesy of Pat Pringle.
500 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

frame potential eruptive scenarios within the context of how 

they differed from 1980. Geologists described the comparatively lower hazards associated with an open volcanic crater 

versus the magma-induced “bulge” that formed during the 

months preceding the 1980 eruption and failed so catastrophically. In the days preceding the Volcanic Alert and resulting 

October 2, 2004, evacuation of Johnston Ridge, there was 

considerable discussion about the presence or absence of 

gas-rich magma beneath the volcano. Geologists constructed 

probability trees in an effort to connect monitoring data with 

potential scenarios and to quantify the probability and magnitude of an explosive eruption.

The range of potential eruptive outcomes and potential for 

explosive activity formed the context within which local, State, 

and Federal agencies organized the emergency-response system. 

As live media coverage (fig. 2) fed a growing public interest 

and thousands of visitors gathered on Johnston Ridge, agency 

managers and geologists were actively discussing a potential 

pullback and the closure of an area 8 km (5 mi) radially around 

the volcano. Officials were concerned about potential human 

responses to a significant explosion or ashfall and the challenges posed by the evacuation of a large numbers of visitors on 

a single highway in steep, mountainous terrain. State and local 

officials who remembered the 1980 eruption were concerned 

about maintaining viable evacuation routes and providing an 

appropriate level of response in adjacent communities.

The steadily escalating tempo of earthquakes and rapid 

progression from Volcanic Advisory to Volcanic Alert (days 

7 through 9; table 2) tested the emergency-response system. 

Implementation of the JOC and multiagency Unified Command 

was complicated by the fact that participating local, State, and 

Federal agencies each brought their own understanding of ICS, 

Unified Command, and its application to the process (table 2). 

Considerable effort was expended to work out the delegation 

of authority and lines of supervision associated with having a 

Unified Command composed of local, State, and Federal CoIncident Commanders (lead IC from the IMT–2 for Federal; a 

representative of the Washington Emergency Management Division for State; and a rotating representative of the four county 

sheriffs for local jurisdictions). Communications and coordination were further challenged by the number of agencies, 

jurisdictions, and geographic locations involved in the response 

(CVO, JOC, GP National Forest headquarters, Monument visitor centers, and other remote sites).

The importance of instituting a Unified Command structure early in the process was a key lesson learned during the 

renewed eruptive activity of autumn 2004. Early activation 

allows collaborators time to familiarize their agencies with 

emerging issues, to make needed adjustments, and to implement needed agreements and delegations of authority. Early 

collaboration is important given the number of State and local 

agencies and jurisdictions involved in emergency response 

on National Forest and adjacent lands. Interagency response 

planning and periodic implementation drills can contribute to a 

shared understanding of roles, responsibilities, and supervisory 

structures and result in a more effective response.

Extended Eruptions Pose Challenges

During the fall of 2004, high levels of public and media 

interest were driven largely by the novelty of renewed eruptive activity and the opportunity to witness small steam and 

ash explosions. As the frequency of explosions diminished 

and winter weather increasingly obscured crater views, public 

interest evolved into fascination with the steaming volcano 

and amazement at the pace and longevity of the continuing 

eruption. Agency managers established closure boundaries 

and gate systems that restricted access to within 8 km (5 mi) 

of the volcano while providing access to adjacent areas for 

traditional forest activities (table 3). Since September 2004, 

the Monument has engaged in a sustained effort to manage 

evolving area closures, ensure adequate staffing during periods 

of increased visitation, and maintain a level of emergency 

preparedness among Monument employees. As geologists’ 

confidence that any large-scale change in behavior will be 

detected by the monitoring network has increased, National 

Forest managers have reopened facilities and trails, restoring 

public access to most of the area surrounding the volcano.

The quiet, nonexplosive nature of the 2004 eruption provided a relatively safe opportunity for the public to watch volcanic processes firsthand and to increase their awareness of volcanic hazards in the Pacific Northwest. Since the fall of 2004, 

millions of people have learned about the ongoing eruption and 

monitoring through media coverage, Web sites, and personal 

experiences at the volcano. The continuing challenge for agency 

managers is to reaffirm the lessons learned during the 2004 

eruption response and to periodically review and update interagency response plans and procedures. Given the intermittent 

nature of eruptive activity, we must adopt a long-term view and 

be prepared to act appropriately as volcanic events and associated public and media responses occur in the future.

Acknowledgments

The response to the 2004 eruption was truly a group effort. 

Our sincere thanks go to staff of the Gifford Pinchot National 

Forest; Mount St. Helens National Volcanic Monument; USGS 

Cascades Volcano Observatory; Washington Incident Management Team No. 4; Washington Emergency Management 

Division; Sheriff’s Departments and supporting emergency 

responders from Clark, Cowlitz, Lewis, and Skamania Counties; 

Washington State Patrol; Washington Department of Transportation; Washington Department of Natural Resources; and 

the Federal Emergency Management Agency. Our response to 

worldwide media interest was greatly assisted by the availability 

of information officers from local, State, and Federal agencies 

whose voluntary assistance made the Joint Information Center 

possible. The manuscript benefited greatly from early reviews 

by Lynn Burditt and Tom Knappenberger and formal reviews by 

Drs. Shigeo Aramaki and Chris Newhall.
23. Managing Public and Media Response to a Reawakening Volcano: Lessons from the 2004 Eruptive Activity 501

Table 2. Management actions and lessons learned following activation of the Emergency Coordination Center and type-3 Incident 

Management Team at the Gifford Pinchot National Forest Headquarters during days 7 through 11 of the 2004 eruption of Mount St. 

Helens, Washington.

[USGS, U.S. Geological Survey; CVO, Cascades Volcano Observatory; FS, U.S. Department of Agriculture Forest Service; Md, earthquake coda magnitude.

See table 1 for sources of volcanic events.]

Day Date Volcanic events Management challenges and responses Lessons learned

7 Sept. 29 Shallow seismicity accelerates overnight with four events per minute 

and increasing number of Md

 2–3 

events. CVO issues Alert Level 

2: Volcano Advisory, cautions 

that explosions, crater ballistics 

and ash clouds could occur at 

any time. GPS equipment detects 

northward movement of 1980–86 

lava dome. No magmatic gas 

detected.

Emergency Coordination Center 

(ECC) and forest-level incidentmanagement team (type 3) is 

activated, not all positions are filled. 

After 7 days, media-response capability at CVO and Visitor Centers 

begins to be strained. Federal and 

State officials discuss establishing a 

Joint Information Center to handle 

increasing demand for information 

and to allow CVO staff to focus on 

monitoring and eruption forecasting. Additional information officers 

arrive to help with media at Visitor 

Centers.

In 2004, public and media response 

to volcanic events was rapid and 

posed the biggest challenge to land 

managers. Current National Volcanic 

Monument response plans recognize 

the need for logistical support to field 

operations independent of volcano 

alert level. Plans include a trained 

cadre of local volcano-information 

officers to assist with phone calls and 

media inquiries at the Visitor Centers 

during the critical early phases of a 

volcanic or media-response event.

9 Oct. 1 CVO issues Information Statement.

Seismicity continues at 1–2 events 

per minute, with largest up to Md

 3. 

Observations reveal uplift of Crater 

Glacier by several meters. At noon 

a small, 20-minute steam and ash 

explosion opens a vent in uplifted 

glacier. Elevated CO2

 detected on 

1980–86 dome, and weak sulfurous 

odor but no SO2

 or H2S.

Local, State, and Federal officials 

meet to discuss implementation of 

Unified Command and establishment of a Joint Operations Center. 

CVO and FS officials discuss 

potential hazards and trigger points 

for closure of the Johnston Ridge 

Observatory and viewpoints closest 

to the volcano. Visitor Center staff 

request additional help to handle 

expected crowds of volcano watchers for the coming weekend.

Maintaining timely flow of information to FS managers and field sites 

is the key to maintaining situational 

awareness and preparedness. Since 

2004, CVO and FS have developed 

streamlined calldown procedures. As 

the eruption has continued, daily contacts have been replaced by weekly 

conference calls to keep Monument 

staff up-to-date on the latest monitoring information.

10 Oct. 2 Vigorous ~1-hour-long, lowfrequency tremor occurs at 1215 

PDT. Given that such tremors may 

indicate magma movement or 

pressurization, CVO issues Alert 

Level 3: Volcano Alert. Following the tremor, shallow seismicity 

continues at 1–2 per minute, with 

largest event Md

 3.

Visitor Center staff and State and local 

counterparts evacuate Johnston 

Ridge Observatory (JRO) and State 

Route 504. In less than an hour, 

2,500 visitors and 14 satellite trucks 

are safely relocated. State highway, 

lands, and airspace within 8 km of 

volcano are closed. Some State and 

Federal officials express concern 

that evacuation of JRO exceeded 

the pace of official calldown procedures, but others view the 2004 

evacuation of JRO as an example 

of front-line employees acting 

decisively to provide for public and 

employee safety.

The Monument emergency plan recognizes that field employees may need 

to act decisively in the interests of employee and visitor safety. Empowering 

field employees to implement clearly 

defined procedures can also be useful 

in the event that lines of communication fail. While empowerment of 

front-line employees is important, it is 

also critical to ensure that actions are 

well coordinated and communications 

are maintained across collaborating 

agencies and up and down the chain 

of command.

11 Oct. 3 25-minute low-frequency tremor 

occurs at 0250 PDT. Magnitude 3 

earthquakes occur at a rate of one 

every 5 minutes. Large-scale uplift 

and fracturing of Crater Glacier 

continues.

Joint Information Center (JIC) is established at Gifford Pinchot National 

Forest Headquarters, providing relief 

for CVO staff and reducing mediacall volume at Visitor Centers. CVO 

has difficulty reaching ECC afterhours contacts and finally reaches 

Visitor Center housing.

Afterhours calldown procedures have 

been amended to ensure that Visitor Centers are contacted directly 

by CVO. Maintaining vigilance in 

response to numerous seismic events 

and steam emissions over the long 

term is challenging.
502 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Day Date Volcanic events Management challenges and responses Lessons learned

12 Oct. 4 22-minute-long steam and ash 

emission (3,700 m). Vent 

area is a bubbling lake. Visual observations assess tens 

of meters of uplift of Crater 

Glacier. Magma is at shallow 

level and could soon reach 

surface. Increased likelihood 

of larger steam and ash emissions. Gas flight detects CO2

and low levels of H2

S.

FS brings in a regional Incident Management Team (IMT type 2) because 

Monument and Forest resources are 

becoming overextended. Unified 

Command coordinates the multiagency 

response effort. Providing resources 

to support IMT and field operations is 

challenging because a potential future 

eruption does not fit within the normal 

criteria for emergency-response funding 

or trigger an emergency declaration.

IMT-2 provides needed logistical support, 

organization, and supervision for increasingly complex multiagency effort. 

A key lesson learned is the importance 

of defining when Monument, National 

Forest, or regional-response officials 

are in charge. Since 2004, the Monument response plan clearly states that 

the Monument Manager is Incident 

Commander until the Forest Supervisor 

activates a Forest-level IMT.

13 Oct. 5 At 0905 PDT, a 70-minutelong steam and ash emission 

(4,500 m) deposits dusting 

of ash 60 miles to northeast. 

The ash plume is visible on 

Doppler weather radar. Seismicity drops and remains at 

low levels. Status remains 

at Alert Level 3: Volcano 

Alert.

Emergency response is directed by Incident Commander (IC) and two other 

co-ICs representing Washington State 

Emergency Management Division and 

four county sheriffs. Joint Operations Center is established. Resources 

include gate guards, traffic-control 

personnel at Visitor Centers, and fixedwing observation aircraft.

IMT-2 provides welcome relief for 

Monument employees, many of whom 

have been on duty for more than 12 

days. Current response plans recognize that Monument staff may need to 

be assisted by other Forest employees 

and resources during fast paced media 

events and prior to activation of a Forest or Regional IMT.

14 Oct. 6 Seismicity remains at reduced 

level. Probability of eruption 

that threatens life and property is decreased, so CVO 

steps back to Alert Level 2: 

Volcano Advisory. Rainfall 

overnight generates small 

debris flows in the crater. 

Low clouds and rain limit 

visibility and air operations.

IMT-2 and Forest representatives meet 

with local and State law enforcement 

and emergency managers to define 

closure zones; considerable discussion 

about the value of linking closure zones 

to specific alert levels. CVO and FS 

stress the importance of maintaining 

flexibility so closures can be adjusted 

according to current eruptive behavior 

and potential threats.

The Incident Command System (ICS) 

and Unified Command provide a useful structure for organizing a complex, 

multiagency response. However, in the 

absence of an emergency declaration or specific response funding, 

cost containment is a real concern. 

Response plans must ensure that key 

ICS functions are accounted for in a 

local response organization.

15 

to 

25

Oct. 7

 to

Oct.17

Shallow seismicity continues. 

After magma surfaces (day 

19), seismicity gradually decreases. Small crater debris 

flows occur with rainfall. 

Status remains at Alert 

Level 2: Volcano Advisory.

Public and media interest declines as explosive activity subsides and weather 

obscures volcano. IMT-2 departs and 

operations are transitioned back to 

local IMT-3. As hazards diminish, 

closures are lifted and staffed temporary gates are replaced with unstaffed, 

permanent gates.

In absence of explosions, media and 

public interest in ongoing eruptive 

activity decreases. Outreach and 

information efforts are key to maintaining awareness of ongoing events 

and potential future hazards. Periodic 

press conferences and field visits aid 

outreach effort.

26

 to 

575

Oct.18, 2004

 to

June 2006

Continuous eruption and formation of spines. Extrusion 

rate and seismicity gradually 

decline. Periodic small steam 

and ash emissions with dome 

rockfall. Status remains at 

Alert Level 2.

Maintaining calldown procedures and 

response capability over months and 

years of continuous eruptive activity is 

a challenge. As comfort level with continuing nonexplosive eruption grows, 

additional areas around the volcano are 

reopened.

Confusion can occur when staff unfamiliar with volcanic processes and 

terminology relay technical information during calldowns. Need to ensure 

that contacts are knowledgeable about 

monitoring terminology and volcanic 

hazards.

Table 3. Management actions and lessons learned following activation of a short type-2 Incident Management Team from day 12 and 

onward in the 2004 eruption of Mount St. Helens, Washington.

[USGS, U.S. Geological Survey; CVO, Cascades Volcano Observatory; FS, U.S. Department of Agriculture Forest Service; Md, earthquake coda magnitude.

See table 1 for sources of volcanic events.]
23. Managing Public and Media Response to a Reawakening Volcano: Lessons from the 2004 Eruptive Activity 503

References Cited

Driedger, C.L., Neal, C.A., Knappenberger, T.H., Needham, 

D.H., Harper, R.B., and Steele, W.P., 2008, Hazard information 

management during the autumn 2004 reawakening of Mount St. 

Helens volcano, Washington, chap. 24 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. Geological Survey Professional Paper 1750 (this volume).

Federal Emergency Management Agency, 2004, National Incident Management System: U.S. Department of Homeland 

Security, 139 p. [http://www.fema.gov/emergency/nims/

nims_compliance.shtm; last accessed May 22, 2006].

Moran, S.C., Malone, S.D., Qamar, A.I., Thelen, W.A., 

Wright, A.K., and Caplan-Auerbach, J., 2008a, Seismicity associated with renewed dome building at Mount St. 

Helens, 2004–2005, chap. 2 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Moran, S.C., McChesney, P.J., and Lockhart, A.B., 2008b, 

Seismicity and infrasound associated with explosions at 

Mount St. Helens, 2004–2005, chap. 6 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Scott, W.E., Sherrod, D.R., and Gardner, C.A., 2008, Overview of the 2004 to 2006, and continuing, eruption of 

Mount St. Helens, Washington, chap. 1 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

U.S. Department of Agriculture Forest Service, 1984, Mount 

St. Helens National Volcanic Monument, Final Environmental Impact Statement, Comprehensive Management Plan: 

Vancouver, Wash., Gifford Pinchot National Forest, 450 p.

U.S. Department of Agriculture Forest Service, 1992, Mount 

St. Helens Contingency Plan: Vancouver, Wash., Gifford 

Pinchot National Forest, prepared in cooperation with 

the U.S. Geological Survey; Clark, Cowlitz, Skamania 

and Lewis County Sheriffs; Washington State Emergency 

Management Division; Federal Emergency Management 

Agency; U.S. Army Corps of Engineers; Pacific Power; and 

Portland General Electric, 48 p.

U.S. Department of Agriculture Forest Service, 2003, Mount 

St. Helens Volcanic Activity Response Plan: Vancouver, 

Wash., Gifford Pinchot National Forest, 30 p.

U.S. Department of Agriculture Forest Service, 2005, Mount 

St. Helens National Volcanic Monument, Initial Emergency 

Response Plan: Vancouver, Wash., Gifford Pinchot National 

Forest, Mount St. Helens National Volcanic Monument, 30 p.
504 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 24

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, 4200 University Drive, Anchorage, AK 99508

3

 USDA Forest Service, Gifford Pinchot National Forest, 10600 NE 51st 

Circle, Vancouver, WA 98682; now at USDA Forest Service Pacific Northwest 

Region Office, Box 3623, Portland, OR 97208.

4

 Clark Regional Emergency Services Agency, 710 W 13th St., Vancouver, 

WA 98660

5

 Washington Military Department, Emergency Management Division, 

Bldg. 20, Camp Murray, WA 98430

6

 University of Washington, Pacific Northwest Seismic Network, Box 

351310, Seattle, WA 98195

Hazard Information Management During the Autumn 2004 

Reawakening of Mount St. Helens Volcano, Washington

By Carolyn L. Driedger1, Christina A. Neal2, Tom H. Knappenberger3, Deborah H. Needham4, Robert B. 

Harper5

, and William P. Steele6

Abstract

The 2004 reawakening of Mount St. Helens quickly 

caught the attention of government agencies as well as the 

international news media and the public. Immediate concerns 

focused on a repeat of the catastrophic landslide and blast 

event of May 18, 1980, which remains a vivid memory for 

many individuals. Within several days of the onset of accelerating seismicity, media inquiries increased exponentially. 

Personnel at the U.S. Geological Survey, the Pacific Northwest 

Seismic Network, and the Gifford Pinchot National Forest 

soon handled hundreds of press inquiries and held several 

press briefings per day. About one week into the event, a 

Joint Information Center was established to help maintain a 

consistent hazard message and to provide a centralized information source about volcanic activity, hazards, area closures, 

and media briefings. Scientists, public-affairs specialists, and 

personnel from emergency-management, health, public-safety, 

and land-management agencies answered phones, helped in 

press briefings and interviews, and managed media access to 

colleagues working on science and safety issues. For scientists, in addition to managing the cycle of daily fieldwork, 

challenges included (1) balancing accurate interpretations of 

data under crisis conditions with the need to share information 

quickly, (2) articulating uncertainties for a variety of volcanic 

scenarios, (3) minimizing scientific jargon, and (4) frequently 

updating and effectively distributing talking points. Success 

of hazard information management during a volcanic crisis 

depends largely on scientists’ clarity of communication and 

thorough preplanning among interagency partners. All parties 

must commit to after-action evaluation and improvement of 

communication plans, incorporating lessons learned during 

each event.

Introduction

In late September 2004, a sudden and rapidly accelerating increase in seismicity beneath Mount St. Helens heralded 

the onset of the first volcanic event of consequence in the 

contiguous 48 states since 1986 (Dzurisin and others, 2005). 

In addition to its volcanologic significance, the 2004 eruption captured the attention of the public worldwide, who, even 

after 24 years, vividly recalled the images and impacts of 

Mount St. Helens’ catastrophic eruption on May 18, 1980. The 

2004–2006 eruption was also the first in the contiguous 48 

states since the establishment of widespread Internet use and 

development of the 24-hour international news cycle and its 

attendant around-the-clock demands. An Internet Web camera 

(VolcanoCam) installed and maintained by staff at the Gifford 

Pinchot National Forest (GPNF) gave unprecedented access to 

visual images of a Cascade Range eruption in near-real time. 

Real-time seismic information in the form of Internet-available 

seismograms maintained by the Pacific Northwest Seismic 

Network (PNSN) enabled the general public to monitor the 

volcano from home computers. At Mount St. Helens National 

Volcanic Monument (MSHNVM), scenic overlooks, such as 

the Johnston Ridge Observatory (JRO), provided an unprecedented view for thousands of visitors (fig. 1). All of these 

factors combined to create an enormous and urgent demand 
506 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

for information and commentary regarding the volcano and 

its activity, hazards, likely outcomes, the daily activities of 

scientists, and any other available information about Mount 

St. Helens, one of the world’s most famous volcanoes. This 

chapter describes some of the pre-event planning, the real-time 

development of strategies to respond to this demand, significant challenges, and lessons learned. Additional details regarding the first few weeks of unrest and eruption can be found in 

other contributions in this volume (Scott and others, chap. 1; 

Moran and others, chap. 2; Qamar and others, chap. 3; Moran 

and others, chap. 6).

Pre-Event Coordination and Planning

Since the mid-1990s, the U.S. Geological Survey (USGS) 

has produced a series of modern volcano-hazard assessments 

for each potentially active volcano in Washington and Oregon, 

including Mount St. Helens (Wolfe and Pierson, 1995). Each 

assessment was written in a format accessible to both technical and lay audiences and was distributed to public officials, 

educators, and public libraries in areas at risk.

Soon after this series of publications was completed, 

USGS staff began working with partner agencies in Washington and Oregon to develop hazard-response plans based on the 

assessments. These response plans define the roles of individual agencies and protocols for cooperation during volcanic 

unrest, such as the plan for Mount St. Helens (Gifford Pinchot 

National Forest, 2003). As of spring 2006, volcano-response 

plans now exist for Mounts Baker, Rainier, St. Helens, and 

Hood, and for Glacier Peak. Similar plans are in progress 

for Mount Adams and for the volcanoes of central Oregon. 

Figure 1. Visitors at Mount St. Helens National Volcanic 

Monument’s Johnston Ridge Observatory, Wash., view 

the October 1, 2004, explosion of steam and ash from vent 

approximately 9 km south, in crater. USGS photo by E.Y. Iwatsubo.

Ironically, the USGS was working with the GPNF and other 

agencies to update the Mount St. Helens plan when the 2004 

volcanic unrest began.

In addition to collaborating on volcano response protocols, the USGS maintains active cooperation with its partners 

in hazard communication on the Federal, State, and local 

levels. This includes planning and practicing rapid initiation 

of a Joint Information Center (JIC), sharing expertise at news 

briefings, and assembling interagency-communication calldown lists. Since the 1980s unrest at Mount St. Helens, staff 

at the University of Washington’s (UW) PNSN in Seattle have 

coordinated with the USGS in the development and release 

of Volcano Information Statements and Volcanic Alert Level 

changes. This close collaboration, refined over the years, 

enabled separate but well-synchronized media responses to 

the rapidly evolving unrest at Mount St. Helens that began on 

September 23, 2004.

Chronology of Events

The initial earthquake swarm on the morning of September 23, 2004, now commonly viewed as the onset of 

volcanic unrest, was noted in a Cascade Range update posted 

on the USGS-Cascades Volcano Observatory (CVO) Web 

site at 1800 PDT, but it brought little immediate inquiry from 

the media. The following morning CVO and PNSN released 

an Information Statement regarding the earthquake swarm, 

and veteran public viewers of seismograms on the PNSN 

Web site began to recognize the potential significance of the 

seismicity. This recognition culminated in approximately 

one dozen media inquires per day during the next few days 

at CVO. The PNSN, which provided information to Seattle 

television and radio stations and newspapers, was inundated 

by requests regarding the earthquakes. The growing public 

interest in the unrest resulted in a deluge of requests for 

information from members of the local press and other media 

providers. The PNSN published a press release on its Web 

site that described the onset of the earthquake swarm and 

provided links to seismicity pages on its own Web site and 

to Web sites of CVO, GPNF, and other sources. The PNSN 

Seismology Lab staff extended their operation to seven days 

per week, often working late into the night. As seismicity 

increased, the PNSN placed a disclaimer on its Web site, 

noting that postings for recent earthquakes were incomplete 

because staff could not keep up with processing these seismic events during regular working hours.

On September 26, the USGS issued an official Alert 

Level 1: Notice of Volcanic Unrest (unusual activity detected; 

Dzurisin and others, 2005), prompting inquiries to increase at 

both CVO and UW from about one dozen to more than 40 per 

day. Scientists at the USGS and PNSN consulted one another 

early each day about the increasing seismicity. After these conversations, USGS staff wrote formal talking points for use by 

information scientists, who were scientists recruited to speak 
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 507

to the media. On September 29, the USGS issued an Alert 

Level 2: Volcano Advisory (eruption likely but not imminent) 

and began to offer formal media briefings.

Staff at CVO and the USGS Office of Communications 

worked hard to address all inquiries in the face of increasing 

interest. Scientists identified the range of eruption potential. 

The ability to compile, synthesize, and understand the significance of the events was strained by mounting demands 

for information. The U.S. East Coast news cycle resulted in 

interviews at 0400 PDT, challenging USGS personnel as they 

struggled to maintain adequate 24-hour staffing. Only a few 

individuals rotating through the informal position of information scientist had the breadth of knowledge about all aspects 

of the volcanic unrest to answer all questions. Scientists found 

it challenging to interpret and distill the continuous stream of 

field data into timely public statements.

At the PNSN, television satellite trucks filled the parking lot, and staff and student volunteers struggled to keep up 

with phones and interview requests. In an attempt to make 

the monitoring and media response more sustainable, staff 

were ordered home after a 12-hour shift, and a number of key 

staff alternated on 12-hour shifts. By Monday, September 

27, the PNSN Web site became completely congested, and 

system administrators were forced to configure high-capacity 

servers on the UW network backbone to serve the PNSN site.

During this time, USGS staff supplied information to the 

news media (many of whom kept satellite trucks present all 

day and night), critical operational partners, and other scientists, primarily by telephone and by the release of once- or 

twice-daily text updates by e-mail, faxes, and postings on 

the CVO Web site. Independently, colleagues at other USGS 

offices, the PNSN, the GPNF, and the Washington State Emergency Management Division (EMD) handled other inquiries. 

The communications landscape faced by USGS scientists is 

shown schematically in figure 2.

By September 28, five days into the event, USGS communications and technical staff were fielding 60 or more 

inquiries per day. It became apparent that media and public 

demands for information would soon exceed the capacity 

that could be managed effectively by any single agency. That 

notion, and the concern for the consistency and centralization 

of hazard messages, prompted the USGS to work with agency 

partners from the GPNF and the Washington EMD to plan and 

establish a JIC, described more completely later in this chapter. On September 29, GPNF officials and local emergency 

managers arriving at CVO to discuss the JIC encountered 

numerous media trucks in the parking lot and had to step over 

media cables that snaked through the CVO lobby, preventing 

the doors from closing at this normally secure government 

facility. It was clear that, in order for CVO staff to function, 

a JIC facility would need to be located away from CVO. The 

GPNF headquarters in Vancouver, only six miles distant, was 

chosen as a logical site.

During this time at the PNSN, UW Computing and 

Communications staff installed fiber-optic television circuits 

in the Seismology Lab to help meet the continually increasing demand for staff interviews, which now included national 

and international media providers. To help manage the more 

than 100 daily phone calls, a six-line phone bank was installed 

in the conference room adjacent to the lab. This facility was 

staffed by students and also was used for interviews with 

national and international media members.

As the rate and magnitude of earthquakes under Mount 

St. Helens continued to build, speculation about the nature of 

the coming eruption became the focus of many press reports. 

Daily teleconferences between PNSN and USGS–CVO 

scientists were crucial in formulating and sharing assessments 

of the volcano. Frequent consultation helped keep the public 

message from CVO and PNSN reasonably consistent.

On October 1, with JIC provisioning still in progress, a 

20-minute-long steam and ash emission sent ash more than 3 

km above the crater (Dzurisin and others, 2005; Moran and 

others, this volume, chap. 6). News of this event escalated 

inquiries to more than 80 per day at CVO. At PNSN, requests 

for information lessened slightly as attention shifted to CVO 

and the MSHNVM Castle Lake Viewpoint.

On October 2, intensified seismicity, including an hourlong period of energetic seismic tremor, prompted USGS–

CVO to issue its highest warning level for the Cascade 

Range, an Alert Level 3: Volcano Alert, which means that 

a volcanic event threatening to life and property appears 

imminent or is underway. That afternoon millions of viewers 

worldwide watched a media briefing at CVO presented to 

U.S. Department of the Interior Secretary Gale Norton and 

a host of congressional dignitaries. Its political significance 

and public interest confirmed, Mount St. Helens unrest had 

become an incident of national and international prominence.

Late on the afternoon of October 2, the JIC opened its 

doors at GPNF headquarters. The media, at first reluctant 

to move from CVO, were eventually persuaded to relocate 

owing to the disciplined refusal of CVO staff to answer their 

inquiries onsite and by the ready availability of expertise 

at the new JIC. All information sources, including publicinformation officers (PIOs; a term that we use in a general 

sense to include a variety of professionals, other than volcanologists, who communicated with the media and public), 

scientists, staff at the USGS Office of Communications, and 

agency phone recordings, referred media to the JIC as the 

principal source of volcano information.

The JIC was fully activated from October 3 to October 

13, spanning the buildup to and start of lava extrusion (Dzurisin and others, 2005). After October 13, media attention 

diminished sufficiently to allow JIC staff to return to normal 

agency facilities. Remaining calls were addressed at CVO by 

the outreach staff, information scientists, and USGS Office of 

Communications staff. The latest information, usually from 

the daily update, was made available on a 24-hour media line 

maintained and updated by the GPNF. Importantly, according to Incident Command principles for an evolving, ongoing 

situation, former JIC employees maintained contact so that in 

the event of later heightened activity, the information systems 

could be reestablished quickly. As of this writing in October 
508 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Live and phone media interviews 

Other scientific groups 

Informal inquiries by phone and e-mail 

Airlines 

Media 

Emergency and land-management agencies 

Partnerships—Federal, State, local 

Links to monitoring data 

Updated interpretations 

Updated photographs 

Mount St. Helens National Volcanic Monument 

Pacific Northwest Seismic Network (PNSN) 

USGS Office of Communication 

Non-USGS Science Groups 

Non-JIC Media Centers 

Individual Contacts 

Official Groups 

Web-Based Data 

E-mail 

Internet fax 

Web site 

Daily Updates 

Develop consensus on activity and hazards—USGS and PNSN 

Develop talking points 

Science Staff Meetings 

Joint Operations Center 

Press briefings at CVO, Joint Information Center, and at volcano 

Formal Briefings 

Alert Level Changes 

Daily ashfall maps 

Scenario maps 

Probability trees 

Communication Tools 

Primary Functions 

USGS Scientists 

Monitoring 

and 

communicating 

activity to public 

officials 

and the public 

2006, no escalation of activity requiring JIC reactivation has 

occurred at Mount St. Helens.

Media interest in Mount St. Helens decreased owing to 

cessation of dramatic steam and ash explosions and the onset 

of cloudy weather, which blocked views of the volcano. Media 

managers were unwilling to pay for satellite trucks sitting idle 

and were less willing to report daily about lava dome-building 

events when interesting visual images were unavailable.

Development of Joint Operations 

Center and Joint Information Center

During the onset of volcanic unrest at Mount St. Helens, 

USGS and PNSN scientists and other public officials conducted their own hazard information dissemination according 

to each agency’s protocols. As unrest escalated and the demand 

Figure 2. Graphic illustrating shared responsibilities of U.S. Geological Survey (USGS) scientists for 

communication with a broad range of individuals and agencies. CVO, Cascades Volcano Observatory.
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 509

for a more coordinated message grew, the agencies formally 

pooled their resources in a unified response that employed the 

precepts of the Incident Command System (ICS), a standard, 

on-scene, all-hazards incident management system already in 

use by emergency responders and codified nationally in 2004 

as the National Incident Management System (NIMS; Federal 

Emergency Management Agency, 2004; Frenzen and Mataresse, this volume, chap. 23). Under this arrangement, agencies 

work together through designated members of a unified command (often the senior person from each agency) to establish 

a common strategy and a single Incident Action Plan (Federal 

Emergency Management Agency, 2004). In the case of the 2004 

Mount St. Helens event, those stakeholder agencies created a 

Joint Operations Center (JOC; the term coined in present-day 

terms by NIMS is an “Incident Command Post”; Christine 

Jonientz-Trisler, oral commun., 2006). The JOC focused on 

operational aspects of the response (not scientific or research 

aspects) and took responsibility for the creation of the JIC.

This system of JOC and JIC (under NIMS terminology the JIC is called a “Joint Information System”) is a 

well-known tool used in the ICS today. The JOC-JIC system 

enables interagency coordination, support for decision makers, 

flexibility based on changing circumstances, and, through the 

JIC, development and delivery of consistent messages. The 

system includes plans and protocols to provide information 

during incidents. A JIC (there can be several, although one 

is preferable) generally exists in a location where PIOs from 

involved organizations colocate to provide critical emergency 

information, crisis communications, and public-affairs functions. Although the public can receive information from many 

sources, a JIC allows the various organizations with responsibility during an incident to come together to ensure clear, 

timely, and consistent hazard messages.

The JIC can be viewed as a central hub for communication, a “one-stop shopping” facility where representatives from 

cooperating agencies can address inquiries with a single voice. 

A JIC can be assembled when communications needs exceed 

the capacity of individual agencies. The JIC provides interagency coordination and integration, development and delivery 

of coordinated messages, support for decision makers, and flexibility to meet demands based on changing circumstances. The 

use of a JIC reduces confusion, inaccuracies, and duplication of 

efforts and can help address rumor control. Each organization 

maintains its own authority and policies but still contributes to 

an overall unified message to the public (Federal Emergency 

Management Agency, 2004). A general call center and other 

office resources are shared, and technical representatives from 

specific agencies address inquiries germane to their agency’s 

interest. The JIC is dynamic and can be resized to meet incident 

needs. To the detriment of the host agency, the JIC can displace 

normal agency operations for extended periods, such as the two 

weeks during 2004 that the headquarters conference room was 

unavailable to GPNF staff.

Physical Description, Responsibilities, 

and Operation of Joint Information Center

The Mount St. Helens JIC leadership (fig. 3) included the 

JIC Manager who reported to the Incident Commander at the 

JOC. The JIC Manager was responsible for functioning of the 

JIC and for maintaining records. Liaisons from Washington 

EMD, GPNF, and USGS were in constant contact with the JIC 

Manager or served in that role at some point during JIC operation. The Media Briefings Facilitator, Products Coordinator 

(news releases, for example), Call Center Supervisor, and PIOs 

at MSHNVM all reported to the JIC Manager, as did information scientists and other technical specialists. As needed, each 

agency in the command structure provided PIOs to coordinate 

their home agency’s information (both internally and externally), act as spokespeople, and provide appropriate technical 

expertise. The PIOs also sent talking points to other information 

outlets, such as the PNSN, USGS staff in other cities, and to 

MSHNVM visitor facilities.

Media Briefings Products Call Center Technical Specialists Other Information

Centers

Agency 

Representatives

GPNF, USGS, State

Incident 

Commander

Joint Information

Center (JIC)

Manager

Advisory

Figure 3. Organizational chart 

for the Joint Information Center 

indicating the external and internal 

command structure. GPNF, Gifford 

Pinchot National Forest; USGS, 

U.S. Geological Survey.
510 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Fax Machines

Podium with audio bridge

Cameras

Briefing Room

Messages

Whiteboard

Television

Rumors

Whiteboard

Public 

Information

Officers

InformationNeeds

Whiteboard

Call 

Center

JIC

Management

References/Supplies

Technical

Experts

Whiteboard-Screen-Backdrop

Seating Seating

The JIC staff used a conference room at the GPNF headquarters (fig. 4). The room was dividable by heavy, soundproof curtains. During operation of the JIC, one-half of the 

room served as the Media Briefing Room, with areas reserved 

for television cameras, seating, and a podium. The other half 

of the conference room was arranged by function; one side 

housed the Call Center, and the other side housed the JIC 

Manager, technical experts, the news release production team, 

and the media briefings team.

The JIC Call Center consisted of a temporary six-line 

switchboard system that allowed call takers to simultaneously 

answer multiple incoming lines and to transfer media and public-affairs calls to the technical experts as needed. Call Center 

staff included employees of the GPNF and other national 

forests, USGS, Washington EMD, Oregon’s Department of 

Geology and Mineral Industries, Washington’s Department of 

Geology and Earth Resources, and many emergency-management and nontechnical public-affairs professionals from across 

the greater Portland-Vancouver metropolitan area. The Federal 

Emergency Management Agency (FEMA) also assisted with 

staffing and could have provided more formal assistance and 

resources had the incident escalated to an emergency or had 

an official Federal disaster declaration been issued. FEMA 

also pre-positioned management staff at the JOC and provided 

liaisons to the States of Washington and Oregon Emergency 

Operations Centers and a liaison to the USGS at CVO.

Call Center staff conveyed basic information about the 

current incident to media representatives, reading from daily 

talking points, updates compiled on a dry-erase whiteboard, 

and fact sheets compiled by the technical experts (fig. 5). This 

information satisfied many of the media callers. When an 

interview was requested and a technical expert was unavailable, the request was given to the Media Interview Coordinator. Detailed information (reporter name, media organization, 

publication or broadcast schedule or deadline, and nature of 

the questions) allowed the Media Interview Coordinator to 

prioritize pending requests and to assign USGS or other staff 

to the appropriate interview. This system was highly effective 

in meeting the vast majority of live interview requests. Technical experts, along with the JIC management from Washington 

Figure 4. Layout of Joint Information Center (JIC), housed in conference room of Gifford 

Pinchot National Forest headquarters in Vancouver, Wash. The conference room was divided 

into three parts: a Media Briefing Room, a Call Center, and a section for JIC management and 

technical specialists.
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 511

EMD, the U.S. Department of the Interior Bureau of Land 

Management, and the GPNF and other national forests, posted 

new information prominently on the whiteboard so that it 

could be provided to media callers almost instantaneously. 

The Call Center’s close proximity to scientists ensured that the 

nontechnical PIOs used the appropriate scientific descriptions.

Whereas CVO provided daily updates on volcanic activity and alert levels on its Web site, the JIC distributed news 

releases by e-mail or fax from information derived largely 

A

B

C

from CVO. Staffers combined rosters of media contacts from 

several agencies and amended them continuously in an effort 

to maximize distribution of information and to assure media 

that they would be informed of changes of volcano status. 

The JIC also established a recorded phone message line that 

summarized the current volcano conditions. This telephone 

number was widely advertised to the media, and the message 

was updated daily and provided in English and Spanish. The 

JIC also provided general volcano-hazards information that 

previously had been translated into eight languages.

News briefings at the JIC became an important regular 

source of authoritative information. The JIC initiated a rigorous, 

controlled protocol for briefing time, participation, and followup 

interviews with scientists or other officials. Briefings were 

conducted live on camera by designated scientists from USGS, 

often with equipment and graphics to illustrate content. Outof-area media and radio station representatives called in during 

live briefings by way of a telephone audio link. The audio link 

enabled remote listeners to ask questions of presenters after the 

briefing. USGS Office of Communications staff often managed 

the media briefings. The JIC conducted two briefings on most 

days and a single briefing as demand decreased. During a period 

of several weeks, approximately 38 formal briefings were given 

at the JIC and at MSHNVM (see below).

Information Centers at Mount St. 

Helens

In addition to the JIC in Vancouver and the PNSN in 

Seattle, several other information centers provided updates in 

close coordination with the JIC. These included the interpretive facilities of MSHNVM, especially the Coldwater Ridge 

Visitor Center and nearby Castle Lake Viewpoint (fig. 6). 

At its peak, the Castle Lake Viewpoint hosted 24 television 

satellite trucks and accompanying media. Six GPNF PIOs 

and USGS information scientists staffed the viewpoint and 

held twice-daily news briefings. The multitude of media 

and public visitors at this remote viewpoint was substantial, 

and MSHNVM staff were soon overtaxed with the strain of 

conducting media interviews and providing crowd control for 

thousands of tourists who flocked to the monument. Media 

preferred to operate from Castle Lake Viewpoint because of 

its proximity to Mount St. Helens; however, scientists and the 

JOC preferred indoor briefings at the JIC in Vancouver, in 

close proximity to CVO.

Figure 5. Mount St. Helens Joint Information Center (JIC) housed 

at Gifford Pinchot National Forest headquarters, Vancouver, 

Wash. (USGS photos by C.L. Driedger, October 2004). A, Call 

Center, staffed by public-information officers from local agencies. 

B, JIC management and technical specialist center. C, Media 

Briefing Room.
512 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Help in providing information also came from several 

colleges and universities with knowledgeable Earth science 

faculty, from State geological surveys in Washington and 

Oregon, and from other USGS offices throughout the nation. 

We have no satisfactory method to quantify the load of inquiries on these outlets.

Joint Information Center and Web Site 

Statistics

To document and evaluate the work of the JIC, staff 

tracked the number and type of inquiries (table 1) and 

maintained a master list of participants. During its 11-day 

operation, the JIC was staffed by 70 people from 30 different 

emergency-management, health, safety, and land-management agencies and the USGS, most of whom came from the 

Portland-Vancouver area. These individuals responded to more 

than 750 e-mail inquiries and more than 800 telephone inquiries from media staff in 12 countries.

Not surprisingly, the number of inquiries increased with 

a rise in real-time seismic amplitude measurements (RSAM) 

that reflected increased seismicity at the volcano (fig. 7). 

Inquiry frequency spiked when the alert level was raised, when 

volcanic tremor occurred, and when there were visible events, 

such as steam and ash explosions. USGS–CVO Web site statistics illustrate the same intense and event-driven demand for 

online information.

The PNSN Web site, where information about seismic 

activity related to volcanic unrest was available, received 31 

million hits on the Web servers (equaling approximately 10 

million pages viewed) between September 28 and October 5. 

Early during the event, in anticipation of this intense interest, 

UW staff separated the public Web site servers from computer 

servers that process scientific information, potentially saving 

both systems from failure.

As at PNSN, the record high number of hits on the 

CVO Web site prompted USGS system administrators to 

add servers and then to contract with temporary commercial 

service providers to accommodate the demand for online 

information. In September 2004, before onset of the erupSources of inquiries to the Joint 

 Information Center

Number of 

inquiries

Television (local and affiliate) 278

Print (specific news publications) 228

Other (Associated Press wire service, 

public, and others)

163

Radio (local and affiliate stations) 137

Web-based news services 6

Table 1. Sources and numbers of inquiries to the Joint 

Information Center, Gifford Pinchot National Forest 

headquarters, Vancouver, Wash., from October 3 to 

October 13, 2004.

Figure 6. Castle Lake Viewpoint, Mount St. Helens National Volcanic Monument, Wash., located 

approximately 15 km from Mount St. Helens, served as a nexus for media representatives during October 

2004. Monument staff directed all media representatives to this viewpoint, dubbed “Satellite City,” where 

media queries could be addressed efficiently by local staff and U.S. Geological Survey scientists. As many as 

24 media trucks were on site at the height of media interest. USGS photo by L.G. Mastin, October 5, 2004.
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 513

22 26 30 04 08 12 16 20

0 0

20

40

60

80

100

120

140

160

180

200

220

200

400

600

800

1,000

1,200

1,400

1,600

2004

Swarm 

onset

Unrest

Advisory

Alert

Advisory

Dome 

extrusion

September October

10-MINUTE RSAM, IN COUNTS

Phone inquiries

Explanation

Tremor events

Steam-and-ash explosions Web site inquiries Lahar activity

0

200

400

600

800

1,000

1,200

1,400

1,600

CVO WEB SITE PAGES REQUESTED,

IN THOUSANDS OF INQUIRIES

MEDIA PHONE INQUIRIES

tion, there were approximately 34,000 Web pages requested 

per day. During the steam and ash explosion on October 4, 

2004, 1.43 million Web pages were requested—a 42-fold 

increase in daily CVO Web page requests. In the 2-week 

time period between September 24 and October 7, Web users 

requested 8.4 million pages, and 11.2 million pages were 

requested during October 2004. As of this writing in October 

2006, the number of Web pages requested remains high. The 

volume of Web page requests has, on occasion, inexplicably 

risen tenfold, far exceeding October 2004 levels. CVO Web 

site access has yet to return to pre-October 2004 levels (Lyn 

Topinka, oral commun., 2006).

The majority of GPNF Web site hits were to the Mount 

St. Helens VolcanoCam Web page, which received an estimated 131 million hits (equaling approximately 18 million 

Web pages requested) between September 23 and October 31, 

2004. The VolcanoCam’s popularity resulted in a manyfold 

increase in Web page requests to the GPNF Web site.

Update Protocols at Cascades Volcano 

Observatory

USGS scientific and communications staff generated text 

updates on the status of the volcano each morning on the basis 

of consensus at daily scientific meetings. Although the JIC 

staff relied upon these daily updates, volcanic events sometimes rendered them obsolete. For example, during a scheduled news briefing at the JIC, scientist Willie Scott was interFigure 7. Relation between volcanic activity, alert levels, volume of media inquiries to the 

Cascades Volcano Observatory, and number of Web pages requested from the CVO Web site 

(Lyn Topinka, oral commun., 2005) during the first weeks of unrest at Mount St. Helens. Realtime seismic amplitude measurement (RSAM), in red, is a proxy for the level of seismicity at 

volcano. Media phone inquiries (daily counts, in tan) and public Web activity (hachured) generally 

increased with the rise in RSAM and spiked following changes of alert level and prominent 

volcanic events, such as steam and ash explosions.
514 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

rupted by a news reporter who announced that a steam and 

ash explosion was in progress. Scott was forced to abandon 

his prepared text and interpret the eruption as seen on a small 

monitor provided by one of the television stations.

Before the onset of the 2004 volcanic unrest at Mount 

St. Helens, CVO issued occasional written statements concerning volcanic activity and posted these to its Web site. 

Such statements were posted daily in the fall of 2004 and 

were supplemented by media advisories, information about 

news briefings, and notification of events released by fax and 

e-mail. In October 2004, CVO broadened daily update dissemination by e-mail to include media and aviation contacts; 

previously, e-mailed updates were sent primarily to government agencies. Broad update distribution has continued 

throughout 2006.

Hazard Information Management 

Through October 2006

Since cessation of the JIC on October 13, 2004, staff 

at CVO has replied to about 800 phone inquiries from the 

media, principally about the status of Mount St. Helens. 

As of October 2006, the combined number of media phone 

inquiries received by CVO and by the JIC during its 11 days 

of operation in 2004 exceeded 2,000. In addition to these 

general media inquiries, representatives of 41 documentary 

film projects contacted USGS–CVO for interviews, information, and graphics, mostly about the ongoing and 1980 

eruptions at Mount St. Helens. Each year, the CVO staff has 

provided additional media briefings on the anniversary of the 

2004 eruption and after significant explosive events. USGS 

staff assembled 10 rolls of video for use by media (b-rolls) 

between November 2004 and October 2006. As of this 

writing, the level of media interest is variable. The number 

of media inquiries rises with sightings of rockfall, with its 

subsequent suspension of dust and ash particles in the air, 

and with the onset of cold and clear weather, which initiates 

visible condensation plumes above the crater rim. CVO staff 

has taken the time to educate local media about these two 

phenomena. It is commonly accepted that this education has 

reduced the number of media inquiries considerably.

Challenges and Lessons Learned

The USGS and its partner agencies in emergency 

management faced a number of challenges in this episode 

of unrest at Mount St. Helens. Here, we present the most 

important challenges and lessons into four groupings: media 

and message management, JIC management, special needs of 

university cooperators, and organizational constraints. Many 

of these observations were compiled at an interagency afteraction review of the eruption response, an important step in 

identifying lessons learned following any eruption crisis.

Media and Message Management

Pre-Event Planning for Communications with 

Media

A pre-event hazard communication and media management plan reduces the inevitable scramble to establish an effective information pipeline to the public and other constituencies 

(Peterson and Tilling, 1993; Newhall and Punongbayan, 1996). 

The plan should be written by agency representatives most 

likely to be in the “hot seat” during an event. Volcano observatories and cooperating scientific agencies need full-time scientists 

whose major roles are to manage media relations, community 

outreach, and education. These scientists should achieve a working understanding of emergency roles codified in NIMS, including their role in a JIC. The roles should address the following 

responsibilities:

1. Engage the media and educate them about hazards and agency 

roles in hazard response.

2. Prepare background information about a volcano and its hazards 

before a crisis begins and arrange for rapid development and dissemination of updated information, graphics, and maps required 

for briefings.

3. Maintain relations with media representatives so that they are 

prepared to reach broad audiences with agency messages. Media 

representatives should feel comfortable asking questions so that 

they can “get the story right.”

4. Establish, exercise, and review phone, e-mail, fax, and other 

crisis-communication protocols.

5. Identify scientists who can dedicate their time to communication 

needs and offer media training for key staff.

6. Establish a system for obtaining photography and video footage 

for inclusion in media b-rolls.

7. Acquire official agency logos and apparel.

8. Consider a system for tracking media inquiries and workload.

9. Isolate public Web sites, by serving them from computer systems independent of the computers used for processing scientific 

data, to reduce the potential that increasing demand will crash 

either system.

Tools for First Moments Following an Event

After months or years of volcanic quiet, an eruption or 

onset of dramatic unrest can be a chaotic time with respect 
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 515

to media-communication protocols. Background fact sheets, 

FAQs, and other pre-event resources can help fill the immediate need for information. CVO developed a “Volcano Rapid 

Response Reference Page” that lists steps to be taken within the 

first few minutes of a visible volcanic event (fig. 8). The Reference Page can be customized for eruptions at any volcano.

Fighting Imaginations and Misperceptions

Agencies must devise a system to prepare, review, and 

update talking points that reinforce appropriate messages and terminology. A recurring challenge during the 2004 volcanic unrest 

TASKS - First 10 MINUTES of volcanic event 

ARRANGEMENTS - During first HOUR of volcanic event 

Include contact information for additional help within office, within team and communication 

offices, and with partner agencies. 

Volcano Rapid Response Reference Page 

1- Before Speaking To Media: 

Visit operations room—obtain three known facts about situation for the development of 

 talking points: 

1- News Briefing Arrangements: Establish time and place; announce the briefing; arrange 

 for audio bridge; choose speakers; develop messages to greater depth and information; 

 assemble agenda; address needs for graphics. 

2- Front Desk Duty: Arrange for staffing in evenings and early mornings as necessary. 

3- Joint Information Center Planning: Initiate interagency arrangements as necessary. 

6- B-Roll: Encourage scientific observers to take video that can be used for media b-roll; 

 arrange for b-roll preparation and distribution. 

7- Web Site Information: Provide any necessary information about media briefings and 

 b-roll to webmaster for b-roll preparation and distribution. 

4 - Review Personal Needs For Outreach Staff: Food, family schedules, for example. 

5- Off-Hours Arrangements: Advise duty scientists to check the front desk recorder for 

 phone messages and respond to inquiries during nonwork hours. If inquiries require 

 additional attention, contact outreach staff. s t n i o P g n i k l a T 

What Is Happening? 

What Is The Impact? (Example - Ash from plume will fall east of volcano). 

What Are We Doing About It ? (Example - USGS closely monitoring this event 

and advising local officials). 

2- Update Talking Points: Add some situation background. 

4- Recorded Message: Place temporary messages on pertinent phone recorders, 

 information lines, and Web sites. 

3- Arrange Additional Help: For liaison between operations room and outreach staff. 

5- Answer Inquiries: Keep talking points visible while providing telephone interviews; 

 maintain a record of inquiries. 

6- Request Assistance From Other Offices/Agencies: If it appears that a media event 

 might overwhelm staff on hand. 

Contact Information for Outreach Staff: 

(Example - At 1725 PDT a small explosive eruption began). 

Figure 8. Volcano Rapid Response Reference Page illustrating how the Cascades 

Volcano Observatory addresses media inquiries during sudden volcanic events, such 

as steam-and-ash explosions and resultant plumes.

concerned the widespread perception that any activity at Mount 

St. Helens was a prelude to another cataclysmic eruption resembling the event on May 18, 1980. Dispelling this image was a 

challenging exercise in wording and repetition. Once lodged in 

the public’s mind, such misperceptions are hard to remove.

Disseminating Updates to Remote Information 

Centers

Maintaining contact, continuity, consistency, and completeness of message with remote information centers is both a 

technical and a human challenge. Communication and sharing 
516 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

of talking points with PIOs and information scientists at Castle 

Lake Viewpoint were hampered by technological difficulties (fax 

and phone line service between information centers was intermittent) and the Viewpoint’s physical distance from the JIC.

The Media’s Need for Information, Now!

Agency representatives must be prepared to address 

nearly continuous information demands from the media for 

the 24-hour international news cycle. Scientists must create a 

steady stream of updated information to satisfy the appetite of 

media outlets and their constituencies. During the Mount St. 

Helens unrest, media photographers found it difficult to get 

close-range photographs because of temporary flight restrictions. At the request of the media, the 10 b-rolls released by 

the USGS provided close-up video coverage of the eruption, 

volcanologists working in the crater, and time-lapse and 

thermal-image video. Photographs taken by USGS observers 

were posted on the CVO Web site within a day of their being 

taken. There was a constant need for charts and illustrations 

for news briefings, especially for diagrams of volcano cross 

sections and schematics showing the relative sizes of the lava 

dome through time. Any media management plan should cover 

how these graphical products will be distributed efficiently.

Near-Simultaneous Observations by USGS Staff 

and Media

Scientists must often react to near-simultaneous observations by scientific staff and media. Today’s media outlets have 

access to and budgets for fly-bys of the volcano by helicopters, which often are outfitted with modern infrared sensors 

and cameras similar to those used by scientists. At Mount St. 

Helens, there was pressure to interpret observations made by 

media representatives before the extended group of monitoring 

scientists had seen or analyzed the information.

Maintaining Appropriate Distance Between 

Media and Operational Staff

Maintaining appropriate distance between media representatives and operational staff is a primary reason for positioning 

the JIC away from science facilities. Prior to the 2004 volcanic 

unrest, many media members had developed links with scientists and contacted them at will. The creation of the informationscientist role freed operational scientists who had responsibilities for monitoring and analysis. The founding of the JIC and its 

media briefings allowed controlled access to scientists.

Consistent and Careful Use of Nomenclature

Another important concern in sharing information with the 

media and public is consistent use of terminology and nomenclature. In one particular case, USGS scientists unwittingly fed 

public concern through their casual use of geographical terms. 

Constant repetition of the “bulging crater floor...on the south 

side...of the old lava dome” gave rise to a mistaken notion that 

the south flank of the entire volcano was bulging in a manner 

similar to the north flank prior to the catastrophic May 18, 1980, 

event. Another source of terminology confusion concerned 

use of the word “eruption.” For example, there were references 

to the events of early October 2004 as explosions, emissions, 

and eruptions. Although such distinctions appear trivial, and in 

many senses are purely semantic, to the reporting media representatives, these are important facts to get right.

The word “eruption” has a fairly specific image to many, 

and that includes visible lava, ash, and activity that is full of 

motion and potential danger. At this writing in October 2006, 

many people are unaware that Mount St. Helens continues 

to erupt. The same phenomenon occurred in Alaska during 

the 2005–6 eruption of Augustine Volcano. Once the volcano 

ceased ash-cloud production and began quietly producing 

lava flows, many people assumed the eruption had stopped 

altogether, despite the local volcano observatory’s constant 

repetition of messages that the eruption continued.

Joint Information Center Management

Development and Concept

Agencies must recognize the value of a JIC and then 

plan for it. At Mount St. Helens there was a hesitancy to form 

a JIC, caused in part by immediate needs but in greater part 

by not understanding a JIC’s value to all involved agencies. 

Volcano-response plans for each Cascade Range volcano 

refer to the use of a JIC, but, at the time of the 2004 event, no 

detailed plan existed for rapid development and staffing. It is 

important to understand the basics of the NIMS, each agency’s 

role during a crisis, and how a JIC may help.

Adequate Space for Necessities

A JIC requires large and small rooms with secure entry 

points. The Mount St. Helens JIC lacked distinct and secure 

entry points that were out-of-view of JIC operations. Access to 

the Media Briefing Room required passage of media representatives through the inner workings of the JIC, which exposed 

sensitive information not intended for public release, such as 

private phone numbers. This required JIC workers temporarily to “sanitize” whiteboards of sensitive information. Such 

security issues can be addressed with good JIC design. The 

Mount St. Helens JIC suffered from a lack of small rooms for 

individual media interviews.

Staffing and Logistics

Agencies must plan for the details of assembling people 

and technological facilities rapidly during establishment and 

maintenance of a JIC. At the Mount St. Helens JIC, staff from 
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 517

the Clark Regional Emergency Services Agency took the initiative to locate call center personnel by announcing the need 

on an extensive PIO listserve in the Portland-Vancouver area. 

Many people responded to the call for help, but maintaining a 

refreshed staff of trained and knowledgeable call takers, technical specialists, and JIC managers was a constant challenge. 

Also of value was participation of volunteers from the City of 

Vancouver who could assist traveling media with requests for 

general local information. A JIC handbook of procedures and 

resources was extremely helpful for educating existing and 

newer personnel. Lines of authority were a subject of frequent 

discussion, with personnel adjusting to the roles placed upon 

them within the JIC. Security firewalls presented continuous 

technical roadblocks for communications, both within the JIC 

and to outside agencies. A general e-mail account could not be 

established, so staff had to use individual e-mail accounts for 

transmitting JIC business information.

Maintaining Consistent and Timely Messages

Agencies must consider that maintenance of a consistent, 

current message requires vigilance. At Mount St. Helens, volcanic events frequently eclipsed the Daily Update and necessitated hurried development of updated talking points, education 

of call takers and staff at other information outlets, and rumor 

control. An official JIC Daily Update required review by a 

representative of each agency. The duration of the production 

process often outlived the usefulness of the product. Early 

morning requests from East Coast media were most challenging because daily updates were not yet available. News briefings came too late for local morning news shows. There was 

demand for hourly updates, but no official process was available to provide this need. As a result of such media demands, 

the Alaska Volcano Observatory provided hourly status reports 

on its Web page during the height of the eruption of Augustine 

Volcano in early 2006. These were informal and nonreviewed 

snapshots of what was happening at the volcano, and they 

served to underscore that scientists were actively watching the 

volcano. Even remarking that nothing had changed in many 

hours apparently served to reassure Web users that the Observatory was aware of current conditions.

Special Needs of University Cooperators: 

Pacific Northwest Seismic Network

University cooperators must find the resources to protect 

monitoring capabilities while addressing media response. The 

greatest challenge faced by the PNSN was to find ways to 

satisfy the numerous requests for interviews while maintaining 

network operations and analyzing and interpreting data.

The PNSN PIO addressed media needs with support from 

student staff, who answered phones and organized requests for 

information and interviews. The PIO reviewed these requests 

and PNSN scientists monitoring the eruption were scheduled, 

when appropriate, to participate in interviews. The PNSN also 

received assistance from the UW Office of News and Information who assisted in scheduling interviews and disseminating 

press releases. At key moments, such as a change in alert level 

or following steam and ash explosions, reporters were pooled 

and the PNSN Director made a statement, answered questions for a few minutes, and then returned to work. National 

news outlets did not always get to talk to the scientist they 

requested, but all media requests were addressed. The PNSN 

concentrated their response onsite at the UW and turned down 

almost all requests for television-studio interviews. The PNSN 

also frequently referred many requests to the JIC, particularly 

those not directly related to seismicity.

Organizational Constraints

Staffing

The USGS volcano observatory system consists primarily of technical specialists with defined scientific roles and 

expertise to monitor volcanoes and assess hazards. For most 

staff, outreach and interaction with the media are ancillary 

duties; therefore, technical tasks may go unfinished when 

scientific staff are consumed by providing information. Even 

for observatories with professional communication and publicinformation specialists, a single eruption crisis can quickly 

overwhelm slim resources. To respond effectively and not 

diminish the technical efficacy of the observatory, it is key to 

call quickly for reinforcements from within and outside the 

USGS. It is helpful to have predetermined lists of staff available for temporary short-term duties to assist in this capacity. 

Crisis-related staffing should include information scientists, 

Web and illustration staff, and information technology support.

Cost

Establishing a JIC is not cost free. According to the 

GPNF, the total JIC operation costs for 11 days, including 

extra staffing by interpreters at MSHNVM, was approximately 

$88,400 in salaries and $7,500 in equipment and supplies. 

Most of the personnel who worked at the JIC either donated 

their time, or their agencies absorbed salary costs as a contribution to the regional response. For the Call Center, where 40 

people rotated through various shifts to answer phones, the 

estimated cost of staff time alone totaled more than $17,000.

The State of Washington does not have a system to pay 

for JIC support without declarations of emergency or disaster, 

a situation common in other states. Once such a declaration 

exists, the State can request Federal resources from FEMA 

to support staff costs. Cultivating partnerships in advance, 

including the development of formal mutual-aid agreements, 

can help local offices deal with the staffing shortage during 

an information crisis. Although some agencies or jurisdictions were unable or unwilling to commit paid staff time to 

volunteer at the Mount St. Helens JIC, many others did donate 

personnel to the public-information effort. This volunteer 
518 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

opportunity provided PIOs with skill-building advantages, 

because it is rare to get the chance to work within a JIC.

Another expense of JIC operation is technological outfitting of the facility. The Mount St. Helens JIC at the GPNF 

headquarters already had phone lines and extensions available, 

so the cost of installing a switchboard and other phone lines 

was minimal. However, it was necessary to install fiber-optic 

cable to serve the needs of the media satellite trucks, and 

this involved some major contract work. By pre-identifying 

a facility with critical infrastructure in place and securing an 

agreement in advance for free or low-cost emergency use of 

that facility, technical costs can be minimized.

Media Training Needs

Scientists rarely, if ever, are trained to work with the 

media. For those who are likely to be directly involved in providing information to the media, many resources are available 

to provide an appropriate level of training, including classes in 

basic crisis communication available through local emergencymanagement offices.

Conclusions

Preeruption planning contributed significantly to the 

readiness of the USGS, PNSN, GPNF, and other key agencies 

to deliver timely and effective hazard information about the 

evolving eruption at Mount St. Helens, one of the world’s most 

famous volcanoes. The success of the response speaks well of 

decisions by the USGS, a largely scientific agency, to enable 

communication planning to proceed on par with scientific 

response planning. Training of media representatives over the 

long term; practicing response plans; and ongoing communication among scientists, emergency managers, and community 

leaders about volcano hazards in the Pacific Northwest brings a 

cushion of support that bolstered effectiveness of the response. 

Despite this preparation, considerable flexibility, creativity, 

and rapid development of strategies to deal with unanticipated 

issues were essential. In particular, today’s widespread use of 

the Internet and the around-the-clock news cycle required a fast, 

sustained pace of information delivery that frequently pushed 

the limits of staffing and internal communications. Establishing a JIC within the ICS structure contributed substantially to 

managing these expectations. The JIC also insulated the JOC 

and CVO and PNSN scientists from media attention, freeing 

managers and scientists to deal with critical aspects of the crisis.

To our knowledge, this was the first time that a fully 

developed JIC was used to address volcanic unrest. Earlier 

volcanic responses have tended toward multiple agencies 

providing information independently by specialists, such as at 

Ruapehu volcano, New Zealand (David Johnston, oral commun., 2005) and Volcán Santa Ana, El Salvador (John Ewert, 

oral commun., 2005). An exception was a single source at 

Pinatubo volcano, Philippines (Chris Newhall, oral commun., 

2005). Our 2004 experience at Mount St. Helens confirms that 

volcano and seismic observatories benefit from the availability 

of a full-time scientist focusing on the needs of media relations, community outreach, and education.

Acknowledgments

We thank all those who contributed to successfully managing the hazard message during the 2004 unrest and eruption. 

We also acknowledge the critical feedback and constructive 

criticism from our partners in the media, at other agencies, and 

our own colleagues, all of whom helped improve our communication with the public. This manuscript was reviewed by 

Maggie Mangan and Chris Jonientz-Trisler.

References Cited

Dzurisin, D., Vallance, J.W., Gerlach, T.M., Moran, S.C., and 

Malone, S.D., 2005, Mount St. Helens reawakens: Eos (American Geophysical Union Transactions), v. 86, no. 3, p. 25, 29.

Federal Emergency Management Agency, 2004, National Incident Management System: U.S. Department of Homeland 

Security, 139 p. [http://www.fema.gov/emergency/nims/

index.shtm, last accessed November 17, 2006].

Frenzen, P.M., and Mataresse, M.T., 2008, Managing public 

and media response to a reawakening volcano; lessons from 

the 2004 eruptive activity of Mount St. Helens, chap. 23 of

Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).

Gifford Pinchot National Forest, 2003, Mount St. Helens 

volcanic activity response plan update: Vancouver, Wash., 

Gifford Pinchot National Forest, 42 p.

Moran, S.C., Malone, S.D., Qamar, A.I., Thelen, W.A., 

Wright, A.K., and Caplan-Auerbach, J., 2008a, Seismicity associated with renewed dome building at Mount St. 

Helens, 2004–2005, chap. 2 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Moran, S.C., McChesney, P.J., and Lockhart, A.B., 2008b, 

Seismicity and infrasound associated with explosions at 

Mount St. Helens, 2004–2005, chap. 6 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).
24. Hazard Information Management During the Autumn 2004 Reawakening of Mount St. Helens Volcano 519

Newhall, C.G., and Punongbayan, R.S., 1996. Successful 

volcanic-risk mitigation, in Scarpa, R., and Tilling, R.I., 

eds., Monitoring and mitigation of volcano hazards: Berlin, 

Springer-Verlag, p. 807–838.

Peterson, D.W., and Tilling, R.I., 1993, Interactions between 

scientists, civil authorities, and the public at hazardous volcanoes, in Kilburn, C.R.J., and Luongo, G., eds., Monitoring active lavas: London, UCL Press, p. 339–365.

Qamar, A.I., Malone, S.D., Moran, S.C., Steele, W.P., and 

Thelen, W.A., 2008, Near-real-time information products 

for Mount St. Helens—tracking the ongoing eruption, chap. 

3 of Sherrod, D.R., Scott, W.E., and Stauffer, P.H., eds., A 

volcano rekindled; the renewed eruption of Mount St. Helens, 2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).

Scott, W.E., Sherrod, D.R., and Gardner, C.A., 2008, Overview of the 2004 to 2005, and continuing, eruption of 

Mount St. Helens, Washington, chap. 1 of Sherrod, D.R., 

Scott, W.E., and Stauffer, P.H., eds., A volcano rekindled; 

the renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Wolfe, E.W., and Pierson, T.C., 1995, Volcanic-hazard zonation for Mount St. Helens, Washington: U.S. Geological 

Survey Open-File Report 95–497, 12 p.
520 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 521

Tracking the flow of gases from magma is perhaps the most vexing task 

of physical volcanology. To play upon a phrase, “gas waits for no man,” 

because the gaseous species are volatile. Also, things get in the way, not the 

least of which are the volcanic edifice itself and ground water that permeates 

it.

The 2004–6 Mount St. Helens dacite was a “flat” magma—greatly 

depleted in excess (exsolved) volatiles compared to the May 18, 1980, dacite. If new magma entered the system as part of the 2004 eruption, it must 

have been gas poor, on the basis of modeling constrained by cumulative CO2

emissions, cumulative dacite production, and measurement of dissolved H2

O 

and other volatile concentrations within the dacite glass.

This view, that the eruption was driven by degassed magma, is supported by chemical trends in hot springs and creek waters that drain the 

Mount St. Helens crater. Since 1994 these waters have shown decreasing 

SO4

 and Cl concentrations, most likely reflecting changing release rates of 

sulfur gases, HCl, and CO2

 from the magma and, to a varying degree, the 

efficiency of gas scrubbing by the water. Still unanswered is the question 

of what triggered the eruption of such gas-poor magma—was it influx of a 

small amount of hotter, gas-rich magma that was otherwise compositionally 

similar to the 1980s magma, infiltration of rainwater into the upper part of 

the conduit system, or some other mechanism?

The lessons from the 2004 eruption are clear: Scrubbing or simply having a magma already depleted in gases can mask the more typical signal 

of shallow magmatic sulfur-species degassing, so early monitoring of CO2

should accompany early SO2

 monitoring. Indeed, the monitoring of CO2, 

SO2

, and H2S can track the drying out of a volcano and help distinguish the 

effects of gas scrubbing from the loss of permeability when pathways for 

gas are sealed by precipitates.

This eruption also provided an opportunity to further test new gas-emission monitoring tools. Results from open-path Fourier-transform infrared 

spectroscopy support the interpretation that closed-system degassing occurs 

as shallow as 1–2 km and that open-system degassing characterizes the shallowest part of the magmatic system.

Volcanic Emissions 
522 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

View to south-southwest on September 19, 2005, as spine 6 grew westward 

away from disintegrating spines 4 and 5 (left of center) and compressed the 

highly crevassed west arm of Crater Glacier. Rocky terrain in lower right is 1980s 

dome. USGS photo by J.W. Ewert.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 25

1

 U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Pre- and Post-Eruptive Investigations of Gas and Water 

Samples from Mount St. Helens, Washington, 2002 to 2005

By Deborah Bergfeld1, William C. Evans1, Kenneth A. McGee2, and Kurt R. Spicer2

Abstract

Samples of gas and water from thermal springs in 

Loowit and Step canyons and creeks that drain the crater at 

Mount St. Helens have been collected since October 2004 

to monitor the flux of dissolved magmatic volatiles in the 

hydrologic system. The changing composition of the waters 

highlights a trend that began as early as 1994 and includes 

decreasing SO4 and Cl concentrations and large increases in 

HCO3. Geochemical models indicate that mineral sources and 

sinks are not the main controls on the changing water chemistry, and carbon and helium isotopes indicate that their sources 

in the gases and waters have remained unchanged during 

this time. The present-day molar ratios of C, S, and Cl in the 

springs approximate ratios measured in plume emissions in 

August 2005 and provide supporting evidence that changes 

in water chemistry most likely reflect changes in the release 

rates of sulfur gases, HCl, and CO2 from the magma and a 

varying degree of efficiency of gas scrubbing by the overlying 

water. Results from coupled chemical analyses and discharge 

measurements on the creeks yield an estimate of the dissolved 

flux of magmatic HCl, SO2

, and CO2 of around 5.2, 4.7, and 

22 metric tons per day, respectively.

Introduction

Airborne gas measurements were a valuable tool in 

tracking magmatic emissions for Mount St. Helens during 

the 1980–86 eruptive sequence, but flights eventually were 

suspended as emission rates dropped to low levels. Between 

September 1988 and September 2004, attempts to measure the 

flux of magmatic volatiles at Mount St. Helens were restricted 

to analyses of gases and condensates from vents on the lava 

dome and thermal waters from springs in Loowit and Step 

canyons3

 in The Breach. Results from those analyses generally have shown a progressive decrease in magmatic volatile 

concentrations with time. Airborne gas-flux measurements 

over the crater at Mount St. Helens were resumed in September 2004 following the onset of seismic activity (Doukas 

and others, 2005). To date, those measurements have shown 

that plume CO2 and SO2 emission rates are much lower than 

emission rates during the 1980s eruptions (Gerlach and others, 

2005, and this volume, chap. 26). Plume emission rates were 

one line of evidence that supported an early hypothesis that 

degassed magma was driving the current eruption.

To track the flux of dissolved magmatic volatiles in the 

hydrologic system during the current eruption, we established 

a surface-water sampling program on October 13, 2004. The 

first samples were collected on the Pumice Plain from two 

creeks that drain the crater. In mid-April 2005, as conditions 

nearer the crater stabilized, we expanded the monitoring network to include three locations in The Breach (fig. 1). These 

five sites were the main focus of our monitoring campaign 

during the first year after the 2004–5 eruption, but at times we 

also have collected samples from a warm spring on the west 

end of the Pumice Plain and from hot springs in Loowit and 

Step canyons in upper parts of The Breach. Thus far, our sample collection intervals have been irregular, but since October 

2005 some continuous temperature and conductivity data have 

been collected by a probe installed near the top of Step creek.

In this report we present geochemical data from the 

monitoring sites; data from field campaigns in the crater and 

3

 Many geographic features in the Mount St. Helens area are referred to 

informally by Earth scientists. Capitalization of “canyon,” “creek,” or “spring” 

indicates names formally adopted and listed in the Geographic Names 

Information System, a database maintained by the U.S. Board on Geographic 

Names. Noncapitalized names are applied informally—eds.
524 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The Breach in 2002 and 2005, which focused on the chemistry 

of crater gases and hot spring waters; and data from a survey 

of regional cold springs. We also include data from previous 

investigations of the geochemistry of the hot springs at Mount 

St. Helens in order to evaluate long-term changes in gas 

sources or fluxes and, ultimately, to determine how magmatic 

processes at an active volcano are reflected in the hydrothermal system. We have adopted the names of the Loowit canyon 

hot springs that were used by Shevenell and Goff (1993) but 

have adhered to the policies of the U.S. Board on Geographic 

Names regarding capitalization of formal and informal names. 

Additional chemical data not discussed in the paper can be 

found in appendix 1, which appears only in the digital versions 

Gas sample

Hot-spring sample

Cold/warm-spring sample

Monitoring site

Creek sample

19

1,2

17,18

14-16

21A,B 25

24

9-11

3-8

20

23

22

28

27

26

12,13

2

1

13

7

8

6

4 5

3

10 17

18

14

16

12

15

11 9

The 

Breach

Pumice Plain

Loowit 

Creek

Step

creek

560000 570000

5120000

5110000

0 3

KILOMETERS 

Figure 1. Maps showing sample locations for this study, 2002–2005, Mount St. Helens, Washington. Numbers for sample sites 

correspond to numbers listed in table 1. Inset shows crater and The Breach sample locations in more detail. Datum for geographic 

grid and UTM coordinates referenced to WGS84, zone 10.

of this work (in the DVD that accompanies the printed volume 

and online at http://pubs.usgs.gov/pp/1750).

Background

After the May 1980 eruption of Mount St. Helens, gas 

samples were collected from cracks and fumaroles near the 

dome in the fall of 1980 (Evans and others, 1981; Gerlach 

and Casadevall, 1986). Due to dome growth and collapse and 

extrusions of new lava, the gas-sampling locations varied 

during the subsequent years. Before our work in 2002, the last 
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 525

gas sample from the crater was collected in June 1998 from an 

86°C vent on the southwest side of the 1980–86 lava dome on 

a flow called the September 1984 lobe, herein referred to more 

briefly as “September lobe.” The sampling site was formerly 

the hottest fumarole on the dome, but in 1998 it had cooled 

by 260–370°C relative to temperatures recorded in 1994–95 

(Symonds and others, 2001).

Downcutting of avalanche deposits in The Breach of 

Mount St. Helens created Loowit and Step canyons, which 

drain to the north and represent the main discharge of surface water from the crater (fig. 1). Thermal springs were first 

observed discharging into Loowit Creek during the summer 

of 1983, and the first sample of spring water was collected in 

October 1983 (Thompson, 1990). Lower-temperature, low-discharge springs in Step canyon were discovered at a later date, 

and the first published analyses of the Step canyon springs are 

from 1988 (Shevenell and Goff, 1993). It has been estimated 

that more than 95 percent of the hot spring discharge from 

both canyons is provided by the springs in Loowit canyon 

(Shevenell and Goff, 1993). During the 1980s to mid-1990s, 

the hot springs in Loowit canyon were located in three areas 

along the length of the canyon. Each area contained clusters 

of hot springs of varying size and temperature. From south 

to north, the spring groups are called (1) Loowit source, (2) 

Loowit boulder basin, and (3) Loowit travertine. The name 

Loowit travertine is a misnomer that stems from the presence 

of gypsum-bearing precipitates on the rocks around the vent 

orifices (Shevenell and Goff, 1993). There are no names for 

the springs in Step canyon. The waters from both Loowit and 

Step canyons were sampled throughout the 1980s and 1990s, 

but because of the larger volume of thermal water, the Loowit 

canyon springs were sampled with the greatest frequency.

Post-2000 Field Campaigns

Gas and Water Samples from the Crater and The 

Breach

In August 2002 there were no fumaroles in the crater at 

Mount St. Helens, but diffuse steam was rising from several 

locations on the 1980–86 dome. We sampled gas at two locations from the strongest vents we could find on September 

lobe (fig. 1; table 1). Because of safety concerns we were 

unable to resample the sites in 2005. In August 2002 and July 

2005 we also collected gases from bubbling hot springs in 

Loowit canyon. To our knowledge these are the first free-gas 

samples collected from springs in Loowit canyon.

Water samples were collected from the Loowit canyon hot 

springs in 2002 and 2005, and samples of the Step canyon 

springs were collected in July and October 2005. At springs 

with multiple orifices, our strategy was to choose the hottest 

spring for sample collection, and at times we also sampled 

water from a nearby spring. In 2002, the hottest vents at the 

Loowit boulder basin and Loowit travertine sites were along 

the west bank of Loowit Creek. The outflow from the southernmost spring, which we assume to be the Loowit source spring, 

was from a single vent on a bench above the canyon floor. The 

water cascaded down the west canyon wall into Loowit Creek. 

In July 2005 there was no visible outflow from the Loowit 

source spring; instead, a new large spring issued from the base 

of a small fan that abuts the west wall of the canyon. Additional 

cooler waters discharged from an upper part of the fan and 

small vents along the west bank of the creek.

The thermal waters in Step canyon occur in clusters and 

line the banks of the headwaters of Step creek. In 2005 there 

were no high-discharge vents, and all of the waters issued 

from areas composed of low-discharge seeps. Intermittent 

degassing was observed at both of the sites in July, but no gas 

was visible in October 2005.

Creeks and Regional Springs

The five sites that make up the monitoring network are 

along creeks that provide the major surface drainage from the 

crater at Mount St. Helens. The network includes two sites 

(LCBF and SCBF) on the Pumice Plain (fig. 1; table 1) and 

three sites (LCAF, EFSCAF, and WFSCAF) closer to the 

crater above the waterfalls on Loowit and Step creeks. The 

Pumice Plain sites are the most accessible and serve to monitor the surface drainage from locations just above the confluence of the two creeks. Of the higher-altitude sites, LCAF 

receives input from all of the hot springs in Loowit canyon and 

variable seasonal input of glacial meltwater; EFSCAF is on the 

east fork of Step creek, which is the fork that contains the Step 

canyon springs; and WFSCAF is a tributary to Step creek that 

drains the west wall of the crater. This channel (WFSCAF) 

carries seasonal runoff from the west arm of the Crater Glacier. We have not walked the length of this creek, but there are 

no obvious thermal waters that discharge to this drainage. In 

this report we also present water data from the probe locations 

(fig. 1; table 1) and three other locations on Loowit Creek, 

which were sampled in 2002 and 2005.

Several studies have used carbon and helium isotopes, 

together with water chemistry, to demonstrate that cold 

ground waters in volcanic regions can carry large amounts of 

dissolved magmatic carbon (Rose and Davisson, 1996; Chiodini and others, 2000; Evans and others, 2002). To investigate 

outflow of magmatic volatiles from the volcano to regional 

springs, we sampled eight mostly low-temperature springs in 

the area surrounding Mount St. Helens in July 2005. Kalama 

Spring, located southwest of Mount St. Helens, has the largest 

discharge of any spring in our sample set, followed by carbonate spring, a warm spring on the west end of the Pumice 

Plain (fig. 1; tables 1 and 2). Because the chemistry of carbonate spring has some similarities with the springs from The 

Breach, future discussion of this spring will be included with 

discussions of the hot springs. Of the seven other cold springs 

in this report, the most well-studied are Kalama, Willow, and 

Moss Springs, which have been sampled infrequently since 
526 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Location Map 

No. Northing Easting Altitude (m)

Old dome Mount St. Helens

September lobe, landing site 1 5116595 562436 2,122

September lobe, upper dome 2 5116600 562487 2,130

Springs in Loowit canyon

2005 degassing spring 3 5118045 562819 1,664

Loowit source spring 4 5118136 562831 1,635

New spring on fan of Loowit canyon 5 5118157 562856 1,617

Loowit boulder basin spring 6 5118275 562882 1,604

Loowit boulder basin degassing spring 7 5118273 562891 1,604

Loowit travertine spring 8 5118442 562921 1,584

Springs in Step canyon

Step canyon hottest spring 14 5118396 562453 1,644

Step canyon source spring 16 5118324 562453 1,645

Monitoring network sites

Probe site, Loowit Creek above falls 9 5119128 562929 1,507

Loowit Creek above falls 11 5119198 562981 1,509

Probe site, east fork Step creek above falls 15 5118401 562446 1,617

West fork Step creek above falls 17 5119212 562519 1,512

East fork Step creek above falls 18 5119224 562546 1,512

Loowit Creek below falls 21A 5121306 562300 1,219

Step creek below falls 21B 5121306 562300 1,219

Other creek locations

Lower Loowit Creek 10 5119150 562935 1,509

Upper Loowit Creek 12 5117947 562819 1,662

Upper Loowit Creek 13 5117784 562886 1,709

Pumice Plain spring

Carbonate spring at gage 19 5122487 559926 935

Carbonate spring 20 5122421 560010 945

Regional cold springs

Springs west of Cold Spring Creek 22 5115219 557268 1,132

Kalama Spring 23 5110368 557390 832

Spring on Loowit trail 24 5119654 564283 1,330

Willow spring 25 5121399 564292 1,125

Spring on Pine Creek 26 5109610 569886 731

Moss spring 27 5112493 569935 882

Spring west of Swift Creek 28 5109062 563021 756

Table 1. Map number and location coordinates for samples collected during 2002–2005, 

Mount St. Helens, Washington.

[UTM coordinates and altitude are referenced to the WGS84 datum.]
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 527

the late 1980s. Four of the springs are unnamed and were 

identified on 24,000-scale quadrangle maps or were brought 

to our attention by Mike Clynne of the U.S. Geological Survey (USGS), who has mapped extensively in the area around 

Mount St. Helens. In the text that follows we refer to the 

regional springs by their associated map number (fig. 1; table 

1). The eight springs we report on here are likely the largest 

on the margins of the volcano, although assuredly more cold 

springs are in the Mount St. Helens area.

Field and Laboratory Methods

Water and gas samples were collected for a suite of 

chemical analyses using standard field and laboratory techniques. At all sites the water temperature and conductivity 

were measured in the field using portable meters. Spring-water 

samples collected for bulk chemistry were preserved in the 

field by filtering them through a 0.45-μm filter into plastic 

bottles that were first rinsed with filtered water. Samples for 

cation analyses were preserved with high-purity nitric acid by 

dropwise addition to a pH < 2. Other nonfiltered (raw) springwater samples were collected in glass bottles for alkalinity 

and δD and δ18O analyses and in preevacuated glass bottles 

for δ13C analysis of dissolved inorganic carbon (DIC), which 

consists of dissolved CO2 + HCO3 + CO3. Depending on the 

sample size, waters collected for δ13C–DIC were acidified with 

0.5 to 1.0 mL of high-purity hydrochloric acid to convert DIC 

to CO2. Values for δ13C–DIC and 14C–DIC were determined 

on purified CO2 extracted from the water samples following 

methods described in Evans and others (2002).

Samples of creek waters from the monitoring sites 

typically were collected into prerinsed plastic bottles as raw 

waters. Filtering, sample preservation, and pH measurements were performed in the laboratory, at Menlo Park, Calif. 

Because samples from the monitoring sites were sometimes 

stored before they were received at the lab, we have not 

determined the bicarbonate concentrations of many of these 

water samples. A few samples of creek waters were collected 

for δ13C analysis of DIC as described above. Creek discharge 

rate was measured using a pygmy meter. Visual estimates of 

discharge are identified in table 2.

Anion and cation concentrations in water samples 

were determined using ion chromatography and inductively 

coupled argon plasma spectrometry at USGS laboratories in 

Menlo Park, Calif. Isotope analyses were performed at USGS 

laboratories in Menlo Park, Calif., and Reston, Va., and at 

the Center for Accelerator Mass Spectrometry at Lawrence 

Livermore National Laboratory, Calif. Isotope analyses were 

performed using techniques outlined in Epstein and Mayeda 

(1953), Coplen (1973), Vogel and others (1987), and Coplen 

and others (1991).

Gas samples for bulk composition and δ13C–CO2

 analyses were collected into preevacuated glass bottles through a 

hollow titanium rod (from gas vents) or a funnel fitted with 

Table 2. Spring and creek temperature and discharge 

measurements 2003–2005, Mount St. Helens, Washington.

[Letters following discharge rate indicate field method: EV, width and 

depth measured, velocity estimated; E, visual estimate; M, measured. 

Some locations coded to specify sites: ps, probe site; LCAF, Loowit 

Creek above falls; LCBF, Loowit Creek below falls; EFSCAF, east fork 

Step creek above falls; WFSCAF, west fork Step creek above falls; 

SCBF, Step creek below falls; CARB gs, carbonate spring gage site.]
528 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Tygon tubing (from bubbling springs). Prior to sampling, the 

entire collection system was purged of atmospheric gases. 

Gas samples were analyzed for bulk composition at the USGS 

in Menlo Park, Calif., using gas chromatography methods 

reported in Evans and others (1981). Carbon dioxide for stable 

isotope analysis was separated from the bulk gas sample using 

standard cryogenic techniques on a vacuum line (Evans and 

others, 2002). The spring gas samples for 3He/4He analysis 

were collected from the funnel and tubing apparatus into 

copper tubing that was then sealed at both ends with refrigeration clamps. The 3

He/4He determination was run at Lawrence 

Berkeley National Laboratory in Berkeley, California, following methods outlined by Kennedy and van Soest (2006). 

Helium isotope ratios are reported as R/RA values, which represent the 3

He/4He ratio in the gas relative to the 3He/4He ratio 

in air. Values for R/RA >1 indicate that a percentage of the gas 

is derived from a mantle or magmatic source.

Results and Discussion

Chemistry of the Hot Springs, 1983 to 2005

The chemistry and temperature of the Loowit canyon hot 

springs changed rapidly after the springs appeared in 1983 

(Thompson, 1990; Shevenell and Goff, 1993, 1995). The highest water temperature recorded was 92.6°C from the Loowit 

travertine spring group in 1986, and by then, concentrations of 

many of the dissolved constituents in the Loowit springs had 

peaked (Shevenell and Goff, 1993). The subsequent declining temperatures and variable chemistry were attributed to 

a rapidly cooling hydrothermal system that had formed just 

after the 1980 eruption. Several important conclusions were 

reached by Shevenell and Goff (1993, 1995, 2000) and Goff 

and McMurtry (2000), including (1) that the Loowit canyon 

hot springs discharged mixtures of recent meteoric water and 

about 10 percent magmatic water, (2) that the underground 

residence time of the mixed water was too short to allow 

water-rock equilibration or reliable geothermometry, and (3) 

that Cl, SO4

, and HCO3 in the springs were mainly derived 

from magmatic volatiles.

More recent samples show that the chemistry of the hot 

springs at Mount St. Helens is still changing (F. Goff and the 

USGS, unpub. data, 1984–94), although water temperatures 

since 1990 appear to have stabilized (fig. 2). The early trend 

in anions from SO4

 toward Cl dominance stopped, and water 

samples from 1994, 2002, and 2005 show declines in both 

SO4 and Cl relative to HCO3 (table 3; fig. 3). The post-1989 

data for the Step canyon springs follow a similar trend toward 

HCO3

-dominated water.

Overall, most elements in the Loowit springs show the 

effects of dilution by meteoric water over time (fig. 4). The 

plots in figure 4 may understate the dilution effects because 

of the increased focus on sampling the hottest springs in 

2002 and 2005. Dilution effects also are seen in the δD and 

δ18O composition of the waters from 2002 and 2005, which 

plot on or near the world meteoric water line (table 4; fig. 

5). Previous δ18O values of Loowit hot spring waters through 

1994 were shifted by as much as 2 per mil to the right of the 

meteoric water line.

Simple dilution, however, does not fully explain all 

of the data. Concentrations of Na and SO4 have decreased 

since 1984, whereas Cl and B concentrations remained high 

throughout the mid- to late 1980s and declined in later years 

(fig. 4). Concentrations of Mg increased in the 2002 and 2005 

samples, and HCO3 concentrations showed a large increase to 

concentrations higher than was recorded previously.

A clear response to the renewal of activity in 2004 has 

not yet been recognized in the hot springs. Just prior to the 

onset of seismicity in 2004, the discharge and water temperature at a spring in Loowit canyon appeared to be higher than 

normal (J.S. Pallister, USGS, oral commun., 2004), but no 

measurements were made. However, any possible temperature 

increase was short lived, as comparisons between the 2002 and 

2005 samples from the hottest springs in Loowit canyon show 

little change in temperatures (table 3). Similarly, in July 2005 

the temperature measured at the hottest spring in Step canyon 

was substantially higher than temperatures reported from 1994 

(fig. 2), but three months later, in October, the temperature at 

the same location had declined 18°C, although the specific 

conductance had decreased only slightly. Because the flow 

from the Step canyon vents is low, this temperature decline is 

conceivably related to the change in surface air temperature 

and conductive cooling. The possible causes for rapid swings 

in temperature at the hot springs need further investigation, but 

at this point, it is difficult to attribute any temperature changes 

Figure 2. Change in maximum temperature and specific 

conductance of hot springs in The Breach of Mount St. 

Helens since 1984. Data from dates before 2002 are from 

Thompson (1990), Shevenell and Goff (1993), Goff and 

McMurtry (2000), and unpublished data from the 

U.S. Geological Survey.

Max. temperature, Loowit springs

Max. temperature, Step springs

Conductivity, Loowit springs

Conductivity, Step springs

10

0

20

30

40

50

60

70

80

90

100

1983 1987 1991 1995 1999 2003

0

1

2

3

4

5

6

7

Onset of 

seismic unrest

MILLISIEMENS PER CENTIMETER

TEMPERATURE, IN °C

DATE
Sample No. Date Map No. Temp. (°C) pH Cond (mS/cm) B Br Ca Cl F Fe HCO3 K Li MSprings in Loowit canyon MSH-02-01 08/16/02 4 58.9 6.4 2.6 4.9 0.56 55 310 2.3 3.5 727 44 0.74 24 

MSH-02-02 08/17/02 6 61.0 6.3 2.7 5.7 0.61 62 330 2.2 4.0 769 46 0.82 28 

MSH-02-03 08/17/02 7 62.3 6.3 2.8 5.9 0.61 68 340 2.2 3.9 844 46 0.83 28 

MSH-02-04 08/18/02 8 57.6 6.5 2.8 5.6 0.59 68 350 2.6 2.3 779 45 0.75 31 

MSH-05-02 07/26/05 3 59.1 -- 2.6 6.0 0.44 67 281 2.5 4.0 -- 47 0.94 36 

MSH-05-03 07/26/05 3 57.0 6.5 3.0 5.9 0.44 66 280 2.6 4.2 761 46 0.90 35 

MSH-05-04 07/26/05 5 61.9 6.7 3.4 6.7 0.48 69 291 2.4 4.3 807 48 1.01 37 

MSH-05-05 07/26/05 5 60.5 -- 2.7 6.5 0.48 72 288 2.4 3.3 -- 48 1.00 38 

Springs in Step canyon MSH-05-07 07/26/05 14 56.6 6.6 2.4 5.7 0.32 67 217 1.3 4.7 752 30 0.95 49 

MSH-05-06 07/26/05 16 45.1 6.6 1.9 4.3 0.24 73 171 1.2 7.7 706 22 0.74 50 

MSH-05-22 10/25/05 14 38.4 6.4 2.2 5.7 0.36 66 241 1.9 17 -- 21 0.84 57 

Monitoring network sites LCAF 042105 04/21/05 11 27.5 8.5 2.1 3.8 0.32 48 190 2.2 0.34 -- 30 0.57 27 

LCAF 063005 06/30/05 11 27.9 8.6 1.6 -- 0.24 -- 172 2.6 -- -- -- -- -- 

MSH-05-11 07/26/05 11 29.1 8.8 1.6 3.4 <0.003 54 172 2.2 1.1 487 27 0.52 24 

LCAF 082605 08/26/05 11 30.1 8.3 1.6 2.9 0.28 32 169 2.3 0.29 -- 24 0.46 20 

LCAF 101105 10/11/05 11 30.0 8.6 1.9 -- 0.32 -- 197 2.3 -- -- -- -- -- 

MSH-05-24 10/25/05 9 25.9 8.7 1.8 3.5 0.32 55 193 2.5 0.85 533 28 0.54 24 

LCBF 101304 10/13/04 21A 16.3 8.5 1.5 -- 0.24 -- 147 1.8 -- -- -- -- -- 

LCBF 102004 10/20/04 21A 10.2 7.8 1.6 2.9 0.30 15 151 2.2 <0.08 -- 21 0.45 16 

LCBF 120204 12/02/04 21A 8.4 8.5 1.8 -- 0.24 -- 167 2.0 -- -- -- -- -- 

LCBF 021005 02/10/05 21A 11.5 8.6 1.8 -- 0.24 -- 170 2.0 -- -- -- -- -- 

LCBF 042105 04/21/05 21A 16.9 8.7 1.9 3.7 0.32 31 180 2.1 0.48 -- 28 0.56 24 

LCBF 063005 06/30/05 21A 18.6 8.7 1.6 -- 0.28 -- 174 2.6 -- -- -- -- -- 

MSH-05-12 07/26/05 21A 22.5 8.9 1.6 3.5 0.26 45 167 2.2 0.47 -- 27 0.53 24 

LCBF 082605 08/26/05 21A 19.8 8.5 1.6 -- 0.28 -- 178 2.2 -- -- -- -- -- 

EFSCAF 042105 04/21/05 18 19.5 8.6 1.5 -- 0.32 -- 190 1.6 -- -- -- -- -- 

EFSCAF 063005 06/30/05 18 23.5 8.5 1.6 -- 0.24 -- 190 1.0 -- -- -- -- -- 

MSH-05-09 07/26/05 18 26.6 8.8 1.5 4.7 0.24 29 172 1.0 0.13 533 20 0.78 52 

EFSCAF 082605 08/26/05 18 24.1 8.2 1.4 3.4 0.20 17 154 0.9 0.32 -- 15 0.59 50 

EFSCAF 101105 10/11/05 18 17.6 8.3 1.5 -- 0.28 -- 197 1.4 -- -- -- -- -- 

Table 3. Water chemistry data from springs and creeks, 2002-2005, Mount St. Helens, Washington.[Acidity (pH) and conductivity data in bold are from laboratory measurements, HCO3 data in bold is determined on the basis of the measured amount of gas from the dissolved 

Unless specified, all units are in parts per million. mS/cm, millisiemens per centimeter; --, no data. Additional chemical analyses are available in the digital appendix that accomthis work.]
MSH-05-23 10/25/05 15 25.4 7.2 2.0 5.7 0.36 67 232 1.9 21 787 16 0.81 54 

WFSCAF 042105 04/21/05 17 4.2 5.7 0.19 0.09 <0.003 12 47 2.2 0.0 -- 0.90 0.01 6

WFSCAF 063005 06/30/05 17 6.6 5.7 0.09 -- 0.01 -- 20 3.2 -- -- -- -- -

MSH-05-10 07/26/05 17 10.5 5.8 0.06 0.03 <0.003 3.3 10 2.4 0.01 2.0 0.56 0.004 1

WFSCAF 082605 08/26/05 17 16.4 6.3 0.05 0.02 0.01 2.4 6 1.7 0.03 -- 0.58 0.004 1WFSCAF 101105 10/11/05 17 7.5 6.6 0.10 -- 0.01 -- 17 2.2 -- -- -- -- -

SCBF 101304 10/13/04 21B 14.5 8.5 0.79 -- 0.100 -- 84 0.9 -- -- -- -- -SCBF 102004 10/20/04 21B 9.1 7.8 0.20 -- 0.03 -- 26 1.1 -- -- -- -- -SCBF 120204 12/02/04 21B 4.1 8.5 1.0 2.6 0.14 31 94 1.3 0.10 -- 8.6 0.39 24 

SCBF 021005 02/10/05 21B 5.6 8.6 1.4 -- 0.20 -- 140 1.4 -- -- -- -- -

SCBF 063005 06/30/05 21B 14.5 7.9 0.34 -- 0.06 -- 48 3.0 -- -- -- -- -

MSH-05-13 07/26/05 21B 21.2 8.4 0.46 1.2 0.07 12 52 2.0 <0.04 -- 5.6 0.20 13 

SCBF 082605 08/26/05 21B 19.2 8.7 0.89 -- 0.12 -- 93 1.3 -- -- -- -- -

Other Loowit Creek locations MSH-02-05 08/18/02 10 31.7 7.9 1.9 3.6 0.38 50 220 1.9 0.23 -- 29 0.50 20 

MSH-02-06 08/18/02 12 2.8 6.0 0.23 <0.02 0.004 23 1.3 1.4 0.62 92 3.4 0.03 6

MSH-05-01 07/26/05 13 7.7 6.4 0.08 0.04 0.010 7 16 2.3 0.02 4 0.54 0.005 2

Pumice Plain springCARB 081204 10/13/04 20 17.4 7.8 0.9 -- 0.12 -- 81 1.2 -- -- -- -- -

MSH-05-08 07/26/05 20 25.4 7.3 1.2 2.1 0.18 48 107 2.0 0.08 243 11 0.30 12 

MSH-05-25 10/25/05 19 15.7 7.4 0.8 1.3 0.12 28 71 1.1 0.27 187 6.9 0.19 7

Regional cold springsMSH-02-10 08/23/02 23 3.2 7.2 0.05 <0.02 0.003 3.5 1.0 0.1 0.01 25 0.73 0.004 1

MSH-05-15 07/29/05 22 4.4 5.7 0.04 0.004 <0.003 3.4 0.7 0.0 0.01 16 0.66 0.001 0MSH-05-14 07/29/05 23 4.7 5.8 0.05 0.01 <0.003 3.6 1.1 0.1 0.01 28 0.74 0.006 1MSH-05-17 07/30/05 24 3.0 8.2 0.08 0.01 <0.003 4.9 3.1 0.2 0.01 11 0.80 0.004 1MSH-05-16 07/30/05 25 5.4 6.3 0.13 0.02 <0.003 13 1.4 0.2 <0.04 8 1.5 0.007 3MSH-05-19 07/31/05 26 7.2 6.2 0.12 0.10 0.010 4.7 6.5 0.6 <0.04 54 1.8 0.048 2MSH-05-18 07/31/05 27 6.3 6.5 0.09 0.08 <0.003 4.3 5.4 0.6 0.01 33 1.0 0.028 2MSH-05-20 07/31/05 28 5.2 6.2 0.09 0.03 <0.003 7.0 2.1 0.2 <0.04 52 1.3 0.011 2

 Sample No. Date Map No. Temp. (°C) pH Cond (mS/cm) B Br Ca Cl F Fe HCO3 K Li Mg 

Table 3. Water chemistry data from springs and creeks, 2002-2005, Mount St. Helens, Washington.—Continued[Acidity (pH) and conductivity data in bold are from laboratory measurements, HCO3 data in bold is determined on the basis of the measured amount of gas from the dissolved 

Unless specified, all units are in parts per million. mS/cm, millisiemens per centimeter; --, no data. Additional chemical analyses are available in the digital appendix that accomthis work.]
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 531

Figure 3. Ternary plot of anion proportion by mass for spring 

waters at or around Mount St. Helens, Washington (after 

Giggenbach, 1992). Arrows show change in major anions in 

Loowit springs since 1983 and Step springs since 1988. Data 

from dates before 2002 are from Thompson (1990), Shevenell 

and Goff (1993), Goff and McMurtry (2000), and unpublished 

data from the U.S. Geological Survey and Fraser Goff 

(University of New Mexico).

or the minor differences in the chemistry of the Loowit springs 

between 2002 and 2005 (table 3) to the current eruption.

Chemistry of the Regional Springs

All of the regional springs have temperatures <8°C and 

conductivities <150 μS/cm but can be divided into three 

groups based on variations in major anion concentrations 

(fig. 3; table 3). One group is defined by relatively high SO4

concentrations, and the other groups are characterized as 

HCO3-dominated waters but are subdivided on the basis of 

Cl concentrations.

The SO4 waters consist of springs 24 and 25, which are 

north of Mount St. Helens on the eastern border of the Pumice 

Plain (fig. 1). Spring 25 has higher SO4 concentrations than 

spring 24, and samples collected since 1988 show that anion 

compositions have been variable (fig. 3). In the late 1980s, 

SO4

-rich warm springs occurred in the cooling pyroclastic 

deposits of the Pumice Plain. Shevenell and Goff (1995) noted 

that these springs were not “steam-heated” waters, which are 

characteristically SO4 rich, but instead probably were leaching SO4

 from CaSO4 alteration minerals or from encrustations 

around the vents of short-lived fumaroles on the pyroclastic 

flows. Springs 24 and 25 are likely to be cold, dilute examples 

of this type of water.

The HCO3-rich waters consist of springs 22, 23, and 28, 

which are solely HCO3 waters, and springs 26 and 27, which 

have slightly elevated Cl in comparison with the other three 

springs (fig. 3; table 3). Springs 26 and 27 are southeast of 

Mount St. Helens and are of interest because (1) their anion 

chemistry lies along the dilution trend defined by the hot 

springs and (2) their 14C values are greatly reduced relative to 

modern carbon (table 4), as discussed in more detail below. 

The elevated Cl in these springs may provide evidence for 

Na

Ca

K (×10)

2,000

1,600

1,200

800

400

0

B (×25)

Cl

SO4

800

600

400

200

0

HCO3

Mg (×10)

6/83 6/88 6/93 6/98 6/03 6/83 6/88 6/93 6/98 6/03

1,200

1,000

800

600

400

200

CONCENTRATION, IN PARTS PER MILLION

0

CONCENTRATION, IN PARTS PER MILLION

CONCENTRATION, IN PARTS PER MILLION

DATE DATE

6/83 6/88 6/93 6/98 6/03

DATE

Figure 4. Chemistry of Loowit springs waters from 1983 to 2005, Mount St. Helens, Washington. Data from dates before 2002 are from 

Thompson (1990), Shevenell and Goff (1993), Goff and McMurtry (2000), and unpublished data (Fraser Goff, University of New Mexico).

HCO3

–

Cl–

SO4

-2

Mature waters

Volcanic

waters

1983–1989 1994–2005

1983–1985

1986–1989

1994

2002

2005

 Loowit springs Step springs

EXPLANATION

1988–1989

1994

2005

Carbonate spring

Regional springs

23 26

27

22

28 25 24

Meteoric

waters

Steam-heated

waters
532 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

-110

-108

-106

-104

-102

-100

-98

-96

-94

-92

-90

-88

-86

-84

-82

-80

-78

-76

δD, IN PER MIL

δ18O, IN PER MIL

-15 -14 -13 -12 -11 -10

World meteoric water line

22

25

24

28

23

27

26

1984–1985 1988–1989

1986–1991 1994

1994 2005

2002

2005

Carbonate spring

Snow

Loowit/Step creeks

Loowit springs

EXPLANATION

Step springs

Regional springs in north

Regional springs in south

Figure 5. Stable-isotope composition of Loowit springs since 

1980s, Mount St. Helens, Washington. Data from dates before 

2002 are from Thompson (1990), Shevenell and Goff (1993, 

1995), Goff and McMurtry (2000), and unpublished data from 

the U.S. Geological Survey.

some leakage of Loowit-type water through the southeastern 

wall of the crater.

Gas Transfer from Magma into Water

Gases released from magma can be variably dissolved into overlying cold water or released to a brine phase 

(Fournier, 1986) that can then mix with cold water. The 

resulting fluid is acidic, hot, and highly reactive with the 

surrounding rock. Mineral dissolution converts the acidic gas 

species H2S, SO2, HCl, and CO2 into the anions SO4, Cl, and 

HCO3. The concentration of these anions in the spring waters 

can be related back to magmatic degassing only if the effects 

of variable gas solution and mineral dissolution/precipitation 

are constrained. Some constraints can be derived from the 

water chemistry.

Many metallic ions (such as Ca) are not transported 

in volatile form (Symonds and others, 1993), requiring that 

the major source of the cations in the springs be water-rock 

interaction. This process could ultimately remove SO4 and 

HCO3 from solution via precipitation. Attempts at geothermometry by Shevenell and Goff (1993; 1995) yielded maximum temperatures between 120 ̊C and 175 ̊C for the Loowit 

waters. For an assumed temperature of 150 ̊C, calculations 

using SOLMNEQ88 (Kharaka and others, 1988) indicate 

saturation with anhydrite for waters collected in the 1980s but 

undersaturation for samples from 2002 and 2005. From the 

1980s through 2005, Ca, Na, K, and SO4

 in the spring waters 

all show similar patterns of decline that are consistent with 

dilution (fig. 4), so there is no evidence that anhydrite was ever 

a significant control on SO4 concentrations. Samples from all 

collection years are saturated with calcite at all temperatures 

above the discharge temperature. Calcite formation/dissolution may exert some influence on Ca and HCO3

. However, the 

large increase in HCO3 after the 1980s cannot be attributed to 

calcite dissolution, which would also produce both higher Ca 

concentrations and higher pH values. Ca and pH have in fact 

dropped during this time (table 3).

Shevenell and Goff (2000) argued that halite formation in 

the magmatic fluid component was probably not a significant 

sink for Cl. Although Br and Cl concentrations in the springs 

have declined steadily over the years, it is noteworthy that 

for the past two decades the ratios of these two components 

have remained at a remarkably constant value around 0.002 

(1984–2005, 44 samples, ± 0.001). The Step canyon springs 

and carbonate spring have lower concentrations of Br and Cl 

than the Loowit springs but fall along the same trend (fig. 6). 

The constant Cl/Br ratio shows that halite was never a major 

control on Cl in fluid sources to the hot springs.

In contrast, fumarole condensates have shown a large 

range in Cl/Br (fig. 6), possibly indicating that deposition 

of metal halides in the vent throats (Keith and others, 1981; 

Edmonds and others, this volume, chap. 27) can be a significant control on fumarolic Cl/Br ratios. The Cl/Br ratio 

in the springs, which lies within the range of the fumarolic 

condensates, may better represent the HCl/HBr ratio in gases 

released from magma. Additional Cl and Br could be leached 

from the volcanic rocks. Reported Br and Cl ratios in seven 

samples of andesite and dacite from the crater, The Breach, 

and Pumice Plain are variable, ranging between 0.002 and 

0.01 (table 1 of Shevenell and Goff, 1993). However, the SO4

-

rich nature of the transient hot springs in the Pumice Plain 

(Shevenell and Goff, 1995) indicates that rock leaching is not 

a major source of halides.

Overall, the data indicate that mineral sources and sinks 

for SO4, Cl, and HCO3 are not the main controls on the anions 
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 533

Sample No. Date Map 

No.

Temperature

(°C)

δD

per mil

δ18O

per mil

13C-DIC

per mil

14C-DIC

percent

modern carbon

Springs in Loowit canyon

MSH-02-01 08/16/02 4 58.9 -94 -12.9 -12.3 --

MSH-02-02 08/17/02 6 61.0 -95 -12.7 -- --

MSH-02-03 08/17/02 7 62.3 -94 -12.6 -12.2 --

MSH-02-04 08/18/02 8 57.6 -92 -12.5 -11.9 --

MSH-05-02 07/26/05 3 59.1 -91 -12.4 -- --

MSH-05-03 07/26/05 3 57.0 -92 -12.5 -12.1 --

MSH-05-04 07/26/05 5 61.9 -91 -12.3 -11.9 --

MSH-05-05 07/26/05 5 60.5 -91 -12.3 -- --

Springs in Step canyon

MSH-05-06 07/26/05 16 45.1 -90 -12.6 -11.4 --

MSH-05-07 07/26/05 14 56.6 -91 -12.3 -10.9 --

MSH-05-22 10/25/05 14 38.4 -90 -12.2 -- --

Monitoring network sites

MSH-05-23 10/25/05 15 25.4 -92 -12.1 -7.5 --

MSH-05-24 10/25/05 9 25.9 -89 -12.4 -8.1 --

Other Loowit Creek locations

MSH-02-06 08/18/02 12 2.8 -98 -13.8 -11.2 --

MSH-05-01 07/26/05 13 7.7 -93 -12.8 -- --

Pumice Plain spring

MSH-05-08 07/26/05 20 25.4 -89 -12.4 -11.5 5.5

MSH-05-25 10/25/05 19 15.7 -88 -12.5 -- --

Regional cold springs

MSH-05-15 07/29/05 22 4.4 -78 -11.4 -- --

MSH-02-10 08/23/02 23 3.2 -82 -11.9 -14.2 93.8

MSH-05-14 07/29/05 23 4.7 -81 -11.8 -15.6 --

MSH-05-17 07/30/05 24 3.0 -95 -13.4 -- --

MSH-05-16 07/30/05 25 5.4 -93 -13.0 -- --

MSH-05-19 07/31/05 26 7.2 -84 -12.1 -16.5 50.0

MSH-05-18 07/31/05 27 6.3 -87 -12.6 -16.0 54.8

MSH-05-20 07/31/05 28 5.2 -80 -11.7 -15.9 53.1

Table 4. Isotope values determined for water samples collected during 2002–2005, Mount St. Helens, 

Washington.

[DIC, dissolved inorganic carbon; --, no data.]
534 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

in the hot springs. We conclude that the changing anion proportions shown in figure 3 reflect changes in the release rates 

of sulfur gases, HCl, and CO2 from the magma and a varying 

degree of efficiency of gas scrubbing by the overlying water.

The 1983–89 trend in Loowit waters from SO4 toward 

Cl (fig. 3) can be explained by the cessation in the supply of 

magma at the end of the 1980–86 dome-building eruptions. 

With declining output of sulfur gas from magma, less sulfur 

was available to be scrubbed into overlying meteoric water. The 

simultaneous increase in Cl can be attributed to crystallization 

of extruded and shallowly emplaced magma, with degassing of 

HCl or partitioning of Cl into an aqueous phase that then mixed 

with meteoric water. The ensuing trend toward HCO3 dominance after 1989 (fig. 3) would logically be attributed to continuing declines in the supply of S and Cl as cooling and crystallization of the 1980–86 dome and shallow underlying magma 

progressed toward completion; CO2, due to its low solubility 

even at high pressure, continued to exsolve from magma at 

greater depths. This scheme does not explain the large increase 

in HCO3

 concentrations in Loowit springs (fig. 4; table 3). 

Coupled with the drop in pH, the rise in HCO3 concentrations 

between 1989 and 2002 represents an order-of-magnitude jump 

in DIC, an event of particular interest because it could indicate 

the ascent of fresh, CO2-rich magma from depth beneath the 

chamber, leading up to the renewal of eruptive activity in 2004. 

0.0

0.5

1.0

1.5

2.0

2.5

0 200 400 600 800 1,000

Br CONTENT, IN PARTS PER MILLION

Cl CONTENT, IN PARTS PER MILLION

Seawater trend

Loowit-Step

trend

1980s 1988–1989

1994

2005

Carbonate spring

1994

2002

2005

Fumarole condensate

Loowit springs

EXPLANATION

Step springs

Figure 6. Trends in Br and Cl concentrations for 2002–2005 

spring waters, compared with early data from springs and 

fumarole condensates from the 1980–86 dome at Mount St. 

Helens, Washington. One 1989 fumarole condensate sample that 

plots off the scale is not shown. Data from fumarole condensates 

and waters before 2002 are from Shevenell and Goff (1993, 2000), 

Goff and McMurtry (2000), and unpublished data (Fraser Goff, 

University of New Mexico).

However, the HCO3 increase seems to have begun by 1994 (fig. 

4), and carbon isotope data (discussed below) do not indicate 

CO2 release from a new source of magma.

By 2002, the molar C:S:Cl ratio in Loowit waters had 

shifted from ~1:3:8 in the late 1980s to 10:1.4:3.6. This ratio 

is much closer to the 10:1.1:1.2 ratio measured in plume emissions by various techniques in 2005 (Gerlach and others, 2005; 

Edmonds and others, this volume, chap. 27). Our preferred 

explanation for the shift in the ratios is that for many years 

after dome emplacement in the 1980s, temperatures above 

boiling in the rocks of the 1980–86 dome prevented shallow waters from completely sealing off escape pathways for 

gas. Highly soluble HCl and SO2 were effectively scrubbed 

by infiltrating water, but the CO2 was still largely able to 

escape due to its much lower solubility. Once gas pathways 

were mostly blocked by liquid water, presumably after 1994, 

CO2 scrubbing became efficient, and dissolved C:S:Cl ratios 

approached those of degassing magma. Note that the timeframe for these changes in chemistry is in fair agreement with 

the estimate that the 1980–86 dome would cool through its 

magnetization temperature in 18–36 years after emplacement 

(Dzurisin and others, 1990).

Carbon and Helium Isotope Evidence on 

Magmatic End Members

Gas from the Crater and The Breach

In spite of declining fumarole temperatures and the 

increasing amount of air in gases venting in the crater at 

Mount St. Helens, the carbon and helium isotopic compositions give evidence of a prolonged magmatic input. Gas 

collected in November 1980 from a >400°C fumarole near the 

dome provided the most representative δ13C−CO2 value for the 

early magma at Mount St. Helens, −10.5 per mil (Evans and 

others, 1981; table 5). Over time the carbon isotope composition of CO2

 in fumarole gases shifted to lighter values (fig. 

7A). The shift cannot be related to the increased air concentrations because the δ13C value of atmospheric CO2

 is around 

−7 per mil (Faure, 1986). Since 1994 CO2 concentrations in 

gases from the 1980–86 dome have declined, but δ13C–CO2

values have stabilized (fig. 7). In June 1994, gas from a 560°C 

fumarole and other lower temperature fumaroles on the dome 

contained about 32 percent CO2

 and had δ13C values between 

−11.7 and −12.0 per mil (Goff and McMurtry, 2000; F. Goff, 

unpub. data, 1994). In 1998, gas samples collected from an 

86°C vent on September lobe contained 5 percent CO2 with a 

δ13C value of −11.9 per mil (Symonds and others, 2003; table 

5), and gas samples collected from vents on September lobe 

in 2002 had less than 3 percent CO2

 with δ13C values between 

−12.0 and −11.8 per mil (table 5).

Gas from the bubbling springs in 2002 and 2005 contained mostly CO2

 but also had considerable N2 and Ar (table 

5). The N2/Ar ratio in the gas falls between the ratios for air 

and air-saturated meteoric water, but low O2 values show that 
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 535

direct air contamination was not a factor. The R/RA values for 3

He/4He ratios in the gas were 5.5 and 5.8, respectively, similar 

to a 1989 hot spring water sample that had a value of 5.7 (Goff 

and McMurtry, 2000). The δ13C value for the CO2 in both 

years was −14.7 per mil, significantly lower than the isotopic 

composition of CO2 in the vent gases. Analysis of the 2002 

bubbling gas indicates that it is essentially 14C dead, ruling out 

any significant biogenic CO2 source (table 5; figs. 7A, 7B). 

An important consideration is that equilibrium fractionation 

processes call for lower δ13C values in the gas-bubble CO2, 

relative to the DIC, consisting of HCO3 and dissolved CO2.

Dissolved Inorganic Carbon in Spring Waters

The carbon isotope composition of DIC from the springs 

in Loowit and Step canyons and carbonate spring has a small 

range in values that is independent of water temperature (fig. 

7A). The δ13C–DIC values in all these springs are similar to 

0 10 20 30 40 50 60 70

-22

-20

-18

-16

-14

-12

-10

-8 A B 

-22

-20

-18

-16

-14

-12

-10

-8

0 20 40 60 80 100

DIC-CO2

equilibration

14 TEMPERATURE, IN °C C CONTENT, IN PERCENT MODERN CARBON

δ13C, IN PER MIL

28

23 (2002)

27

26

1980 range,T ≥ 93°C

1994–2002 range, T ≥ 85°C

28

27

26

Atmospheric 

CO2

Biogenic 

CO2

δ13

C, IN PER MIL 

Step springs DIC range

Loowit springs DIC range

23 (2005)

mixing

23 (2002)

23 (2005*)

CO2 Loowit springs

EXPLANATION

CO2 September lobe

DIC carbonate spring

DIC Loowit springs

DIC Step springs

DIC regional springs

the δ13C composition of fumarolic CO2 from recent years 

but are lower than the 1980 fumarolic δ13C–CO2 values. The 

δ13C–DIC data thus provide additional evidence for a change 

in the δ13C of the CO2 released from the underlying magma. 

This carbon isotope shift to lower values since 1980 is most 

easily attributed to fractional loss of 13C-enriched CO2 during 

degassing of the magma (Gerlach and Taylor, 1990) and is a 

strong argument against the involvement of fresh magma. A 

new pulse of gas-rich magma from great depth would likely 

cause a return toward heavier δ13C values in gas vents and 

DIC, and a new pulse of basalt from mantle depths might 

bring an increase in 3He/4He ratios.

The regional springs have much lower HCO3 concentrations than springs in the crater and The Breach and have lower 

δ13C–DIC values (tables 3, 4). The 14C content of the DIC 

in springs 26, 27, and 28 contains about 50 percent modern 

carbon, and the 13C–14C isotope values fall on a mixing line 

between typical biogenic carbon and magmatic carbon in the 

dome vents and hot springs (fig. 7B). Springs 26, 27, and 28 

Figure 7. Water temperatures and carbon isotope compositions of dissolved inorganic carbon (DIC) and CO2

from vents and springs at and around Mount St. Helens, Washington 1980–2005. A, δ13C–DIC values for hot and 

cold springs from this study compared to CO2

 gas collected from the 1980–86 dome. Horizontal arrows represent 

the range in δ13C–CO2

 values reported for 1980s dome gas samples from Evans (1981), Goff and McMurtry (2000), 

Symonds and others (2003), and this study. B, Comparison of δ13C–DIC and 14C–DIC values for some regional cold 

springs with values of DIC and CO2

 from hot springs and atmospheric and biogenically derived CO2. 14C values for 

DIC in Loowit and Step springs and CO2

 from the September lobe are assumed to be zero. The 2005 analysis of 

sample 23, shown with asterisk, lacks a 14C–DIC value, so it is arbitrarily assigned the 2002 analytical value, for 

purposes of plotting the δ13C results.

25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 537

may contain some magmatic carbon, and the Cl concentrations in springs 26 and 27 support this idea. The DIC in spring 

23 contains 94 percent modern carbon, and two samples from 

different years had variable δ13C values that are higher than the 

other regional springs. The variation in the 13C composition may 

indicate that some CO2 degassing occurs from the water prior 

to its point of emergence. A small magmatic component in the 

spring is possible, but without additional study (for example, 

He isotopes), we cannot eliminate the possibility that the dead 

carbon in all four of these springs is derived from Tertiary-age 

hydrothermal carbonate in the volcanic rocks that underlie the 

edifice of Mount St. Helens (Evarts and others, 1987).

Total Discharges of Water, Heat, and Magmatic 

Volatiles

Creek-Water Discharge

During the course of this investigation, the measured surface discharge from Loowit Creek was about three to six times 

greater than the discharge from Step creek (table 2). Maximum seasonal discharge at the monitoring sites occurred in 

the midsummer to fall months. There are no historical data on 

discharge at Step creek, but measurements at LCAF show that 

flow rates have increased since 1989 (fig. 8; tables 2 and 6). 

The increase does not reflect variations in annual precipitation 

and more likely results from increased input of meltwater as 

Crater Glacier grew from a persistent snow bank in the shadow 

of the crater rim to an area of about 1 km2 by September 2001 

(Schilling and others, 2004).

Since October 2004, the emplacement of the new lava 

dome at Mount St. Helens resulted in extensive deformation 

of Crater Glacier and a concomitant decrease in the volume 

of glacial ice, without any large change in the surface flows 

of the creeks, indicating that some meltwater may exit the 

crater from leakage through the crater floor (Major and others, 2005; Walder and others, 2007; Walder and others, this 

volume, chap. 13). A complete hydrologic budget to address 

this idea has yet to be developed for Mount St. Helens and, 

until such time, our multiple measurements of stream flow 

and chemistry for Loowit and Step creeks (table 7) offer the 

most comprehensive picture of discharge and dissolved fluxes 

from the crater area.

Two major contributors to conductivity in the creek 

waters are Cl and SO4

 (fig. 9A). During the course of our 

investigation at both of the Loowit Creek monitoring sites, Cl/

SO4 ratios were essentially constant and mimicked ratios in 

Loowit springs in 2005 (fig. 9B). In contrast, Step creek waters 

show large variations in Cl to SO4

, which can be attributed 

directly to changes in the chemistry and discharge of water 

from WFSCAF. Combining the average flows of Loowit Creek 

and the east fork of Step creek yields a total thermal water 

discharge of 335 L/s with an average Cl concentration of 180 

mg/L (table 7). The creek waters are composed primarily of 

meteoric water from melting ice and snow plus water from the 

0 0

50

100

150

200

250

300

350

400

1984 1989 1994 1999 2004

50

100

150

200

250

300

350

YEAR

DISCHARGE, IN LITERS PER SECOND

Onset of

seismic unrest

YEARLY PRECIPITATON, IN CENTIMETERS

Yearly precipitation

EXPLANATION

Discharge

Figure 8. Discharge at site LCAF from 1985 to 2005 and 

yearly precipitation totals from the SNOTEL station at 

Spirit Lake, Mount St. Helens, Washington, from 1986 to 

2005. Ticks on x axis correspond to month of June. Data 

from August 1988 and 1989 are from Shevenell (1990).

 Date Discharge 

(L/s) Meter type

08/15/85 161 Price AA meter

09/09/85 197 Price AA meter

08/09/88 195 Gurley meter

08/10/89 213 Gurley meter

09/14/89 220 Price AA meter

09/14/89 197 Price AA meter

10/16/89 116 Pygmy meter

01/17/90 109 Pygmy meter

08/24/90 132 Pygmy meter

09/20/90 126 Pygmy meter

10/23/90 109 Pygmy meter

12/07/90 117 Pygmy meter

02/08/91 144 Pygmy meter

05/01/92 161 Pygmy meter

07/28/92 160 Pygmy meter

07/28/92 144 Pygmy meter

05/05/93 118 Pygmy meter

07/07/94 212 Pygmy meter

09/01/98 248 Pygmy meter

Table 6. Discharge measurements and meter 

information for Loowit Creek above the falls, 

1985–98, Mount St. Helens, Washington.
Sample No. Location Date Temp. (°C) Conductivity (mS/cm) Discharge (L/s) 14C (% modern carbon) HCO3 (mg/L) CO2 (t/d) CO2 (mag. t/d) Cl (mg/L) 

Monitoring network sites MSH-05-11 LCAF 07/26/05 29.1 1.63 396 488 12.1 24 172 

LCAF 082605 LCAF 08/26/05 30.1 1.61 408 -- -- -- -- 169 

LCAF 101105 LCAF 10/11/05 30.0 1.94 267 -- -- -- -- 197 

MSH-05-24 LCAF ps 10/25/05 25.9 1.83 215 -- 533 7.1 14 193 

LCBF 101304 LCBF 10/13/04 16.3 1.49 306 -- -- -- -- 147 

LCBF 042105 LCBF 04/21/05 16.9 1.87 187 -- -- -- -- 180 

MSH-05-12 LCBF 07/26/05 22.5 1.56 399 -- -- -- -- 167 

LCBF 082605 LCBF 08/26/05 19.8 1.63 279 -- -- -- -- 178 

Average flux 19.0 MSH-05-09 EFSCAF 07/26/05 26.6 1.45 45 -- 535 1.5 3 172 

EFSCAF 101105 EFSCAF 10/11/05 17.6 1.52 19 -- -- -- -- 197 

MSH-05-23 EFSCAF ps 10/25/05 25.4 1.96 27 -- 787 1.3 2.6 232 

Average flux 2.8 MSH-05-10 WFSCAF 07/26/05 10.5 0.06 142 -- 12 0.1 0.1 10 

SCBF 101304 SCBF 10/13/04 14.5 0.79 112 -- -- -- -- 84 

SCBF 082605 SCBF 08/26/05 19.2 0.89 49 -- -- -- -- 93 

MSH-05-13 SCBF 07/26/05 21.2 0.46 110 -- -- -- -- 52 

Average flux Pumice Plain springMSH-05-25 CARB gs 10/25/05 15.7 0.80 354 5.5 204 4.5 4.3 71 Regional cold springs

MSH-05-14 #23 07/29/05 4.7 0.05 818 93.8 29 1.5 0.16 -- 

MSH-05-18 (nr) #27 07/31/05 6.3 0.08 82 54.8 44 0.2 0.11 -- 

MSH-05-19 nr #26 07/31/05 7.2 0.12 137 50.0 94 0.8 0.42 -- 

MSH-05-20 #28 07/31/05 5.2 0.09 29 53.1 73 0.1 0.07 -- 

Loowit Creek above falls, 1988 and 1989SH-45 LCAF 08/09/88 41.0 3.36 195 -- 87 1.1 1.1 600 

SH-81 LCAF 08/10/89 41.6 3.02 213 -- 88 1.2 1.2 549 

Average flux 1.1 

Table 7. Selected chemistry and discharge measurements for monitoring sites and some springs, 1988–2005, Mount St. Helens, Washington.[CO2 t/d, metric tons per day, calculated as CO2 from measured HCO3 flux; CO2 mag., the corrected magmatic flux of CO2 in metric tons per day based on measured spring pH val

mens per centimeter; ps, probe site; LCAF, Loowit Creek above falls; LCBF, Loowit Creek below falls; EFSCAF, east fork Step creek above falls; WFSCAF, west fork Step creek below falls; CARB gs, carbonate spring gage site. Data from 1988 and 1989 are from Shevenell (1990). The 14C data in bold were collected at a different time from the time the dis
25. Pre- and Post-Eruptive Investigations of Gas and Water Samples from Mount St. Helens, Washington, 2002 to 2005 539

B

A

Loowit springs 2002

Loowit springs 2005

LCAF

LCBF

Carbonate spring

Other Loowit and Step creek sites

Step springs 2005

EFSCAF 

WFSCAF

SCBF

9/04 12/04 3/05 6/05 9/05

0.0

0.5

1.0

1.5

2.0

2.5

3.0

3.5

7.0

8.0

Cl/SO4 RATIO

DATE

r = 0.98

0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5

0

100

200

300

400

500

600

700

Cl + SO4 CONTENT,

IN PARTS PER MILLION

CONDUCTIVITY, IN MILLISIEMENS PER CENTIMETER

Step

Loowit

2005 spring ratios

Figure 9. Cl and SO4 concentrations for Loowit 

spring waters since 2002 and the creek-monitoring 

sites since October 2004, Mount St. Helens, 

Washington. A, Cl and SO4

 are major anions 

controlling conductivity in creek waters. B, Temporal 

plot showing Cl/SO4

 ratios at the monitoring sites. 

Loowit Creek sites show no change during the first 

year after the start of the 2004–5 eruption. Sites SCBF 

and WFSCAF show change in Cl and SO4

 ratios. Gray 

bars represent range in Cl and SO4

 for Loowit canyon 

and Step canyon springs in 2005. The LCAF site plot is 

shifted slightly so data can be shown.

hot springs, and we can normalize the creek data to a source 

spring fluid with 280 mg/L by assuming that the meteoric 

water has no Cl (table 3). This calculation yields total water 

outputs of 215 L/s for the Loowit and Step springs. By using 

the density and enthalpy of water at 56°C, the heat output 

is ~50 MW. For perspective, a heat output of 50 MW would 

completely cool the 1980–86 dome in about 125 years.

Present-Day Flux of Magmatic Volatiles

Assuming that all of the dissolved carbon in the hot 

springs is derived from the underlying magma, we can couple 

the DIC concentration and the discharge to provide an estimate 

of the magmatic carbon flux. Table 7 shows six datasets for 

dates when a DIC sample was collected from a monitoring site 

concurrent with a discharge measurement. The largest flux 

measured was 12 metric tons per day (t/d), expressed as CO2, 

from LCAF in July 2005. On the same day, another 1.5 t/d (as 

CO2) discharged from EFSCAF. Measured pH values show 

that the DIC in Step and Loowit springs consists of approximately equal proportions of dissolved CO2

 and HCO3, but 

nearly all of the dissolved CO2 is outgassed upstream of the 

monitoring sites. The total output of magmatic carbon from 

the springs would, therefore, be double the amounts noted 

above. We did not gage carbonate spring in July 2005, but in 

October 2005 its magmatic CO2

 flux was 4.3 t/d. That value is 

adjusted for a small modern, nonmagmatic carbon contribution 

indicated by the 14C value using mass balance, assuming the 

modern DIC component contains 105 percent modern carbon 

(James and others, 1999). Any potential flux of magmatic CO2

from WFSCAF or the regional springs is less than 1 t/d.

We also estimated Cl and SO4 fluxes for the creeks using 

the same methods. Apart from carbonate spring, which may 

have some nonmagmatic sources of Cl and SO4, the data in table 

7 show that the Loowit Creek sites are the main contributors of 

Cl and SO4. Together, the Loowit and Step creeks yield average Cl and SO4

 fluxes of 5.2 t/d (expressed as HCl) and 4.7 t/d 

(expressed as SO2) and, adjusting the DIC as discussed above, 

22 t/d of CO2. These data provide a first look at the flux of magmatic components dissolved in the waters at Mount St. Helens. 

For comparison, HCl and SO2 fluxes in plume emissions on 

August 31, 2005, were 14 and 22 t/d, respectively (Edmonds 

and others, this volume, chap. 27), whereas CO2 in the craterplume emissions during the course of the 2004–6 eruption were 

in hundreds of tons per day (Doukas and others, 2005; Gerlach 

and others, this volume, chap. 26), and CO2 emissions from the 

plume coming off Loowit springs were 27 t/d on June 9, 2005 

(Gerlach and others, this volume, chap. 26).

Flux of Magmatic Volatiles 1988–1989

Estimates similar to those made for the 2004–5 data can 

be made for the flux of magmatic volatiles at LCAF in the 

late 1980s by using data from Shevenell (1990). Tables 2 and 

7 show that the discharge at LCAF in August 1988 and 1989 

was roughly 60 percent of the flow measured in the months of 

July and August 2003–5. In spite of lower discharges of creek 

water in the late 1980s, dissolved SO4

 and Cl fluxes were 1.4 

and 2.2 times higher, respectively, than fluxes in 2004–5 (table 

7). The dissolved magmatic CO2 flux in 1988–89, however, 

was only 1.1 t/d, in comparison with 19 t/d from this study 

(table 7). Because the pH of the springs was about 7.2, we 

made no correction for outgassing of CO2

.
540 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Conclusions

Data from springs and gas discharges in the crater are 

consistent with plume emission rates (Doukas and others, 

2005; Gerlach and others, this volume, chap. 26) that indicate 

the current eruption is driven by degassed magma. From the 

early 1980s to 2002, gas vents on the 1980–86 dome grew 

weaker and more air-dominated, while magmatic Cl and SO4

inputs to the Loowit springs declined, and the isotopic signal 

(D and 18O) from magmatic water input disappeared. The large 

DIC increase in Loowit springs over time could indicate a 

pulse of new magma, but the δ13C signatures of the DIC and 

of CO2 gas are consistent with degassing of residual magma 

emplaced in the 1980s. No significant changes in Loowit 

springs occurred between 2002 and 2005. However, we note 

that the time for water to flow from the new dome to the 

spring vents could exceed one year, the time between the onset 

of unrest and the latest water sampling, so some caution must 

be applied to interpretations of the Loowit data. Carbon and 

helium isotope data for gas from the new dome might be more 

compelling but are unavailable.

Hot or mineralized springs occur on many volcanoes 

that lack summit plumes or obvious fumaroles and, in these 

cases, may provide the only opportunity to look for magmatic 

gases released at depth. The similarity in C:S:Cl ratios of 

recent Loowit springs samples and the 2005 plume emissions 

strongly supports the idea that spring geochemistry can accurately reflect magmatic degassing under favorable conditions. 

In this context, it is important to continue tracking the chemistry of the Mount St. Helens springs during the next few years 

in a search for responses to the current eruption.

Acknowledgments

We wish to thank several individuals who contributed to 

our research. Fraser Goff (Earth and Planetary Sciences Dept., 

Univ. of New Mexico) provided previously unpublished data 

from his work at Mount St. Helens. Helpful reviews were provided by Fraser Goff and Jon Major (USGS, Cascades Volcano 

Observatory). Peter Frenzen, the monument scientist at the 

Mount St. Helens National Volcanic Monument (USDA-Forest 

Service), provided logistical support and invaluable advice for 

several of our field campaigns. Student volunteers and interns 

Bryn Kimball (2002), Taryn López (2005), and Leif Rasmuson 

(2005) helped with sampling and stream gaging.

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rekindled; the renewed eruption of Mount St. Helens, 

2004–2006: U.S. Geological Survey Professional Paper 

1750 (this volume).
542 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Trace-Element and Isotopic Data from Springs and Creeks in the 

Mount St. Helens Area, Washington, 2002–2005

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

This appendix is a spreadsheet of analytical data.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 26

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Emission Rates of CO2, SO2, and H2S, Scrubbing, and 

Preeruption Excess Volatiles at Mount St. Helens, 2004–2005

By Terrence M. Gerlach1, Kenneth A. McGee1, and Michael P. Doukas1

Abstract

Airborne surveillance of gas emissions began at Mount 

St. Helens on September 27, 2004. Reconnaissance measurements—SO2 column abundances and CO2, SO2, and H2S 

concentrations—showed neither a gas plume downwind of 

the volcano nor gas sources within the crater. Subsequent 

measurements taken during the period of unrest before the 

eruption began on October 1 and for several days after October 

1 showed only small point sources of gas within the crater. 

These sources defined a pattern of scrubbed degassing that 

evolved from near-zero emissions, to scattered CO2-only 

sources, to growing sources of CO2 with minor H2S and SO2, 

and finally to myriad sources of CO2 with increasingly SO2-

dominant sulfur gases. Scrubbing strongly hydrolyzed SO2 but 

also affected CO2 and H2S.

From October 7 on, a coherent plume spilled over the 

crater rim, yielding emission rates for CO2 and SO2, but not 

always for H2S. Virtually all SO2 and most CO2 outgassed 

from the growing dome of new dacite; some CO2 came from 

sources on the 1980–86 lava dome and in the Loowit springs 

area. The 2004–5 emission rates were notably low and variable. Emission rates for CO2 peaked early (10/7/2004) at 2,415 

metric tons per day (t/d), but the median rate was 655 t/d; 

only about 20 percent of the rates were greater than 1,000 t/d 

and about 45 percent were less than 500 t/d. Emission rates of 

SO2 never exceeded 240 t/d, and the median rate was only 72 

t/d; 70 percent of SO2 emission rates were <100 t/d and about 

40 percent were <50 t/d. Emission rates of H2S were <10 

t/d. Cumulative outputs through November 2005 were about 

231,000 t CO2 and 30,000 t SO2.

The CO2 and SO2 emission rates are distinctly lower 

than those of the early 1980s, but they are similar to those of 

the 1980s lava-dome eruptions, possibly back as far as late 

1981 or late 1980. However, the CO2/SO2 ratio of the 2004–5 

emissions (11±1) is higher than that of the 1980s emissions (8) 

because of scrubbing during the early part of the eruption. The 

nonscrubbed CO2/SO2 of the 2004–5 gases is 9±1, similar to 

the1980s emissions.

Modeling constrained by cumulative CO2 emissions, 

cumulative dacite production, and melt H2O concentration 

confirms that the 2004–5 dacite is a “flat” magma—that is, a 

dacite magma greatly depleted in excess (exsolved) volatiles 

compared to May 18, 1980, dacite. The inferred excess-volatile 

content of the current dacite is only 1.2 volume percent (vol. 

percent), or 0.2 weight percent (wt. percent), compared to the 

15 vol. percent (3 wt. percent) of the May 18 dacite at 900°C 

and 220 megapascals (MPa) (8.6 km depth). At the much lower 

pressure of 130 MPa prior to ascent from the shallowest part 

of the reservoir (5.2 km depth), the current dacite’s inferred 

excess-volatile content is 1.2 wt. percent—significantly less 

than the 3 wt. percent of the May 18 dacite and on the low end 

of the 1–6-wt. percent range of deeper intermediate to silicic 

magmas commonly involved in explosive volcanism.

The modeling further indicates that before ascent from 

5.2 km depth, the dacite’s excess volatile phase was H2O rich 

(XH2

O = ~0.96, XCO2

 = ~0.04, mole fraction basis) with H2S/

SO2 > 40, and its rhyolitic melt phase contained about 4.4 wt. 

percent H2O and 37 parts per million (ppm) CO2. After closedsystem ascent to the depth range of groundmass crystallization (~0.5–1 km), XH2

O is 0.98, H2S/SO2 is 4–7, and the melt 

contains 1.1–1.5 wt. percent H2O and 1–3 ppm CO2. Sulfur 

dioxide becomes dominant over H2S at depths less than 0.2 

km. Because of excess-volatile depletion, open-system degassing involves depths shallower than about 2.5 km, weak CO2

and SO2 emissions prone to the effects of scrubbing, and—in 

contrast to early 1980 degassing—no measurable effect on the 

equilibrium (210Pb/226 Ra) values of the 2004–5 dacite.

The 2004–5 gas emissions are incompatible with “new” 

gas-rich magma introduced into the reservoir in the months 
544 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

just before or since the onset of unrest and eruptive activity. However, the gas emissions are compatible with a flat 

magma—either leftover dacite from 1986 or dacite injected 

into the reservoir since 1986 as excess volatile-depleted 

magma, or as gas-rich magma subsequently mixed with larger 

amounts of leftover dacite or stripped of its excess volatiles by 

degassing prior to the current eruption.

Introduction

Airborne surveillance of gas emissions began on September 27, 2004, four days after a shallow earthquake swarm 

beneath the 1980–86 lava dome signaled Mount St. Helens’ 

reawakening after 18 years (Scott and others, this volume, 

chap. 1; Moran and others, this volume, chap. 2). In this 

chapter we report the CO2, SO2, and H2S gas emissions from 

the eruption’s inception to the end of 2005, documenting their 

measurement, sources, and the magnitude and variability of 

the emission rates—presented in the context of gas-emission 

rates at Mount St. Helens in the 1980s and at several Cascade 

Range volcanoes since the late 1990s. The outstanding feature 

of the 2004–5 gas emissions is their persistently low emission rates. We attribute these low rates mainly to the extreme 

depletion of excess (that is, exsolved) volatiles in the current 

dacite compared to the May 18, 1980, dacite and secondarily to gas scrubbing. VolatileCalc modeling (Newman and 

Lowenstern, 2002) quantifies and confirms our hypothesis of 

excess-volatile depletion in the 2004–5 dacite. We exploit the 

modeling to reveal the amount and composition of excess-volatile fluid in the source dacite at depth, to track its evolution 

as the dacite ascends, and to quantify several related aspects 

of the degassing.

Background

Airborne Emission-Rate Measurements at 

Mount St. Helens in the 1980s

Airborne monitoring from a fixed-wing aircraft of SO2

emission rates began at Mount St. Helens in May 1980 by 

making correlation spectrometer (COSPEC) measurements on 

the plume. More than 1,000 COSPEC flights followed during 

the next 8 years (McGee, 1992; McGee and Casadevall, 1994), 

most of which were made 1–2 km downwind of source vents 

during episodes of dome-building eruptions, endogenous 

dome growth, or intervening noneruptive periods. Casadevall 

and others (1981, 1983) describe the measurement technique 

and discuss the SO2 emissions through 1982. McGee (1992) 

describes the structure, dynamics, and SO2 cross sections of 

noneruptive plumes at Mount St. Helens that occurred from 

1980 to 1988. McGee and Sutton (1994) discuss measurements made during four 1984–86 dome-building eruptions 

and compare results with various geophysical monitoring 

data. Early SO2 emission rates frequently exceeded 1,000 

metric tons per day (t/d), but rates declined to negligible levels 

by September 1988 when airborne COSPEC measurements 

ended, nearly 2 years after the last dome-building eruption of 

the 1980s (McGee and Casadevall, 1994). Airborne monitoring also included 119 measurements of CO2 emission rates 

that ranged from >20,000 t/d in July 1980 to <1,000 t/d by the 

end of August 1981, when measurements were halted because 

concentrations of plume CO2 had declined to levels indistinguishable from ambient atmospheric CO2 by the analyzer 

employed at the time (Harris and others, 1981; McGee and 

Casadevall, 1994). McGee and Casadevall (1994) tabulate the 

1980–1988 airborne emission-rate measurements and supplementary data, all of which can be downloaded from the USGS 

publications Web site, http://pubs.usgs.gov/of/1994/of94-212/

(last accessed March 14, 2008).

Airborne Emission-Rate Measurements in the 

Cascades Since the 1980s

We initiated a program of volcanic gas emission-rate 

measurements at Cascade Range volcanoes in the late 1990s 

after developing improved airborne monitoring techniques 

(described below). The program comprised airborne plume 

measurements at major volcanic centers, from Mount Baker in 

Washington to Lassen Peak in California, to establish baseline 

emission rates of CO2, SO2, and H2S for identifying anomalous 

gas emissions at times of volcano unrest in the Cascade Range 

(table 1). Several volcanic centers—Glacier Peak, Mount 

Rainer, Mount Adams, Mount Jefferson, Medicine Lake caldera and Mount Shasta—produce no detectable emissions of 

the three gases; however, these airborne results do not rule out 

minor diffusive degassing and scattered spring and fumarole 

gas discharges derived from underlying magma. Other centers, 

such as Newberry caldera and South Sister, though lacking 

coherent plumes, do produce trace quantities of gas from small 

point sources that can be detected in airborne gas measurements. Besides Mount St. Helens, only three volcanoes in the 

Cascade Range produce coherent plumes with measurable gas 

output at the present time—Mount Baker, Mount Hood, and 

Lassen Peak. All of these volcanoes emit small amounts of 

gas: emission rates are <300 t/d for CO2 and <10 t/d for H2S; 

SO2 is below detection.

Airborne Measurements at Mount St. Helens in 

1998

In May 1998, background seismicity at Mount St. Helens 

increased markedly from the low levels recorded for the 

previous few years. Located directly below the lava dome, 

these small earthquakes clustered mainly in two distinct depth 

bands: 2–5 km and 7–9 km. Throughout June and the first 

half of July 1998, the seismicity continued to increase, with 

earthquake depths spanning the entire range from 2 to 9 km. 

The first of several airborne gas-surveillance flights recorded 
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 545

Volcano Date Platform CO2

(t/d)

SO2

(t/d)

H2S

(t/d)

Mount Baker 07/08/1998 FX 273 0 –

09/13/20001 FX 187 0 5.5

03/21/2001 FX 0 0 0

Glacier Peak 07/08/1998 FX 0 0 –

Mount Rainer 07/08/1998 FX 0 0 –

Mount Adams 08/03/1998 FX 0 0 –

08/10/2005 FX 0 0 0

Mount St. 

Helens 06/22/1998 FX 1,900 0 –

07/08/19982 FX tr 0 –

07/08/19983 FX tr 0 –

08/03/1998 FX 0 0 –

09/14/1998 FX 0 0 –

Mount Hood 06/22/1998 FX 0 0 –

08/03/1998 FX 0 0 –

08/10/2005 FX 0 0 0

09/15/2005 Heli 144 0 6.4

Mount Jefferson 08/11/1998 FX 0 0 –

Newberry 

caldera 09/19/2000 FX tr 0 –

South Sister 04/25/2001 Heli tr 0 –

09/21/2001 Heli 0 0 –

Medicine Lake 

caldera 08/11/1998 FX 0 0 –

Mount Shasta 08/11/1998 FX 0 0 –

Lassen Peak 08/11/1998 FX 110 0 –

09/19/2000 FX 20 0 2

1

Data from McGee and others (2001).

2

Morning flight.

3

Afternoon flight.

Table 1. Gas measurements in the Cascade Range and nearby calderas after 1988.

[Platform indicates aircraft: FX, fixed-wing; Heli, helicopter. t/d, metric tons per day; –, no data; 0, 

measurement taken but gas was not detected; tr, trace.]
546 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

a CO2 emission rate of 1,900 t/d on June 22 (table 1). By July 

8, morning and afternoon gas-measurement flights detected 

only trace CO2 degassing, and by the end of July, seismicity 

returned to levels similar to those prior to May. Additional 

gas-surveillance flights on August 3 and September 14 failed 

to detect CO2 degassing. All 1998 flights failed to detect SO2; 

there were no H2S measurements.

Instrumentation and Configuration of 

the Airborne Monitoring System

In this study, we used an airborne system developed 

during the past decade for measuring volcanic CO2, SO2, and 

H2S emissions and intended primarily for application in the 

Cascade Range and the Aleutian arc. Sites of interest in these 

regions are remote, ground access is difficult, climates are wet, 

and thick snow pack and glaciers are present locally. Large 

volumes of recharge water foster hydrothermal systems and 

bodies of ground water and surficial water capable of scrubbing acidic magmatic volatiles (especially SO2, HCl, HF) and 

masking magma degassing at depth during the early stages of 

unrest—hence the emphasis on an airborne system capable 

of measuring CO2 and H2S emissions (Doukas and Gerlach, 

1995; Symonds and others, 2001).

The airborne system incorporates a LI-COR Model 

LI-6252 nondispersive infrared CO2 analyzer and flow control 

unit to measure CO2 concentrations. This instrument, with a 

fast 1-s response time and a high sensitivity to low concentrations of CO2, is well suited for measuring small changes in CO2

concentration. Its use for determining CO2 in volcanic plumes 

is described in detail elsewhere (Gerlach and others, 1997; 

Gerlach and others, 1999). We used an Interscan Model 4170 

H2S analyzer to measure plume H2S in the 0–1 ppm range. This 

instrument consists of an electrochemical voltametric sensor 

coupled to a 1-L/min sample-draw pump with H2S concentration recorded from calibrated analog output. McGee and others 

(2001) describe the application of the Interscan analyzer to 

measuring H2S in volcanic plumes. Similarly, an Interscan 

Model 4240 SO2 analyzer with a 0–2 ppm range provides direct 

measurements of SO2 in the plume. Assuming ambient air is 

largely devoid of H2S and SO2, our measurements indicate 

typical instrument noise values of 3–4 parts per billion (ppb)for 

these analyzers. Interscan analyzers show some cross sensitivity to other gases, so these gases are removed by appropriate 

chemical scrubbers installed on input lines. Because Interscan 

analyzers measure gas partial pressures, all calibrated output 

is corrected for the pressure and temperature at the altitude of 

measurement.

The airborne system is configured to employ two types 

of optical ultraviolet spectrometers for measuring SO2

 column 

abundances. For some of the early measurements in this 

study, we employed a Barringer correlation spectrometer 

(COSPEC V)—long the standard tool for monitoring volcanic SO2 emission rates (Stoiber and Jepsen, 1973; Crafford, 

1975; Malinconico, 1979; Stoiber and others, 1983). A data 

logger records the analog output of the instrument for later 

processing using USGS software. For the majority of the SO2

flux measurements, we used a FLYSPEC built and configured 

as described in Horton and others (2006). This new, miniaturized, lightweight and low-power, ultraviolet correlation 

spectrometer incorporates an Ocean Optics USB2000 ultraviolet spectrometer with a fiber-optic collimating lens, a UV 

band-pass filter, and two SO2 calibration cells. A subnotebook 

computer provides FLYSPEC power, control, and data collection through a USB port, and a small GPS receiver supplies 

location data for cosine corrections of traverses not perpendicular to the direction of plume travel.

The airborne instrument package also includes a type-K 

thermocouple, shielded from wind and direct sunlight, for 

measuring ambient air temperature and a chart recorder for 

in-flight logistical use. A pressure transducer mounted within 

the LI-COR analyzer provides measurements of ambient 

atmospheric pressure inside the unpressurized aircraft cabin. 

An onboard Rockwell GPS receiver continuously records the 

precise latitude, longitude, and altitude, so that the location 

of each measurement can be retrieved later for data processing. The data from all instruments except the FLYSPEC are 

recorded at 1-s frequency in a handheld PSC Falcon Model 

310 portable data-collection terminal. Figure 1 illustrates 

schematically the configuration of all the instruments of the 

airborne system. Figure 2 displays instrument placement in 

both helicopter and fixed-wing aircraft and the external helicopter mounting of a FLYSPEC.

SO2 SO2 H2S CO2

GPS

COMPUTER COMPUTER

CHART

RECORDER

DATA

LOGGER

COSPEC INTERSCAN LI-COR

Pressure Temperature

Location

FLYSPEC

ABUNDANCE CONCENTRATION

Figure 1. Airborne instrument package showing instruments used 

for measuring SO2 column abundances (FLYSPEC and COSPEC); 

instruments for measuring CO2, SO2, and H2S gas concentration 

(LI-COR and Interscan); sensors for measuring temperature, 

pressure, and location; and other components used during a typical 

airborne mission (computer, data logger, and chart recorder).
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 547

A

B

C

Figure 2. Gas-monitoring equipment installation for airborne 

surveillance, 2004–2005. A, FLYSPEC being attached to the strut 

of helicopter. The instrument’s narrow field of view allows it to be 

mounted near the body of the aircraft without blocking the light. 

B, Typical mounting arrangement of instrumentation in helicopter. 

C, Instrumentation placed in a fixed-wing aircraft. Note COSPEC 

panel (right) protruding from behind copilot’s seat. 

USGS photos by K.A. McGee (A, C) and M.P. Doukas (B).

Procedures and Methods

The gas measurements began on September 27, 2004. 

The five flights during the first week were primarily for 

reconnaissance purposes. At this stage of unrest, the volcano 

lacked a coherent gas plume but produced small fragmentary 

and ephemeral parcels of gas that—although detectable by 

the airborne system—were only useful for rough estimates 

of gas emission rates. Nevertheless, we were able to discover 

and follow the development of point sources of emerging gas 

emissions within the crater by mounting the direct-sampling 

instruments (LI-COR and Interscans) in a helicopter with 

an air intake tube attached to the lower left strut and making 

reconnaissance flights of the 1980–86 lava dome, the crater 

floor, and the crater walls (fig. 3).

After the first week of October 2004, a coherent plume 

developed and spilled over the crater rim, and accurate emission-rate determinations became feasible (fig. 4). We mounted 

and flew the direct-sampling instruments and the COSPEC V 

in a twin-engine aircraft configured for open-flow sampling 

of external air (Gerlach and others, 1997, 1999). The use of 

a twin-engine aircraft prevented contamination by combustion products from engine exhaust. Upward-looking COSPEC 

measurements taken while flying traverses beneath the plume 

perpendicular to the direction of plume transport provided SO2

column abundances across an orthogonal section of the plume 

(fig. 5A). Flying traverses across the plume at different altitudes, each normal to its transport direction, allowed in-plume 

measurements and profiling of entire orthogonal sections of 

the plume by the direct-sampling instruments (LI-COR and 

Interscans) (fig. 5B). Figure 6 illustrates a series of airborne 

traverses across the plume, showing where the Interscan 

analyzer detected SO2 in the orthogonal plume cross section. 

Later processing of these data into emission rates included 

corrections for curvature in the flight paths and for deviations 

from orthogonality. Contour maps of gas concentrations in the 

orthogonal plume cross section were produced using mapping 

software (Surfer v. 8).

By November 10, 2004, we made a transition to the use 

of a helicopter as the airborne platform for gas measurements, 

for these reasons: the availability of a helicopter for field work, 

the superior maneuverability of a helicopter for making plume 

measurements at the crater rim and within the crater, and the 

slower airspeed of the helicopter, which permits more measurements per meter across small plumes. Use of the helicop-
548 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A B

Figure 3. Shaded-relief elevation model of Mount St. Helens’ 

crater indicating the flight path of a gas-measurement flight in 

October 2004. Early gas flights in fall 2004 before the emergence 

of a discrete plume, such as the flight illustrated here, were 

geared toward detecting point sources of CO2, SO2, and H2S. 

View is to the southeast.

Figure 4. Typical gas plumes from Mount St. Helens early in the 2004–2005 eruption. A, Plume observed on October 27, 2004. View 

is to southeast. B, Plume on November 4, 2004. View is to south. Crater diameter is about 2.1 km. 

USGS photos by M.P. Doukas (A) and K.A. McGee (B).

ter also marked the transition to exclusive use of the FLYSPEC 

ultraviolet spectrometer for SO2 emission-rate measurements. 

Although it is possible to utilize the COSPEC in a helicopter, 

“chopping” of light to the COSPEC telescope by the rotor 

blades creates interference in the signal output. This requires 

selection of longer time constants to dampen the interference, 

which negates the advantageous spatial resolution of the helicopter (Caltabiano and others, 1992; Galle and others, 2002; 

Elias and others, 2006). The FLYSPEC, however, is relatively 

immune to interference by the rotor blades, and no special 

filtering or processing of the data is required. The switch from 

COSPEC to FLYSPEC did not introduce inconsistencies into 

the SO2 emission-rate dataset—side-by-side testing of these 

instruments shows they produce statistically equivalent results 

(Elias and others, 2006; Horton and others, 2006).

The relatively small field of view (2.5°) of the FLYSPEC 

allows it to be mounted on an external structural member close 

to the body of the helicopter and positioned to point vertically 

without loss of signal due to light obstruction (fig. 2A). We 

fed the USB signal and power cable through a side window 

of the helicopter to connect with the subnotebook computer. 

Prior to the plume measurements, we landed the helicopter at 

a calibration station high on the flank of Mount St. Helens and 

collected instrument spectra for the high and low calibration 

cells along with spectra for dark current and clear sky. Once the 

FLYSPEC was calibrated, we flew several traverses beneath the 

plume that were averaged to obtain the SO2 column abundances of the plume. We reduced the data with the FluxCalc 

software program written by staff at the Hawai‘i Institute of 

Geophysics and Planetology, University of Hawai‘i at Mānoa. 

The program compares absorption values of as many as nine 

peak-and-trough combinations in each sample SO2 spectrum to 

the two calibration cell spectra and then computes the column 

abundance of SO2.

The standard method for determining emission rates 

from airborne in-plume measurements by the direct-sampling 

instruments is based on the volcanic gas concentrations 

measured in the cross section through the plume normal to 

wind direction, the average plume pressure and temperature 

in that section, and the wind speed (Harris and others, 1981; 

Gerlach and others, 1997). Gerlach and others (1999) adapted 

the method by using orbital traverses around a volcano and 

correcting the data from curved plume cross sections approximately normal to wind direction. McGee and others (2001) 

further adapted the method by using H2S peaks as a guide 

for resolving the boundary between atmospheric-only CO2

and atmospheric plus volcanic plume CO2. In this study, we 

often used the location of SO2 peaks recorded by the Interscan analyzer as markers to distinguish plume CO2 from the 

atmospheric background CO2 because H2S was not always 

detected in the plume. Direct measurement of CO2 and SO2
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 549

A

B

Remote sensing

Plume profiling

Wind direction

Figure 5. Schematic flight patterns for plume monitoring. 

A, Flight pattern for remote measurement of SO2 column 

abundances in the plume. Flight traverses with the upwardlooking FLYSPEC or COSPEC spectrometer are made 

beneath the plume perpendicular to its direction of travel. B, 

Instruments that measure concentration and require direct 

sampling of the plume, such as LI-COR and Interscan, are 

flown directly through the plume, perpendicular to its travel 

direction, in a series of top-to-bottom profiles forming a 

vertical plume cross section.

Figure 6. Illustration 

of a downwind airborne 

measurement of the 

plume from Mount St. 

Helens on July 13, 2005, 

plotted on shaded-relief 

elevation model. View is 

to the southeast. Blue line 

represents GPS tracks of 

several traverses through 

the plume, with hotter 

colors (yellow, orange, 

and red) representing 

locations where SO2 was 

measured. Red arrow 

indicates wind direction; 

wind is from the north.

also allowed calculation of CO2/SO2 ratios at many locations 

in the plume.

Knowing the velocity of plume travel is essential in computing accurate gas-emission rates from either gas concentrations in a plume cross section or column-abundance measurements made beneath the plume. It is assumed that the speed of 

the plume as it moves downwind is the same as the ambient 

wind speed. We used the Doukas (2002) method of windspeed measurement, in which a GPS receiver records the rate 

and direction of drift of the aircraft while flying neutral wind 

circles at the altitude of the plume. This information allows 

determination of wind speed and wind direction at the altitude 

of the plume with a windspeed uncertainty of <10 percent.

 The equation of Gerlach and others (1997),

ECO2 = 0.457329(ASPCO2)/T, (1)

calculates the CO2 emission rate, ECO2, in t/d where A is the 

area of the plume cross section after corrections for curvature 

of orbital traverses (in m2), S is the average plume speed (in 

m/s), PCO2 is the average partial pressure of CO2 in the plume 

(in pascals (Pa), calculated from the product of average barometric pressure in the plume and the average mole fraction 

concentration of CO2 in the plume), and T is the average air 

temperature in the plume (in K). The constant (units: s t K 

Pa-1 m-3 d-1) includes the kilogram molecular weight of CO2

(0.04401 kg/mol), the universal gas constant (8.314510 Pa m3

mol-1 K-1), and the conversion factors 86,400 s/d and 103 kg/t. 

When calculating EH2

S, PCO2

is replaced by PH2

S , and the con-
550 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

stant is 0.354141 (McGee and others, 2001); when calculating 

ESO2, PSO2 and the constant 0.665665 are used.

Flight frequency varied because of a mix of factors, which 

included perceived flight safety, apparent necessity of the data 

for hazard assessments, rate of change of eruptive activity, 

and current and forecasted weather. Starting on September 27, 

2004, we made 11 flights in the first 3 weeks, 1 to 2 flights per 

week during the next 5 weeks, and 1 or 2 flights per month 

thereafter throughout 2005, except during December 2005 

when weather conditions did not permit access to the plume. 

Flight dates and aircraft platforms deployed (helicopter or 

fixed-wing) are recorded in table 2.

Description of Data and Statistical Tests

Gas emission-rate data, which form the backbone data 

of this report, tend to have skewed rather than symmetrical 

distributions. The skewed distributions are a reflection of two 

fundamental properties of gas emission-rate data: they can 

take on only positive values, and their standard deviations are 

typically large—commonly about as large or larger than their 

means. When distributions are skewed, means and standard 

deviations may no longer adequately describe population distributions, and interpreting these statistics in terms of a normal 

distribution can produce a misleading picture. It is preferable 

instead to report the median for skewed distributions (Glantz, 

2005); we also report upper and lower percentile information 

as appropriate to give an indication of the dispersion of values 

in the sampled population.

If populations are skewed from normal distributions, 

parametric statistical tests may become unreliable, and 

nonparametric or distribution-free statistical tests should be 

used (Glantz, 2005). Therefore, we used statistical software 

(SigmaStat v. 3.1) to calculate the nonparametric MannWhitney rank-sum test (MWRST), instead of the parametric 

Student’s t test, to test for significant differences in two groups 

of data. The MWRST can be used to evaluate significant difference in the medians of two groups, significant difference 

in the underlying population distributions of two groups, and 

significant difference between two groups caused by random 

sampling variations. We report the probability level (P) of 

MWRST results. Following the tradition of most physical-science research, the critical significance level for P is 5 percent 

(P = 0.05) in this report; that is, P≤0.05 indicates a significant 

difference. Several of the statistical calculations also included 

Kolmogorov-Smirnov normality testing (SigmaStat v. 3.1).

Results and Observations

Emission Rates, Cumulative Emissions, and 

CO2/SO2 Ratios

The measured CO2, SO2, and H2S emission rates of the 

2004–5 eruption spanned the period from September 27, 

2004, to November 22, 2005 (fig. 7, table 2). Emission rates 

frequently were below detection limits before October 7, 

2004 (fig. 7, table 2). Time series of the 24 CO2 and 28 SO2

emission rates measured from October 7, 2004, until the last 

2005 measurement on November 22 were notably low and 

variable. Emission rates ranged from 2,415 t/d to 146 t/d for 

CO2 and from 240 t/d to 14 t/d for SO2; both CO2 and SO2 had 

low medians: 655 t/d for CO2 and 72 t/d for SO2. Less than 5 

percent of the CO2 emission rates were >2,000 t/d, about 80 percent were <1,000 t/d, more than 40 percent were <500 t/d, and 

about 25 percent were <250 t/d. Of the SO2 emission rates, only 

7 percent were >200 t/d, almost 70 percent were <100 t/d, and 

40 percent were <50 t/d. At times, emission rates fell to <150 

t/d for CO2 and to <30 t/d for SO2. Emission rates of H2S were 

always low (<10 t/d). Multiple traverses under the plume gave 

comparable SO2 emission rates, so the observed variation was 

not simply from puffing during times of measurement.2

Cumulative CO2 and SO2 emissions—calculated from 

the areas under emission-rate time-series plots—indicate total 

outputs of 231,150 t CO2 and 29,700 t SO2 from October 7, 

2004, through November 22, 2005 (fig. 8; table 2). Weighted 

emission rates—calculated from cumulative emissions on 

measurement dates (fig. 8; table 2) divided by the number of 

days since October 7, 2004—are like multiday averages and 

thus smoother than measured emission rates. Although less 

variable, weighted emission rates also have low medians—830 

t/d for CO2 and 99 t/d for SO2. However, these median rates 

are somewhat higher than the medians of measured values. 

Nevertheless, statistical tests confirm that no significant dif2

 Note added in proof: Low emission rates of 150–300 t/d CO2, 5–35 t/d 

SO2, and < 0.1 t/d H2S persisted during 12 gas-monitoring flights from 2006 

until the eruption paused in late January 2008.

CO2

SO2

10

100

1,000

10,000

1

EMISSION RATE, IN TONS PER DAY

Sep Dec Jan Apr Jul Oct

2004 2005

Figure 7. Time-series plots of CO2 and SO2 emission rates from 

September 27, 2004, to November 22, 2005 (table 2). The open 

symbols for some earlier measurements between September 27 

and October 1, 2004, indicate estimated detection limit values of 10 

t/d (CO2) and 1 t/d (SO2).
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 551

Date Platform

CO2

(t/d)

SO2

(t/d)

H2S

(t/d)

 CO2

 (t)

SO2

(t)

09/27/2004 Heli 50 1 1 0.1 1 nd nd

09/29/2004 Heli 10 1 1 1 0.1 1 nd nd

09/30/2004 Heli 50 1 1 0.1 1 nd nd

10/01/2004 Heli 10 1 1 1 0.1 1 nd nd

10/03/2004 Heli 50 1 1 0.1 1 nd nd

10/04/2004 Heli 140 1 1 0.4 nd nd

10/07/2004 FX 2,415 115 8 2,415 115

10/11/2004 FX 1,222 70 4 7,274 370

10/13/2004 FX 186 120 6 8,682 560

10/14/2004 FX 710 148 0.1 1 9,130 694

10/20/2004 FX 652 121 8 13,216 1,501

10/27/2004 Heli 1,060 240 8 19,208 2,764

11/04/2004 FX 332 63 0.1 1 24,776 3,976

11/10/2004 Heli nd 107 6 nd 4,486

11/12/2004 Heli 981 138 nd 30,028 4,732

11/20/2004 Heli 1,053 177 6 38,164 5,993

12/17/2004 Heli 914 221 5 64,718 11,370

01/03/2005 Heli 887 76 0.1 1 80,027 13,893

01/24/2005 Heli nd 78 11 nd 15,509

02/08/2005 Heli 718 70 0.3 108,917 16,619

02/25/2005 Heli 211 32 0.6 116,814 17,485

03/10/2005 Heli 149 30 0.3 119,154 17,888

04/21/2005 Heli nd 40 0 nd 19,347

04/26/2005 Heli 146 94 0 126,086 19,681

05/25/2005 Heli 325 33 0 132,916 21,522

06/09/2005 Heli nd 50 nd nd 22,145

07/13/2005 Heli 485 40 nd 152,760 23,675

07/27/2005 Heli 594.3 51 0.1 1 160,316 24,312

08/12/2005 Heli 174 14 0.1 1 166,462 24,832

08/31/2005 Heli 198 22 0.1 1 169,996 25,174

09/28/2005 Heli 789 45 0.1 1 183,814 26,112

10/18/2005 Heli 658 53 0.3 198,284 27,092

11/09/2005 Heli 1,353 88 0.1 1 220,405 28,643

11/22/2005 Heli 300 75 0.1 1 231,150 29,702

1

Estimated detection limits for CO2, SO2, and H2S emission rates.

Table 2. Emission rates and cumulative emissions for CO2, SO2, and H2S 

during 2004–2005, Mount St. Helens, Washington.

[Emission rates are measured rates used in calculations of this study; significant figures 

were exaggerated to reduce roundoff errors. Cumulative emissions were determined as 

described in the text. Platform indicates aircraft: FX, fixed-wing; Heli, helicopter. t/d, 

metric tons per day; nd, not determined.]
552 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

CO2

SO2

Sep Dec Jan Apr Jul Oct

0

50

100

150

200

250

CUMULATIVE EMISSION, IN KILOTONS

2004 2005

ference exists between the medians of measured and weighted 

CO2 and SO2 emission rates and the populations they sample 

(P >0.05).

The 24 contemporaneous CO2 and SO2 emission rates 

measured from October 7, 2004, to November 22, 2005, have 

a median molar CO2/SO2 ratio of ~12; about half of the CO2/

SO2 ratios are between 7 and 18. The corresponding weighted 

emission rates have a median CO2/SO2 of ~10 and about half 

the CO2/SO2 ratios are between 9 and 11. Statistical tests 

confirm there is no significant difference between CO2/SO2

ratios calculated from measured and weighted emission rates 

(P >0.05). Total emissions for the entire period—231,150 t 

CO2 and 29,700 t SO2—indicate a CO2/SO2 of ~11. Thus, we 

employ a grand median CO2/SO2 ratio of 11±1 in this report.

Early Degassing

From the start of airborne gas measurements on September 27, 2004, throughout the period of unrest before the 

eruption began on October 1 and for almost a week thereafter, a volcanic plume was either absent or present as small 

fragmentary and ephemeral parcels of gas, yielding only 

rough gas-emission rates (table 2). Flights during this period 

served mainly as reconnaissance surveys for gas-emitting 

sources on the 1980–86 dome, the crater floor, and the crater 

walls (fig. 3) and led to the discovery and observation of point 

sources of gas emission that, over time, established a record 

of emergent degassing. The record began with SO2 and H2S 

concentrations at detection limits at all locations and CO2

concentrations at ambient atmospheric levels at most locations 

(fig. 9), although sporadic CO2 sources occurred with minor 

peaks <1.5 ppm above ambient levels. Similar results persisted 

through September 30. After the first steam-and-ash explosion 

at the start of the eruption on October 1, new CO2 concentration peaks above ambient levels appeared—one new source 

Figure 8. Time-series plots of cumulative CO2 and SO2 emissions 

from October 7, 2004, to November 22, 2005, when emission rates 

for both gases were above detection limits (fig. 7; table 2).

on the northwest face of the 1980s dome produced an abovebackground peak of 36 ppm—but SO2 and H2S still remained 

largely at noise levels. There was some evidence, however, 

of the presence of H2S—one minor peak was observed, and 

some observers in the air and on the ground downwind of the 

volcano reported the odor of H2S. 

 The number of fumarole sources increased significantly 

by October 2, as did the number of above-ambient CO2 concentration peaks—one reaching up to 16 ppm—while SO2 and 

H2S continued to be mostly absent, although field crews again 

reported H2S odors (fig. 10). A further increase in the number 

of fumaroles and a broadening of CO2 anomalies along with 

coincident low H2S peaks (<0.02 ppm) appeared in the record 

on October 3. By October 4, CO2 anomalies were as broad as 

a kilometer or more across, and SO2 and H2S formed coinci0

0.015

0.030

0

0.015

0.030

370

375

373 CO2

SO2

H2S

0.0 2.0 4.0 6.0 8.0 10.0 12.0

 KILOMETERS ALONG FLIGHT PATH

GAS CONCENTRATION,

IN PARTS PER MILLION

0

0.015

0.030

0

0.015

0.030

0.0 1.0 2.0 3.0 4.0 5.0 6.0

372

369

366

KILOMETERS ALONG FLIGHT PATH

CO2

SO2

GAS CONCENTRATION,

IN PARTS PER MILLION

H2S

Figure 9. Concentrations of CO2, SO2, and H2S measured 

along a helicopter flight path in Mount St. Helens crater on 

September 27, 2004. Concentrations of CO2 are similar to ambient 

atmospheric levels; concentrations of SO2 and H2S are at sensor 

detection limits.

Figure 10. Concentrations of CO2, SO2, and H2S measured along 

a helicopter flight path in Mount St. Helens crater on October 2, 

2004. Concentrations of CO2 show sporadic spikes above ambient 

atmospheric levels; concentrations of SO2 and H2S remain at 

sensor detection limits. The clipped CO2 peak rises to 382 ppm.
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 553

0

0.015

0.030

0.0 2.0 4.0 6.0 8.0 10.0 12.0

372

369

366

0

0.015

0.030

KILOMETERS ALONG FLIGHT PATH

CO2

SO2

H2S

GAS CONCENTRATION,

IN PARTS PER MILLION

0

0.015

0.030

0

0.05

0.10

0 2 4 6 8 10 12

366

367

368 CO2

SO2

H2S

KILOMETERS ALONG FLIGHT PATH

GAS CONCENTRATION,

IN PARTS PER MILLION

0 0.5 1 1.5 2 2.5 3 3.5 4 4.5 5

2.1

2.3

2.5

Background

0.15

1

1.5

2

CO2

(ppm)

BREADTH OF PLUME, IN KILOMETERS FROM ARBITRARY DATUM

ALTITUDE, IN KILOMETERS

ABOVE SEA LEVEL

Figure 11. Concentrations of CO2, SO2, and H2S measured along 

a helicopter flight path in Mount St. Helens crater on October 

4, 2004. Concentrations of CO2 show broad anomalies above 

ambient atmospheric levels; concentrations of SO2 and H2S show 

sharply coincident anomalies spiking above sensor detection 

limits. Note that SO2 and H2S anomalies have concentration 

levels of roughly similar magnitude.

Figure 12. Concentrations of CO2, SO2, and H2S measured 

along a helicopter flight path in Mount St. Helens crater on 

October 11, 2004. Concentrations of CO2 show a broad anomaly 

above ambient atmospheric levels; concentrations of SO2 and 

H2S show coincident anomalies. Note that concentration levels 

of the SO2 anomaly now are clearly greater than those of H2S.

Figure 13. Concentrations of CO2 in a vertical cross section of a coherent gas plume from Mount St. Helens on November 12, 2004, 

1.7 km downwind of source. Bar on right shows scale for concentration of CO2 above local atmospheric background. Plume cross 

section gives a CO2 emission rate of 980 t/d.

dent anomalies of roughly equivalent concentration levels (fig. 

11). Wet degassing prevailed at this time as large gas bubbles 

were seen streaming through and ejecting from pools of water, 

which often contained chunks of floating ice that calved in 

from the adjacent melting glacier. By October 11, SO2 clearly 

dominated H2S as temperatures rose, pools of water dried up, 

and steaming increased greatly (fig. 12). Coherent plumes 

began to form regularly on or about October 7 and spilled over 

the crater rim (fig. 4), permitting acquisition of reliable emission rates by remote SO2 column-abundance measurements 

and direct measurements of CO2, SO2, and H2S gas concentrations (fig. 13).

Composite Degassing Sources

Mount St. Helens continued to produce buoyant plumes 

throughout the remainder of 2004 and all of 2005. In addition 

to the principal degassing site from the vent area on the north 

side of the new dome, several additional sites degassing both 

CO2 and SO2 were present at various locations on both old 

and new lobes of the growing new dome. By summer 2005, 

gas emissions were declining, and airborne gas measurements began to be conducted closer to the gas sources. These 

measurements provided more spatial resolution of gas concentrations within the proximal plume and revealed a minor 

source of the CO2 within the plume coming from the 1980s 

dome. Since 2005, we have consistently detected CO2 coming 

from the southwest side of the 1980s dome, just north of the 

vent area for the new dome. In addition, the older dome hosts 

a smaller CO2 source on its east shoulder and several scattered 

point sources of CO2. The spatial separation of new-dome and 

old-dome gas sources is best developed when winds are from 

the east or west. The degassing from the 1980s dome involves 

little or no SO2. Nearly all of the SO2 in the plume, and the 

majority of CO2, originates from the hotter vents associated 

with the new dome. Under appropriate wind conditions, airborne traverses through the plume show CO2 clearly degassing 

from both domes, whereas SO2 is virtually absent from the 
554 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1980s dome (fig. 14). Figure 14 shows CO2 concentrations on 

a single plume traverse; most plume traverses show CO2

concentrations are much higher over the new dome.

CO2 Plume from Loowit Springs

During some of the airborne gas-measurement flights, a 

small CO2 plume was observed that did not emanate from the 

new or the old dome. The source of this plume proved to be 

the Loowit springs area (Thompson, 1990), north of the 1980s 

dome and near the mouth of the crater, as determined by a 

series of low helicopter traverses on June 9, 2005, when winds 

from the north or northwest were favorable for distinguishing it. Orbital profiling around the upper part of the Loowit 

drainage defined the extent of the plume and quantified the 

CO2 emission rate at 27 t/d (fig. 15), although the exact spring 

or springs involved could not be distinguished; neither SO2

nor H2S were detected. The spring discharge of magmatic CO2

as inorganic carbon (aqueous CO2 and HCO3

–

 in about equal 

amounts) dissolved in Loowit waters and other thermal features 

draining the crater is ~22 t/d, which is similar to the emission 

rate of CO2 gas around Loowit (Bergfeld and others, this volume, chap. 25). It is not known if the Loowit springs plume is a 

result of the current eruption or a longer-term feature.

-0.05

0.00

0.05

0.10

0.15

0.20

-0.5

0

0.5

1

1.5

2

NEW DOME PLUME OLD DOME PLUME

CO2, IN PARTS PER MILLION

SO2, IN PARTS PER MILLION

DISTANCE ALONG TRAVERSE, IN KILOMETERS

0.0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0 4.5 5.0

MSH November 9, 2005

CO2-SO2 traverse 6 S–N view to West

Wind in your face

South end traverse North end traverse

Discussion

The key fact about the CO2 and SO2 emission rates is 

that they are generally low. This fact was evident early on and 

motivated our hypothesis that the CO2 and SO2 emissions of 

the 2004–5 eruption were derived from source dacite depleted 

in excess (that is, exsolved) volatiles at depth (Gerlach and 

others, 2005), herein termed “flat” magma. The corollary 

hypothesis—that gas emissions from flat dacite are expressly 

vulnerable to scrubbing, particularly during the unrest and 

early stages of the eruption—follows from the depletion in 

excess volatiles. The flat magma and gas scrubbing hypotheses 

underlie much of the discussion that follows.

Comparisons with 1980s Emission Rates

Comparisons of the 1980s and 2004–5 passive emission 

rates of SO2 and CO2 support the interpretation that the current 

eruption involves flat magma, possibly left over in the reservoir since the last lava-dome eruption in 1986. The 2004–5 

SO2 emission rates (fig. 7; table 2) are not significantly different from 1980s SO2 emission rates during the period from 

October 18, 1980, to December 9, 1986 (fig. 16), according 

to the MWRST (P ≥0.05). This time interval matches closely 

Figure 14. Airborne traverse of plume on November 9, 2005, showing above-background CO2 concentration (blue) and SO2

concentration (red). Wind direction from nearly due west allows comparison of plume data from the new dome complex 

with the older 1980–1986 dome. Traverse flown approximately 2.3 km downwind of central source area. The older dome is 

discharging CO2 but little or no SO2.
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 555

BREADTH OF PLUME, IN KILOMETERS FROM ARBITRARY DATUM

ALTITUDE, IN KILOMETERS

ABOVE SEA LEVEL

0.15

1.5

3.5

5.5

1.7

1.75

0 0.5 1.0

CO2

(ppm)

Background

the period of 17 lava-dome eruptions and stable dome growth 

that produced the 1980s dome. October 18, 1980, is the date 

of the first SO2 emission-rate measurement after the start of a 

lava-dome eruption on October 16, 1980—the first lava-dome 

eruption of the 1980s to produce a dome that survived. The last 

lava-dome eruption of the 1980s was on October 21, 1986, and 

the last significant SO2 emission rate at the end of this eruption 

was 80 t/d on December 9, 1986. Of the 113 SO2 emission-rate 

measurements made after this date, until measurements ceased 

on September 6, 1988, 95 percent were <25 t/d and 70 percent 

were <3 t/d—barely above detection limits. We did not include 

these data in the MWRST analysis in order to avoid confounding the emission rates associated with the 1980s lava-dome 

eruptions with emission rates of the protracted period of residual 

degassing at the end of the 1980s eruption cycle. The1980s 

dataset for the calculations started with the data from December 

9, 1986, and was extended back in time by incrementally adding 

earlier SO2 emission rates. Along this time line, the probability 

level decreased steadily from P >0.05 and converged on the 

critical significance level (P = 0.05) when the October 18, 1980, 

data were added to the dataset. When still earlier 1980 data 

were included, P dropped abruptly below 0.05, indicating that 

the difference between the two groups of data became robustly 

significant. Thus, between October 18, 1980, and December 

9, 1986, the two groups of data sample a similar population, 

and the differences between them—for example, median SO2

emission rates of 72 t/d versus 100 t/d—can be explained 

by random sampling variations. However, we stress that the 

MWRST analysis operates only on the rates of the two groups 

of data without regard for the associated dates. This is important 

because if the dates are considered, it is evident that although 

the 10/18/1980–12/9/1986 SO2 emission rates may share a common population with 2004–5 SO2 emission rates, they are not 

randomly distributed in time, as indicated by the conspicuous 

skewing of their higher rates to earlier dates (fig. 16).

Plotting the SO2 emission rates of the 1980s and 2004–5 

together as time series on the same scale reveals that the emission rates of the 2004–5 eruption appear similar to those of 

the 1980s lava-dome eruptions back to at least late 1981 (fig. 

16). MWRST comparisons back to late 1981 produce large 

P values indicating virtually no likelihood of a significant 

difference in the SO2 emission rates of the 1980s and 2004–5 

lava-dome eruptions; for example, comparisons back to October 1, 1981, give a P value of 0.7. Such differences as exist in 

the patterns of the two time-series plots are readily explained. 

The generally lower peak values of the 2004–5 eruption probably are related to the more-crystalline magma and steadier 

character of the present eruption; these properties act to curb 

extrusive surges that frequently caused SO2 emission-rate 

peaks during the 1980s lava-dome eruptions. The 2004–5 

emission rates appear blockier and less spiky on the graph, 

partly because of the steadier nature of the present eruption, 

but also because the steadier activity itself (along with tighter 

funding) led to less frequent monitoring flights—four to five 

times less frequent than in the 1980s. The fact that the current eruption has lasted considerably longer than the 1980s 

lava-dome eruptions also contributes to the blockier pattern 

of its SO2 emission rates. The annualized 2004–5 SO2 outputs also bear a noteworthy similarity to annual SO2 outputs 

that occurred during the later years of the 1980s lava-dome 

eruptions. The 14-month, 30,000-t SO2 output of the 2004–5 

eruption (table 2) indicates an annualized output of 26,000 t, 

which is equivalent to the total Mount St. Helens SO2 output 

for 1984 (Gerlach and McGee, 1994)—a year in which there 

were three lava-dome eruptions over a 6-month period. The 

11-month output for 2005 of 18,000 t (computed from table 

2 data) annualizes to an output of 19,600 t—comparable with 

the 1986 SO2 output of 17,000 t (Gerlach and McGee, 1994). 

Thus, the MWRST analyses, the time-series patterns of SO2

emission rates, and the annualized SO2 outputs strongly suggest that the 2004–5 SO2 emissions are similar to those of the 

1980s lava-dome eruptions, possibly back as far as late 1981 

or late 1980.

The CO2 emission-rate measurements of the 1980s are 

restricted to the period between July 1980 and August 1981 

(fig. 17; McGee and Casadevall, 1994); the measurements 

were terminated after August 1981 because the detection limit 

of the older measurement technique was reached. Applications 

of the MWRST to these data and the 2004–5 CO2 emissionrate data (fig. 7, table 2), following the procedure described 

Figure 15. Vertical cross section of CO2 gas plume 250 m downwind of Loowit springs area on June 9, 2005. The CO2 emission 

rate computed from these data was 27 t/d. Scale bar on right shows CO2 concentration above local atmospheric background.
556 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

above for comparing SO2 emission rates, indicates that the 

2004–5 data are significantly different from the July 1980–

May 1981 data (P ≤0.05). However, the 2004–5 data and the 

June–August 1981 data are not different and appear to be sampling the same population (P >0.05). The similarity probably 

would have persisted beyond August 1981 had measurements 

been made by a more sensitive technique. Plotting the CO2 

emission rates of 1980–81 and 2004–5 together as time series 

on the same scale reinforces the MWRST analysis by revealing a pattern in which the latter CO2 emission rates appear as 

an extension of the lower CO2 emission rates of June–August 

1981 (fig. 17). As in the case of the SO2 emission rates, differences in magma physical properties and frequency of measurement account for the blockier, less spiky, and generally lower 

peak values of the 2004–5 CO2 emission rates.

Thus, the time-series comparisons and MWRST results 

for CO2 and SO2 emission rates support a flat 2004–5 magma 

depleted in excess volatiles and producing generally low CO2

and SO2 emissions like those of the 1980s lava-dome eruptions. The flat magma is distinctly different from the gas-rich 

magma that produced the much higher CO2 and SO2 emissions earlier in 1980 (figs. 16, 17; Gerlach and McGee, 1994). 

The similarity to the gas emissions of the 1980s lava-dome 

eruptions suggests that the flat magma of the current eruption 

could involve leftover reservoir magma from 1986.

Comparison with 1980s CO2/SO2 Ratios

Contemporaneous measurements of CO2 and SO2 emission rates in the 1980s allow calculation of CO2/SO2 ratios 

on 117 days during a 418-day period from July 6, 1980, to 

August 28, 1981. The median CO2/SO2 ratio is 8, but there is 

considerable variation—25 percent of the ratios are ≤5 and 

25 percent are ≥11. As noted earlier, measured and weighted 

CO2 and SO2 emission rates and total emissions of the current 

eruption imply a somewhat higher grand median CO2/SO2

ratio of 11±1. The MWRST indicates that the 117 1980–1981 

CO2/SO2 ratios and the 24 2004–5 CO2/SO2 ratios—whether 

calculated from measured or weighted emission rates—are 

associated with significantly different populations (P <0.05). 

Taking these results at face value would imply a real difference in CO2/SO2 ratios of the older and current gases, but we 

show below that the higher CO2/SO2 ratios of the current gases 

are caused by scrubbing.

Scrubbing in the 2004–2005 Eruption

Doukas and Gerlach (1995) drew attention to the unusually weak SO2 emission rates that both preceded and followed 

eruptions at Crater Peak on Mount Spurr volcano, Alaska, in 

1992 and noted the exceptionally strong and persistent H2S 

odor in the Crater Peak plume during repose periods. They 

argued that liquid water present at or beneath the surface 

interacted with ascending magmatic gases, scrubbing SO2

and other strongly acidic gases (HCl, HF) but impacting 

the weakly and moderately acidic gases (CO2 and H2S) less 

severely. The principal mechanisms for SO2 scrubbing by 

liquid water are the hydrolysis reactions:

4SO2(g) + 4H2O(l) = H2S(g) + 6H+

(aq) + 3SO4

-2

(aq) , and

3SO2(g) + 2H2O(l) = S(l,s) + 4H+

(aq) + 2SO4

-2

(aq).

These equilibria shift strongly to the right below 400°C, mainly 

converting SO2 into dissolved sulfate as sulfuric acid. Some 

H2S gas and native sulfur also are generated, which explained 

the intense H2S odor at Crater Peak. Detailed thermochemical 

modeling (Symonds and others, 2001) confirms the significant 

scrubbing of magmatic SO2, HCl, and HF by contact with 

surface, ground, and hydrothermal waters of volcanoes and indicates that scrubbing of CO2 and H2S also becomes significant at 

low gas-to-water ratios (<0.01, mass basis).

Carbon isotopes of dissolved inorganic carbon in 

hydrothermal, ground, spring, and stream waters reveal the 

scrubbing of magmatic CO2 in the vicinity of several Cascade 

Range volcanoes: Lassen Peak, Mount Shasta, Crater Lake, 

Mount Bachelor, Broken Top, Three Sisters, Belknap Crater, 

and Mount Jefferson (Rose and Davisson, 1996; James and 

others, 1999; Evans and others, 2004). Scrubbing of potentially coexisting, more acidic magmatic gases (SO2, H2S, 

HCl, and HF) also probably occurs. Copious ground-water 

recharge in the High Cascades and subsequent downward 

and lateral flow (Ingebritsen and others, 1989, 1992, 1994) is 

probably chiefly responsible for the scrubbing and provides 

a basis for a conceptual model of scrubbing throughout much 

of the Cascade Range. Scrubbing processes may, therefore, 

be responsible for the currently absent to negligible emission 

rates of CO2, SO2, and H2S at several Cascade Range volcanoes we have investigated with airborne measurements since 

the late 1990s (table 1). Although some of these volcanoes 

(for example, Mount Adams and Mount Jefferson) may simply be dormant, it is likely that much of the CO2, SO2, H2S, 

HCl, and HF degassed from magmas at most of these volcanoes is captured by hydrothermal systems or ground water 

and is thus unavailable to form gas emissions at volcanic 

centers. Boiling of hydrothermal fluid containing captured 

magmatic gases may give rise to persistent degassing of CO2

and H2S (but not SO2 because of hydrolysis to dissolved 

sulfate) at some volcanoes—Mount Baker and Lassen Peak 

being likely examples (table 1).

Symonds and others (2001) applied their thermochemical scrubbing models to several eruptions, including the 1980 

eruption of Mount St. Helens, which they say illustrates SO2

scrubbing in the early stage of an eruption. Low SO2 emission rates—from less than 10 t/d to 48 t/d—during the period 

of phreatic eruptions and nonexplosive degassing that began 

on March 27, 1980, continued for 52 days up to the May 18 

climactic explosion. Observations supporting early scrubbing 

of SO2 include reports by COSPEC investigators of persistent 

H2S odor in the crater at times when SO2 emissions rarely 

exceeded 10 t/d (Stoiber and others, 1980). Emission rates of 
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 557

1,000

500

0

1,500

2,000

2,500

3,000

3,500

SO2 EMISSION RATE, IN TONS PER DAY

1980–1986 lava-dome eruptions

1980 1981 1982 1983 1984 1985

YEAR

1986 1987 1988 2004 2005

10,000

5,000

0

15,000

20,000

25,000

1980 1981 2004

YEAR

2005

CO2 EMISSION RATE, IN TONS PER DAY

1980–1981 lava-dome eruptions

Figure 16. Time-series plots of all SO2 emission rates measured during the 1980s (McGee and 

Casadevall, 1994) and during the 2004–2005 eruption. Scale break after last 1980s measurements in 1988 

separates the two time series. Vertical lines indicate the starting dates of the 1980–1986 lava-dome 

eruptions. All 1980s and 2004–2005 time-series data are passive emission rates, except for the first 1980 

data point—2,600 t/d measured in the May 25, 1980, explosive eruption. The May 18, 1980, explosive 

emission rate of 4,000,000 t/d is not included (Gerlach and McGee, 1994; Symonds and others, 2001).

Figure 17. Time-series 

plots of all CO2 emission 

rates measured during 

the 1980s (McGee and 

Casadevall, 1994) and 

during the 2004–2005 

eruption. Scale break after 

the last measurements in 

1981 separates the two 

time series. Vertical lines 

indicate starting dates of 

the 1980–1981 lava-dome 

eruptions. All 1980s and 

2004–2005 time-series data 

are passive emission rates. 

The 1980 data begin with 

the 12,300 t/d CO2 emission 

rate measured on July 6.
558 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

passively degassed SO2 increased after the May 18 eruption 

and after the second explosion on May 25 but still remained 

relatively low (130–260 t/d). Twenty days later, on June 6, SO2

emission rates jumped to 860 t/d and remained high (>600 

t/d) until the end of October 1980 (fig. 16), finally attaining 

levels comparable to those of other erupting arc volcanoes and 

signaling the end of significant scrubbing. Note that the July 

6, 1980, to August 28, 1981, CO2/SO2 ratios discussed above 

represent postscrubbing values.

Symonds and others (2001) also inferred that SO2 scrubbing was active again during the period of seismic unrest at 

Mount St. Helens in the summer of 1998 when, as described 

above, an airborne survey on June 22 measured a CO2 emission rate of 1,900 t/d (table 1), while concurrent COSPEC 

measurements failed to detect SO2. These results are consistent 

with the degassing and scrubbing of SO2 from an intrusion 

at a time when melting of snow from a permanent snowfield 

growing in the crater since 1986 (Anderson and others, 1998) 

helped create a wet edifice. Thus, an unknown amount of 

degassed CO2 may also have been scrubbed at this time.

Scrubbing played a major role in the period of unrest 

and the early stages of the current eruption. We interpret the 

measurements during the September 27–30, 2004, period of 

insignificant gas emission with little or no detectable SO2 or 

H2S and with CO2 largely at atmospheric levels (fig. 9) to be 

the result of fairly complete gas scrubbing at low gas-to-water 

mass ratios <0.01 (Symonds and others, 2001). Sealing of 

rock permeability in the old dome and the upper part of the 

1980s conduit may also have played a role (Moran, 1994). It 

is likely, however, that scrubbing would soon have dominated 

sealing as a deterrent to gas emission. The concurrent shallow 

seismicity and deformation would tend to reestablish permeability. High seismicity was well established by September 

26 and increased significantly on September 28 at shallow 

levels in and below the old lava dome (Moran and others, 

this volume, chap. 2). Radial fractures were visible in Crater 

Glacier adjacent to the old dome by September 26, and an 

area of deformation south of the dome was clearly evident on 

oblique photos on September 30 (Schilling and others, this 

volume, chap. 8). Furthermore, the 2004 unrest and eruption 

followed a period with an unusually large potential for groundwater recharge by surface water. No August–September 

interval since the cessation of dome-building eruptions in 1986 

has had heavier rainfall than in 2004 (http://www.wcc.nrcs.

usda.gov/snotel, last accessed March 15, 2008). The growth 

of Crater Glacier since 1986 provided increased storage of 

water available for release and ground-water recharge into 

late summer. Moreover, the earlier detection of sharp CO2

concentration spikes on October 1 and 2, while SO2 and H2S 

remained at detection limits (fig. 10), strongly suggests scrubbing, rather than sealing, as the main cause restricting earlier 

gas emissions, since water does not scrub CO2 as effectively 

as it scrubs SO2 and H2S. The subsequent appearance of wet 

degassing with large gas bubbles ejecting through pools of 

water obviously involved scrubbing. However, the subequal 

concentrations of SO2 and H2S at this time (fig. 11) reflected 

incomplete scrubbing of SO2, which is likely because rapid 

transport in large bubbles restricted its interaction with water, 

thus allowing some of it to escape the pools. Wet degassing 

increasingly gave way to dry degassing, and SO2 became the 

major sulfur gas (fig. 12). As temperatures rose and steaming 

increased, rock adjacent to the invading magma progressively 

dried out in the days before the initial emergence of a lava 

spine from the deforming area south of the 1980s dome on 

October 11. This date probably marks the end of significant 

scrubbing in the early stages of the eruption.

The composite degassing sources from the 1980s and 

2004–5 domes bear the mark of gas scrubbing (fig. 14). Nearly 

all of the SO2 in the plume comes from the hotter new dome. 

Degassing from the colder old dome involves minor CO2 but 

little or no SO2. We speculate that the source of gas at depth is 

the same for both domes but that the gas feeding the vents on 

the old dome is scrubbed of its SO2 since water is more likely 

to be present within and beneath this colder dome. Edmonds 

and others (this volume, chap. 27) measured several gases, 

including SO2 and HCl, in hot vent emissions on the new 

dome by remote Fourier-transform infrared spectroscopy from 

high on the east rim of Mount St. Helens. If our hypothesis is 

correct, both SO2 and HCl—being strong acids—should be 

scrubbed in emissions from the old dome. Unfortunately, confirmation of concurrent scrubbing of both SO2 and HCl at the 

old dome by similar measurements was not possible, owing to 

the absence of hot infrared sources.

The 24 CO2/SO2 ratios calculated from cumulative CO2

and SO2 emissions (fig. 8; table 2) show a striking pattern 

when plotted as a time series (fig. 18). The ratios decrease 

regularly and precipitously from October 7 until October 27, 

2004, when they make a abrupt transition to nearly constant 

CO2/SO2 values of 9±1, close to the median value of 8 noted 

above for the 1980–81 CO2/SO2. They remain in this range 

(9±1) until September 28, 2005, after which subsequent 2005 

values drift to CO2/SO2 values as high as 11.3.

We interpret the sharp decline in CO2/SO2 (fig. 18) to 

reflect the drying out of the shallow conduit and crater-floor 

rock adjacent to the invading dacite and the reduction of SO2

scrubbing that dominated the early part of the 2004–5 eruption. 

However, cumulative emissions, although smoothing out shortterm variations, can lag in their response to significant changes. 

In this case, measurements suggest early SO2 scrubbing largely 

ceased on or about October 11, 2004, as discussed above, 

approximately two weeks before the October 27 date suggested 

by the cumulative emission data (fig. 18). After October 27, 

the CO2/SO2 ratio of cumulative emissions stabilized within the 

range of 9±1 for many months, grazing the median 1980–81 

CO2/SO2 value of 8 (fig. 18). We interpret this range of values 

to represent the CO2/SO2 of nonscrubbed gas emissions of the 

current eruption, although it may be slightly high because of 

SO2 scrubbing of gases from the old dome. The CO2/SO2 of the 

cumulative emissions rises above this range again after September 28, 2005 (fig. 18), presumably in response to renewed 

scrubbing. Because of the lag effect, the return to scrubbing 

may have started somewhat earlier. We have no data suggesting 
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 559

9/28/05

10/7/04

10/27/04

10

5

0

15

20

25

30

Sep Oct Nov Dec

2004 2005

Jan Feb Mar Apr May Jun Jul Aug Sep Oct Nov

CO2 / SO2, IN MOLES PER MOLE

Nonscrubbed CO2/SO2 range,

2004–2005

Figure 18. Time-series plot of molar CO2 /SO2 ratios (dots) calculated from cumulative 2004–2005 emissions 

(fig. 8; table 2). Solid and dashed lines indicate nonscrubbed CO2 /SO2 median and range of 9±1 for the 

2004–5 gases. Lower dashed line also represents the median value (8) of the 1980–81 CO2 /SO2 ratios as 

discussed in the text.

a specific cause of this return to scrubbing, but plausible factors 

may include the cracking and development of fracture permeability within the cooling and thickening carapace of the new 

dome and (or) seasonal changes affecting rainfall and groundwater recharge in the crater.

The CO2/SO2 time-series pattern in figure 18 illustrates the 

role of scrubbing in the early part of the current eruption and 

the close similarity of CO2/SO2 values in the nonscrubbed gas 

emissions of the 2004–5 eruption to those in the early 1980s. 

We contend, moreover, that scrubbing is the cause of the significant difference in CO2/SO2 indicated between the 1980–81 

and 2004–5 gases by the MWRST analysis. The striking pattern 

of CO2/SO2 ratios (fig. 18) with their approach to the median 

1980–81 CO2/SO2 ratio, hardly seems the result of chance, 

and it is consistent with flat magma remaining from 1986 that, 

although depleted in excess volatiles, still retains a similar CO2/

SO2 ratio. However, the depletion of excess volatiles causes the 

emissions from the flat magma upon ascent and eruption to be 

generally weak and susceptible to scrubbing if water is present.

Fluid Content of the 2004–2005 Dacite at Depth

Recent studies indicate that the dacite erupted explosively at 

Mount St. Helens on May 18, 1980, had a fluid content (that is, 

excess volatiles) prior to ascent and eruption of ≥15 volume percent 

(vol. percent), which corresponds to ≥3 weight percent (wt. percent) (Wallace, 2001, 2003; P.J. Wallace, written commun., 2005). 

The May 18, 1980, dacite was decidedly not flat magma. We have 

tested the flat-magma hypothesis for the 2004–5 dacite by using 

the current CO2 emission data to quantify the fluid content of the 

2004–5 dacite at depth. If the flat-magma hypothesis is valid for 

the current dacite, its fluid content at depth should be significantly 

lower than that determined for the May 18, 1980, dacite.

Fluid Content of the 2004–2005 Dacite and the 

May 18, 1980, Dacite at 900°C and 220 MPa (8.6 

km Depth)

The May 18, 1980, dacite last equilibrated prior to eruption at a temperature of 900°C and a pressure of 220 MPa 

(Rutherford, 1993), corresponding to 8.6 km depth (from 

crustal density model B-B’ of Williams and others, 1987, used 

in this study to convert lithostatic pressures to depths beneath 

the crater). The rhyolitic melt of the May 18 dacite was relatively H2O rich prior to eruption. Microprobe studies of melt 

inclusions give water-by-difference concentrations of 4.6±1 

wt. percent (Rutherford, 1993), and recent ion-microprobe 

studies of melt inclusions in the May 18 white pumice show a 

clustering of H2O concentrations around 5 wt. percent (Blundy 

and Cashman, 2005).
560 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

To maintain comparability with the May 18 dacite, we 

have used the constraints 900°C, 220 MPa, and a rhyolitic 

melt H2O concentration of 5 wt. percent to estimate the 

fluid content of the 2004–5 dacite at 8.6 km depth. Several 

observations and experimental results support the use of 

these constraints. The current dacite contains phenocrysts of 

plagioclase, hypersthene, amphibole, and oxides—the same as 

the 1980–86 dacite, except for an absence of minor clinopyroxene (Pallister and others, this volume, chap. 30; Rutherford 

and Devine, this volume, chap. 31). The current dacite also is 

similar in bulk composition to the 1980–86 dacite, being only 

slightly more evolved (65 vs. 63 wt. percent SiO2) and distinct 

in U-series isotopic ratios of plagioclases (Cooper and Donnelly, this volume, chap. 36) and in bulk-composition ratios 

of Ti and Cr to SiO2 (Pallister and others, this volume, chap. 

30)—differences that are most readily explained by addition 

of a minor fraction of new magma to the reservoir between 

1986 and 2004 (Pallister and others, this volume, chap. 30). 

Melt inclusions provide no data on the H2O concentration of 

preeruption melt in the current dacite (Pallister and others, 

this volume, chap. 30; Rutherford and Devine, this volume, 

chap. 31; Blundy and others, this volume, chap. 33). However, 

on the basis of phase equilibria experiments, Rutherford and 

Devine (this volume, chap. 31) propose that the new magma 

entered the storage reservoir as dacite with a melt concentration >4 wt. percent—implied by the amphiboles—and at a 

temperature of ~900°C—suggested by high-TiO2 magnetite 

phenocrysts and the maximum anorthite content of cyclically 

zoned feldspar phenocrysts. These conditions are similar to 

those determined for the 1980 magma. To explain the cyclic 

zoning of the magma’s phenocrysts, Rutherford and Devine 

invoke convective circulation of injected dacite, which would 

have involved exposure to pressures of 220 MPa at various 

times before rising to the shallower part of the reservoir, where 

they infer it last equilibrated at about 850°C and 120–140 MPa 

prior to final ascent and eruption.

We used VolatileCalc (Newman and H2O Lowenstern, 

2002) to model CO2 solubility in hydrous rhyolitic melt 

at 900°C and 220 MPa to estimate the fluid content of the 

2004–5 dacite for comparison with the May 18 dacite at 

8.6 km depth. At these conditions, rhyolitic melt with 5 wt. 

percent H2O would contain associated dissolved CO2 at a concentration of ~350 ppm (fig. 19A). The mole fraction composition of the coexisting fluid would be ~0.76 for XH2

O and 

~0.24 for XCO2 (fig. 19B), which is consistent with hydrothermal phase equilibria experiments on the current dacite that 

indicate XH2O would be ≥0.7 in the presence of a CO2-bearing 

coexisting fluid (Rutherford and Devine, this volume, chap. 

31). The fluid also would contain a small amount of sulfur, 

but since CO2/SO2 is ~9 in the nonscrubbed gas emissions—

and CO2/S would likely have been higher in the fluid at 

elevated pressure—we ignore sulfur in the present analysis.

We calculated the fluid content of the dacite at 8.6 km 

and the conditions delineated above for 14 dates in 2004–5 

corresponding to the dates of digital elevation models (DEMs; 

Schilling and others, this volume, chap. 8). Table 3 lists values 

keyed to DEM dates for several parameters (defined in table 3 

headnotes) used in the calculations. The last 2005 DEM date 

(12/15/2005) could not be used because weather conditions 

prevented measurement of CO2 emission rates during December 2005, as noted above.

Table 3 contains values for two melt mass parameters: 

one for dacite like the 2004–5 dacite that contained 53.5 

vol. percent melt prior to eruption (Pallister and others, this 

volume, chap. 30), which we discuss later, and the other for 

preeruption dacite with 70 vol. percent melt. The 70-vol. percent case for melt masses is based on the groundmass abundances of 65–80 vol. percent reported for the somewhat less 

evolved May 18, 1980, white pumice (Cashman and Taggert, 

1983; Rutherford and others, 1985; Cashman, 1992) and a 

200

0

400

600

800

1,200

1,000

1,400

0 1 2 3 4 5 6

MELT H2O CONTENT, IN WEIGHT PERCENT

MELT CO2 CONTENT,

IN PARTS PER MILLION

CO2 saturation (348 ppm) at 5 wt% H2O

A

CO2

0.2

0

0.4

0.6

0.8

1.0

0 1 2 3 4 5 6

MELT H2O CONTENT, IN WEIGHT PERCENT

FLUID COMPOSITION,

IN MOLE FRACTION

H2O

CO2

0.24 CO2

0.76 H2O

B

Figure 19. Concentration of CO2 in hydrous rhyolitic melt 

and coexisting fluid composition at 220 MPa and 900°C 

(Newman and Lowenstern, 2002). A, Solubility of CO2 (solid 

blue curve) in hydrous rhyolitic melts with a range of water 

concentrations. Dotted lines indicate a CO2 saturation limit 

of 348 ppm for rhyolitic melt containing 5 wt. percent H2O. 

Melt H2O content is total H2O—that is, OH and molecular 

H2O. B, Mole fraction concentrations of H2O and CO2 (solid 

curves) in fluids coexisting with hydrous rhyolitic melts 

over a range of water concentrations. Dotted lines indicate 

fluid composition of 0.76 XH2

O and 0.24 XCO2

 coexisting with 

rhyolitic melt containing 5 wt. percent water.
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 561

DEM date

Dacite 

volume

(106

 m3)

DRE 

volume

(106 m3)

Dacite 

mass

(106 t)

Melt mass

for 53.5 vol. 

percent melt

(106

 t)

Melt mass

for 70 vol. 

percent melt

(106

 t)

CO2

mass

(t)

11/04/2004 11.8 10.62 27.82 13.07 17.10 23,482

11/20/2004 18.4 16.56 43.39 20.38 26.66 39,380

11/29/2004 21.3 19.17 50.23 23.59 30.86 48,800

12/11/2004 25.5 22.95 60.13 28.24 36.95 61,162

01/03/2005 30.5 27.45 71.92 33.78 44.19 82,459

02/01/2005 35.1 31.59 82.77 38.87 50.86 102,639

02/21/2005 39.2 35.28 92.43 43.41 56.80 112,188

03/10/2005 41.9 37.71 98.80 46.40 60.71 118,092

04/19/2005 47.5 42.75 112.01 52.60 68.83 127,630

06/15/2005 53.9 48.51 127.10 59.69 78.10 141,901

07/14/2005 57.1 51.39 134.64 63.24 82.74 152,123

08/10/2005 61.7 55.53 145.49 68.33 89.40 163,184

09/20/2005 67.3 60.57 158.69 74.53 97.52 182,139

10/24/2005 70.0 63.00 165.06 77.52 101.43 202,885

Table 3. Parameters for calculation of fluid content of the 2004–2005 dacite, Mount St. 

Helens, Washington.

[Digital elevation model (DEM) dates are the dates of DEMs of Schilling and others (this volume, chap. 8). 

Dacite volumes are estimated cumulative volumes of the newly extruded lava dome (Schilling and others, this 

volume, chap. 8). Dense rock equivalent (DRE) volumes are cumulative dacite volumes adjusted for an average 

porosity of 10 vol. percent (Cashman and others, this volume, chap. 19; Pallister and others, this volume, chap. 

30; C. Thornber, written commun., 2006). Dacite masses are cumulative masses in metric tons (t) of dacite 

derived from the DRE volumes and the dacite DRE density of 2.62 g/cm3 (K. Russell, written commun., 2006). 

Melt masses (t) are cumulative masses of melt estimated from DRE volumes and preeruption melt fractions of 

53.5 vol. percent and 70 vol. percent, as described in the text. CO2 masses (t) are cumulative masses of CO2 output as of each DEM date, calculated from a curve fit to the cumulative CO2 emissions (fig. 8). Significant figures 

were exaggerated to reduce roundoff errors.]

melt density of 2.3 g/cm3 reported for experimental investigations of rhyolite densities at elevated temperature and water 

pressure (Silver and others, 1990). The 70-vol. percent case 

maintains comparability with 1980 conditions of 900°C, 220 

MPa, and rhyolitic melt containing 5 wt. percent dissolved 

H2O at 8.6 km depth.

For each DEM date, the amount of dissolved CO2 was 

calculated from the concentration of CO2 in the melt (fig. 19A) 

and the mass of melt for the DEM date (table 3), as estimated 

from the 70-vol. percent melt fraction discussed above. In all 

cases, the amount of dissolved CO2 was a minor fraction (<20 

percent) of the cumulative (total) CO2 output as of the DEM 

date (table 3), indicating that 2004–5 dacite containing 70 

vol. percent rhyolitic melt with 5 wt. percent dissolved H2O 

would indeed have been fluid saturated in the magma reservoir 

at 900°C and 220 MPa (8.6 km depth). We determined the 

amount of CO2 in the fluid phase by subtracting the amount 

of CO2 dissolved in the melt from the cumulative CO2 emitted 

as of the DEM date (table 3). Having determined the amount 

of fluid CO2, we calculated the amount of fluid H2O from the 

H2O-CO2 fluid compositional relation (fig. 19B). We converted 

the amounts of fluid CO2 and H2O into a fluid volume by 

assuming ideal mixing and by using the modified RedlichKwong (MRK) equation of state (Holloway, 1977, 1981) to 

calculate molar volumes of CO2 and H2O at 900°C and 220 

MPa. The volume of fluid and the DEM-based dacite volume 

(table 3), corrected to dense rock equivalent (DRE) basis (table 

3), gave the fluid content as a volume percent of the dacite. 

The dacite masses (table 3) and the amounts of fluid CO2 and 

H2O also allowed expression of the fluid content as a weight 

percent of the dacite.

Figure 20 shows the fluid content of the dacite for the 14 

DEM dates at 900°C and 220 MPa (8.6 km depth). The results, 

ranging from 0.82 to 1.32 vol. percent (0.15–0.24 wt. percent), 

indicate steady fluid content for the dacite at 8.6 km depth 

after a brief period of regularly increasing values involving 

the points for the first four DEM dates (November–December 

2004). Steady fluid content begins with the January 3, 2005, 

DEM date and continues thereafter; it ranges from 1.16 to 

1.32 vol. percent (0.21–0.24 wt. percent) and averages 1.23 

vol. percent (0.22 wt. percent). These results are consistent 

with those of Mastin and others (this volume, chap. 22), who 

also inferred a low reservoir fluid content of <1.5 vol. percent from the relation between erupted-dacite volume and the 

volume shrinkage of the reservoir obtained from geodetic data 

(Lisowski and others, this volume, chap. 15).
562 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

It is noteworthy that the results are steady throughout 

2005 despite the disparate datasets involved. Several factors 

may be at work during the brief period involving the four 

points of lower but steadily rising fluid content during November–December 2004 (fig. 20). These factors probably reflect 

the declining influence with time of CO2 lost in scrubbing 

prior to November 2004 (or missed if airborne measurements 

were too infrequent) on the cumulative CO2 output. Increasing the cumulative CO2 outputs by 38 percent (11/4/2004) to 

15 percent (12/11/2004) would bring the four points into the 

steady range of values. Alternatively, the lower fluid content 

may reflect a higher porosity of the earlier dacite—somewhere 

in the range 11–32 vol. percent (K. Russell, written commun., 

2006)—compared to the eruption average of 10 vol. percent 

(table 3) used in the calculations, although there is presently 

no evidence that porosity declined in a steady way early in the 

eruption. A higher porosity decreases the DRE volume and 

mass of dacite in proportion to the cumulative amount of CO2

and causes the volume percent fluid for the four 2004 DEM 

dates to increase—porosities of 35 vol. percent (11/4/2004) 

to 22 vol. percent (12/11/2004) would cause the fluid content 

of these dates to fall within the range of steady 2005 results. 

Involvement of shallower, more degassed magma early in the 

eruption might also account for the lower 2004 fluid content, 

but it is not obvious from the CO2 emissions that the earlier 

magma was more degassed (figs. 7, 17). Lastly, the procedures 

used to estimate dacite volumes from DEMs may overestimate 

the smaller dacite volumes of the eruption-startup period, causing erroneously low calculated fluid content. Significant underestimation of early dacite volumes seems more likely, however, 

because of unaccounted for dacite residing in the conduit. 

Apparently, the volume of conduit dacite became a negligible 

fraction of the measurable dacite volume fairly early—sometime prior to the first DEM on 11/4/2004 (table 3).

The sensitivity of the modeling to the volume percent 

melt-fraction parameter is an important concern that is easily addressed. Recalculating the model for a melt fraction of 

53.5 vol. percent (table 3), compared to the 70-vol. percent 

melt fraction used above, illustrates the low sensitivity of the 

results to a plausible range of values for the volume percent 

melt parameter. The resulting steady 2005 fluid content ranges 

from 1.23 to 1.39 vol. percent and averages 1.29 vol. percent—

hardly different from the above results for 70 vol. percent melt.

Recently, Liu and others (2005) derived a new empirical 

model of CO2 and H2O solubility in fluid-saturated rhyolitic 

melt and recommended it over VolatileCalc (Newman and 

Lowenstern, 2002) for eruptive degassing and magmachamber dynamics calculations because it gives better fits to 

experimental solubility data, especially at pressures above 

200 MPa. At the conditions considered here, though, the two 

models agree closely on the dacite fluid content. For example, 

during the period of steady 2005 DEM dates, the model of Liu 

and others (2005) gives 1.17–1.33 vol. percent and a 1.24-vol. 

percent average versus 1.16–1.32 vol. percent and a 1.23-vol. 

percent average by VolatileCalc. The agreement becomes virtually exact in calculations at the lower pressures considered 

below. For these reasons, and because of VolatileCalc’s userfriendly software, we have used it throughout this study.

We conclude that the average steady 2005 fluid content of 

1.23 vol. percent (0.22 wt. percent; fig. 20) is the best estimate 

of the fluid content of the current dacite at a reservoir depth of 

8.6 km, 900°C, and 220 MPa. This result and, indeed, all the 

results obtained above are significantly less than the ≥15-vol. 

percent (≥3 wt. percent) fluid determined for the May 18 dac8/10/05

9/20/05

7/14/05

10/24/05

11/4/04

11/20/04

11/29/04

12/11/04

2/1/05

1/3/05

2/21/05

3/10/05

4/19/05

6/15/05

0.8

0.9

1.1

1.0

1.2

1.3

1.4

Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct

FLUID CONTENT OF PREERUPTION DACITE,

IN VOLUME PERCENT

2004 2005

Figure 20. Fluid content of dacite at 900°C and 220 MPa (8.6 km) calculated as described 

in the text for 14 DEM dates in 2004–5.
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 563

ite; furthermore, they are much less than the 5–30-vol. percent 

(1–6 wt. percent) fluid indicated for several other explosive 

eruptions of silicic and intermediate magma (Wallace, 2001, 

2003). The present analysis, therefore, indicates depletion of 

excess volatiles and thus confirms the flat-magma hypothesis 

for the dacite of the 2004–5 eruption.

Fluid Content of the 2004–2005 Dacite Prior to 

Ascent at 850°C and 130 MPa (5.2 km Depth)

As the dacite migrates to reservoir depths shallower than 

8.6 km, its fluid content increases. Of particular interest is the 

fluid content where the dacite last equilibrated at about 850°C 

and 120–140 MPa in the presence of an H2O-rich fluid (XH2

O

≈ 1) prior to ascent and eruption (Rutherford and Devine, this 

volume, chap. 31). The corresponding depth range of 4.8–5.6 

km (Williams and others, 1987) places the preeruption dacite 

in the shallowest part of the ellipsoidal-cylindrical magma 

reservoir extending from 4.5–5 km to ≥12 km, as indicated 

by deformation data from the current eruption (Lisowski and 

others, this volume, chap. 15). The average volume percent 

groundmass of the 2004–5 dacite is 53.5 (Pallister and others, 

this volume, chap. 30; Rutherford and Devine, this volume, 

chap. 31). As such, it is a reasonable approximation of the 

rhyolitic melt fraction of the current dacite before its ascent 

and eruption from a shallow reservoir depth of 5.2 km (130 

MPa)—the median of the 4.8–5.6-km depth range (120–140 

MPa) where the dacite last equilibrated. This is the basis for 

the melt mass case corresponding to a melt fraction of 53.5 

vol. percent in table 3. Accordingly, we have taken our results 

for all DEM dates obtained above (fig. 20) for dacite containing 70 vol. percent melt at 900°C, 220 MPa, and 8.6 km depth 

and repartitioned CO2 and H2O from melt to fluid to obtain a 

dacite with 53.5 vol. percent melt. To bring this dacite to the 

last equilibrium conditions within the shallow reservoir at 5.2 

km (assuming a closed system), we used VolatileCalc to find 

the dissolved melt H2O and CO2 concentrations at 850°C and 

130 MPa (fig. 21A) that—after further adjusting of H2O and 

CO2 partitioning between fluid and the 53.5-vol. percent melt 

phase—gave a fluid composition consistent with the XH2

O and 

XCO2 predicted by the model (fig. 21B). This iterative process 

consistently gave a melt H2O concentration of ~4.4 wt. percent 

and an associated melt CO2 concentration of ~37 ppm (fig. 

21A) for each DEM date. The coexisting fluid composition 

is XH2

O = ~0.96 and XCO2

 = ~0.04 (fig. 21B), consistent with 

experimental results of Rutherford and Devine (this volume, 

chap. 31) that indicate the XH2

O of coexisting fluid would be 

close to 1.0 in the shallower part of the reservoir. These results 

also constrain the bulk dacite concentrations of CO2 and 

H2O—that is, fluid plus melt amounts of CO2 and H2O as mass 

fractions of the dacite—at average values of 1,167 ppm and 

3.1 wt. percent, respectively, for the ten 2005 DEM dates. The 

corresponding average bulk dacite S concentration calculated 

from the cumulative SO2 emissions (fig. 8; table 2) is 90 ppm 

for the 10 DEM dates.

The fluid content results for all the DEM dates range 

from 10.11 to 10.85 vol. percent (fig. 22); again, there is a 

tight grouping of steady fluid content for the ten 2005 DEM 

dates—10.63 to 10.85 vol. percent with an average 10.72 vol. 

percent. The large increase in the average volume percent fluid 

content from 1.23 to 10.72 vol. percent in going from 8.6 km 

to 5.2 km depth reflects the crystallization of melt (from 70 

vol. percent melt to 53.5 vol. percent melt) and the decreased 

solubility of H2O and CO2 in rhyolite melt at lower pressure; it 

also includes the effect of fluid expansion as the pressure drops 

from 220 MPa to 130 MPa. The fluid expansion factor confounds comparisons of fluid content of magmas from different 

depths. This confusion is avoided by using mass units, which do 

200

0

400

600

800

1,000

0 1 2 3 4

MELT H2O CONTENT, IN WEIGHT PERCENT

MELT CO2 CONTENT,

IN PARTS PER MILLION

CO2 saturation (37 ppm) at 4.4 wt% H2O

A

CO2

0.2

0

0.4

0.6

0.8

1.0

0 1 2 3 4

MELT H2O CONTENT, IN WEIGHT PERCENT

FLUID COMPOSITION,

IN MOLE FRACTION

H2O

CO2

0.04 CO2

0.96 H2O

B

Figure 21. Concentration of CO2 in hydrous rhyolitic melt and 

coexisting fluid composition at 130 MPa and 850°C (Newman and 

Lowenstern, 2002). A, Solubility of CO2 (solid blue curve) in hydrous 

rhyolite melt with a range of water concentrations. Dotted 

lines indicate a CO2 saturation limit of 37 ppm for rhyolitic melt 

containing 4.4 wt. percent H2O. Melt H2O content is total H2O—that 

is, OH and molecular H2O. B, Mole fraction concentrations of H2O 

and CO2 (solid curves) in fluids coexisting with hydrous rhyolitic 

melts over a range of water concentrations. Dotted lines indicate 

fluid composition of 0.96 XH2

O and 0.04 XCO2

 coexisting with rhyolitic 

melt containing 4.4 wt. percent water.
564 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

not reflect fluid expansion; in the present case, the steady fluid 

content at 5.2 km depth ranges from 1.22 to 1.25 wt. percent 

and averages 1.23 wt. percent. Despite the relatively shallow 

depth, these results are still significantly smaller than Wallace’s 

≥3 wt. percent fluid content for the May 18 dacite at 8.6 km 

depth and on the low end of his 1–6 wt. percent range for deeper 

intermediate to silicic magmas involved in explosive volcanism 

(Wallace, 2001, 2003). These results further confirm the flatmagma hypothesis for the 2004–5 dacite.

We have assumed closed-system degassing with respect 

to the excess-volatile fluid phase’s mobility in magma at 

depth. The steady excess-volatile fluid content of ~1.23 wt. 

percent calculated for the dacite at 5.2 km throughout 2005 

is consistent with this assumption. Equilibrium (210Pb)/(226Ra) 

values reported for dacite erupted in 2004–5 (Reagan and 

others, this volume, chap. 37) also support closed-system 

degassing. Unlike those of the 1980 lava and tephra (Berlo 

and others, 2004), (210Pb)/(226Ra) values of the 2004–5 dacite 

were not affected by continuous open flow of fluid carrying 222Rn from depth in sufficient amount or for sufficient 

duration to disturb secular equilibrium and generate 210Pb 

deficits or excesses (Reagan and others, this volume, chap. 

37); in our view, this reflects the depletion of excess volatiles 

compared to 1980 dacite. However, to explain the Li enrichment unique to plagioclase phenocrysts of the October 2004 

dacite, Kent and others (2007) propose preferential partitioning of Li from deep melt into aqueous fluid migrating to the 

uppermost part of the magma chamber at 4.5 km, about 1 

year before eruption; they argue that at this depth, which corresponds to 110 MPa (Williams and others, 1987), the fluid 

8/10/05

9/20/05

7/14/05

10/24/05

11/4/04

11/20/04

11/29/04

12/11/04

2/1/05

1/3/05

2/21/05

3/10/05

4/19/05

6/15/05

10.1

10.2

10.5

10.3

10.4

10.6

10.7

10.8

10.9

Nov Dec Jan Feb Mar Apr May Jun Jul Aug Sep Oct

FLUID CONTENT OF PREERUPTION DACITE,

IN VOLUME PERCENT

2004 2005

unmixed into a low-density vapor and a Li-rich brine that 

fostered the enrichment of Li in what became the October 

2004 dacite’s plagioclase phenocrysts. We suggest that Li 

in the dacite magma is partitioned preferentially into the 

accompanying excess-volatile fluid and that the Li-enriched 

plagioclase phenocrysts resulted when that fluid unmixed a 

Li-rich brine as the dacite ascended from 5.2 km to 4.5 km, 

where the magma stalled for several months before final 

ascent to the surface at the start of the eruption.

Ascent Degassing

We carried out calculations like those described above 

for the current dacite to model its closed-system degassing at 

850°C during ascent from shallow reservoir depths. The calculations began at 5.2 km (130 MPa) and terminated at 0.47–0.95 

km (10–20 MPa) where melt of the 1980s lava-dome eruptions 

crystallized extensively to groundmass (Blundy and Cashman, 

2001). Figure 23 shows the fluid content, fluid composition, 

and melt H2O and CO2 concentrations of the dacite with ascent 

depth. We evaluated these variables for all DEM dates at each 

pressure of calculation and then correlated results to appropriate depths (Williams and others, 1987). The curves of figure 23 

employ the average values of the ten 2005 DEM dates, which 

agreed within 5 percent.

Figure 23A shows the fluid content of the 2004–5 dacite 

during ascent; the fluid content of current dacite at 8.6 km 

and that of the May 18, 1980, dacite (Wallace, 2003) are 

included for comparison. The fluid content of the ascending 

2004–5 dacite ranges from 10.7 vol. percent at the last equiliFigure 22. Fluid content of dacite at 850°C and 130 MPa (5.2 km) calculated as described in 

the text for 14 DEM dates in 2004–5.
Figure 23. Closed-system ascent of 2004–2005 dacite from 5.2-km depth of last equilibration at 850°C and 130 MPa before eruption (Rutherford achap. 31) to 0.47–0.95-km depth range of groundmass crystallization at 10–20 MPa (Blundy and Cashman, 2001). A, Fluid content (exolved fluid phasdacite during ascent, and fluid content of 2004–2005 dacite and May 18, 1980, dacite (Wallace, 2003) at 8.6 km as discussed in the text. B, H2O-CO2 f

2004–2005 dacite during ascent. C, Dissolved H2O in melt of 2004–2005 dacite during ascent; dissolved H2O in six melt inclusions (filled squares) wit

H2O concentrations (Pallister and others, this volume, chap. 30; Blundy and others, this volume, chap. 33); and dissolved H2O in a rare glassy fragm

October 2004 dacite sample SH304 (Pallister and others, 2005). D, Dissolved melt CO2 concentration of 2004–2005 dacite during ascent.

0123456789

0 10 20 30 40 50 60 70 80 90

FLUID CONTENT OF PREERUPTION DACITE, IN VOLUME PERCENT

DEPTH, IN KILOMETERS

May 18, 1980, dacite

2004–2005 dacite

2004–2005 dacite,

closed-system ascent

850°C, 130–10 MPa 

Percolation threshold

Groundmass crystallization

850°C, 20–10 MPa

A

2004–2005 dacite,

last equilibration

850°C, 130 MPa

012345

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 

FLUID COMPOSITION, IN MOLE FRACTI

DEPTH, IN KILOMETERS CO2 

B

2004–2005 dacite,

closed-system ascent

850°C, 130–10 MPa

Groundmass crystallization

850°C, 20–10 MPa

2004–2005 dacite,

last equilibration

850°C, 130 MPa

012345

0 1 2 3 4 5

MELT H2O CONTENT, IN WEIGHT PERCENT

DEPTH, IN KILOMETERS Dissolved H2O in melt inclusions Dissolved H2O in glassy fragment 

C

2004–2005 dacite,

closed-system ascent

850°C, 130–10 MPa 

Groundmass crystallization

850°C, 20–10 MPa

2004–2005 dacite,last equilibration850°C, 130 MPa

012345

0 10 20 

MELT CO2 CONTENT, IN PARTS PER MIL

DEPTH, IN KILOMETERS

Groundmass crystallization

850°C, 20–10 MPa

2004–2005 dacite,

closed-system ascent

850°C, 130–10 MPa 

D
566 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

bration depth to 60–80 vol. percent at depths of groundmass 

crystallization. The corresponding fluid composition of 

2004–5 dacite during ascent (fig. 23B) is H2O rich with XH2

O

always >0.96 and reaching 0.98 at the stage of groundmass 

crystallization. Extrapolating the compositional trends to 

near-surface conditions indicates that H2O forms about 99 

percent (molar basis) of the magmatic gases emitted to the 

atmosphere, whereas the monitored gases CO2 and SO2

account for only about 1 percent of the emitted magmatic 

gases. The implied magmatic H2O/(CO2+SO2) of ~99 is consistent with the value of 113 measured by open-path Fouriertransform infrared spectroscopy (Edmonds and others, this 

volume, chap. 27) on gas emissions from the new dome on 

August 31, 2005, considering that dome gases are likely to 

contain a meteoric component in addition to the magmatic 

component of water. Our CO2 and SO2 emission-rate data 

thus allow calculation of expected magmatic H2O emission 

rates during 2004–5; the expected magmatic H2O emission 

rates range from ~7 kt/d to ~100 kt/d with an estimated 

median of ~30 kt/d. We estimate cumulative H2O production 

through 2005 at ~10 Mt from the cumulative production of 

CO2 and SO2.

The concentration of H2O—that is, total OH and 

molecular H2O—dissolved in the melt of the 2004–5 dacite 

ranges during ascent from the 4.4 wt. percent of last equilibration in the reservoir at 5.2 km to about 1.1–1.5 wt. percent 

at the conditions of groundmass crystallization (fig. 23C). 

The ascent curve for H2O dissolved in melt indicates last 

equilibration depths of about 1.3–1.6 km for two measurements of H2O concentration by Fourier-transform infrared 

spectrometry on a rare glassy fragment found in the October 

2004 dacite sample SH304 (Pallister and others, 2005). Melt 

inclusions in the current dacite contain 0.0–3.6 wt. percent 

H2O (Pallister and others, this volume, chap. 30; Blundy and 

others, this volume, chap. 33). The six highest melt-inclusion H2O concentrations range from 2.4 to 3.6 wt. percent, 

corresponding to depths of only 2.0–3.7 km on the ascent 

curve for H2O (fig. 23C). These results corroborate petrologic findings indicating that the melt inclusions provide no 

data on the volatile content of preeruption melt, apparently 

because slow ascent and eruption caused rupturing and (or) 

crystallization of virtually all melt inclusions, as suggested 

by their frequent mafic-mineral hosts, high crystallinity (≥70 

percent), and variable SiO2 concentrations (in this volume: 

Pallister and others, chap. 30; Rutherford and Devine, chap. 

31; Blundy and others, chap. 33). Alternative explanations 

of the low melt-inclusion H2O concentrations—for example, 

that they formed during ascent from 5.2 km, or that the 

source magma at 5.2 km depth was relatively dry (<4 wt. 

percent H2O, which would imply XH2

O <0.84)—are implausible in view of petrologic evidence. The concentration of 

CO2 dissolved in the melt during ascent ranges from 37 ppm 

at the last equilibration in the reservoir at 5.2 km to about 

1.3–2.9 ppm at the conditions of groundmass crystallization 

(fig. 23D).

The percolation threshold of 30 vol. percent indicated at 

~2.5 km (fig. 23A) is a first-order approximation of the depth 

beyond which further ascent causes interconnected bubbles 

to increasingly dominate in overall degassing, ultimately 

transitioning to open-system degassing (Mueller and others, 

2005). Melt concentrations of H2O and CO2 at the percolation threshold are about 2.8 wt. percent and 10 ppm, respectively, and the associated fluid composition is H2O rich (XH2

O

= 0.975). The transition to open-system degassing during 

ascent at some depth shallower than ~2.5 km is an important 

step in the mobilization and flow of the H2O-rich volcanic 

gases. Measured H2O/HCl ratios of gas emissions on August 

31, 2005, are consistent with a transition to open-system 

degassing at depths of ~1 km and prevailing to the surface 

(Edmonds and others, this volume, chap. 27). Open flow of 

H2O-rich volcanic gases during ascent and dome emplacement purged shallow dacite of other volatile elements, 

including the entrainment of 210Po (Reagan and others, this 

volume, chap. 37) and the loss of Li from melt and the margins of plagioclase phenocrysts in the October 2004 dacite 

(Kent and others, 2007).

Fluid H2S/SO2 During Ascent

The gas-emission data show that SO2 is the main sulfur 

species emitted to the atmosphere during the 2004–5 eruption. However, SO2 may not be the dominant sulfur gas 

species at depth, where increased pressure can strongly affect 

sulfur speciation. Gerlach and Casadevall (1986) proposed 

from thermodynamic calculations constrained by fumarole 

gas compositions that H2S was the dominant sulfur gas in 

the 1980s dacite at depth, whereas SO2 dominated at pressures less than 4 MPa (depths less than 0.2 km). They cited 

reports of high H2S/SO2 in Mount St. Helens plumes from 

explosive eruptions of 1980–1981 in which 35–100 percent 

of the sulfur was present as H2S (Hobbs and others, 1982) 

and suggested that rapid release of gas from depth prevented 

full equilibration to low pressure and allowed high H2S/

SO2 to persist. They further suggested that rapid unloading 

during explosive eruptions of magma with oxygen fugacity 

as much as 1 order of magnitude above the Ni-NiO (NNO) 

buffer could burden the atmosphere with significant H2S, as 

well as SO2, and that SO2 might increase as the H2S oxidized, 

because the lifetime of H2S in the atmosphere is only ~1 day 

(Graedel, 1977). Subsequent examination of TOMS satellite 

measurements made once a day on stratospheric volcanic 

clouds showed higher SO2 masses for the May 18, 1980, 

cloud from Mount St. Helens on the second day than on the 

first day after eruption (Bluth and others, 1995).

We have taken a different approach to that of Gerlach and 

Casadevall (1986) and used the reaction

SO2 + H2O = H2S + 1.5 O2
26. Emission Rates of CO2, SO2, and H2S, Scrubbing, and Preeruption Excess Volatiles at Mount St. Helens, 2004–2005 567

to estimate the H2S/SO2 ratio of the fluid phase during ascent 

of the current dacite from its last equilibration depth to the 

surface. For this reaction, the molar H2S/SO2 ratio is given by

H2S/SO2 = (pXH2

O/KƒO2

1.5)(γH2

OγSO2

/γH2

S), (2)

where p is the total pressure in bars (1,300 bars to 1 bar), 

XH2

O is the mole fraction of H2O in the fluid (0.96 to 0.99, 

fig. 23C), K is the reaction equilibrium constant of 1020.03 at 

850°C (HSC Chemistry for Windows v. 5.1), ƒO2 is the current dacite’s oxygen fugacity of 10-12.29 bar at 850°C (Pallister 

and others, this volume, chap. 30), and γi terms are dimensionless fugacity coefficients correcting for the nonideality 

of H2O, SO2, and H2S. Fugacity coefficients from the MRK 

equation of state (Holloway, 1977, 1981), the corresponding 

state equation of Shi and Saxena (1992), and the ideal gas 

assumption (γi = 1) give H2S/SO2 results that agree within 10 

percent, consistent with the relatively low-pressure, hightemperature conditions. The results confirm that H2S is the 

dominant sulfur gas at depths >0.2 km (fig. 24), as in the 

1980s (Gerlach and Casadevall, 1986); the H2S/SO2 ratio 

is >40 at the last equilibration depth in the shallow reservoir (5.2 km) and 3.3–6.5 in the depth range of groundmass 

crystallization (0.47 to 0.95 km). Compared to the explosive eruptions of the early 1980s, the slow ascent of magma 

from depth in the present eruption favors a closer approach 

to equilibrium with more complete conversion of H2S to 

SO2—consistent with the generally low H2S emissions of 

the current eruption. Nevertheless, field crews often reported 

strong intermittent H2S odor, especially after the early steamand-ash events, and H2S odor was reported by an airborne 

observer downwind of the volcano during and immediately 

after the first explosion of the eruption on October 1, 2004. 

These reports may be the result of rapid release of H2S from 

depth. However, we cannot rule out an alternative H2S origin 

0

1

2

3

4

5

0 5 10 15 20 25 30 35 40 45

DEPTH, IN KILOMETERS 

H2S / SO2, IN MOLES / MOLE

Figure 24. Calculated fluid H2S/SO2 of 2004–2005 dacite at 

850°C ascending from the 5.2-km depth (130 MPa) of last 

equilibration in the shallow reservoir to the surface (0.1 MPa), as 

described in the text.

by scrubbing and hydrolysis of SO2, as discussed above, for 

the early days of the eruption in 2004.

Excess Sulfur

The existence of an excess sulfur problem, although 

clear in the 1980s cycle of volcanism at Mount St. Helens 

(Gerlach and McGee, 1994), is problematic in the 2004–5 

eruption. Assuming that the cumulative SO2 emissions (table 

2; fig. 8) for the DEM dates derived entirely from the melt 

fraction of the dacite prior to ascent implies a melt S concentration of 192 ppm at 5.2 km depth. The rupturing and high 

crystallinity of virtually all melt inclusions and the belowdetection-limit sulfur concentrations of most melt inclusions 

and matrix glasses (Pallister and others, this volume, chap. 

30), however, prevent the use of glass measurements to constrain the preeruption melt sulfur concentration of the current 

dacite prior to ascent from 5.2 km, thus precluding comparisons with the 192-ppm figure. Whether or not the preeruption melt sulfur concentration at 5.2 km depth accounts for 

the 2004–5 SO2 emissions is therefore presently irresolvable, 

as is the question of excess sulfur for the 2004–5 eruption. 

The alternative source of the SO2 emissions—that is, the 

fluid fraction of the dacite at 5.2 km—thus remains viable; 

this option would require a mole-fraction S concentration 

in the fluid (XS) of only 0.004 prior to ascent (recall that the 

fluid XH2

O and XCO2

 at 5.2 km are about 0.96 and 0.04, respectively). Moreover, experimental studies (Keppler, 1999) on 

haplogranitic melt at 850°C, 200 MPa, and ƒO2 of 0.5 log 

units above the Ni-NiO buffer (NNO + 0.5)—close to the 

current dacite’s ƒO2 of NNO + ~1 (Pallister and others, this 

volume, chap. 30)—give fluid/melt partition coefficients for 

S of 47±4. These results strongly favor excess sulfur in the 

fluid fraction of the dacite at 5.2 km as the main source of 

the SO2 emissions.

Concluding Remarks

The gas emissions of the 2004–5 eruption of Mount St. 

Helens provide no compelling evidence that gas-rich “new” 

magma was introduced into the reservoir during the months 

just prior to or since the onset of activity in September 2004, 

although additions of gas-poor new magma cannot be ruled 

out. The gas emissions are instead indicative of a flat source 

magma markedly depleted in excess volatiles compared to the 

May 18, 1980, dacite and the intermediate and silicic magmas 

commonly involved in explosive volcanism. However, the flat 

dacite of this eruption contains rhyolitic melt with appreciable 

dissolved water (4.4 wt. percent) despite its depletion in excess 

volatiles. It remains unclear whether flat dacite of this sort 

can give rise to large and violent explosive eruptions. The flat 

dacite has not done so in the present case, but this fact begs 

the question of how dependent explosive volcanism involving intermediate and silicic magma is on the magnitude of the 

excess-volatile load.
568 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

The 2004–5 gas emissions are plausibly compatible 

with 1980–86 dome-building magma left over in the reservoir 

and (or) with gas-poor new magma that entered the reservoir 

since 1986. Petrologic evidence favors the introduction of new 

magma (Pallister and others, this volume, chap. 30), but why 

is this new magma so gas poor compared to the May 18, 1980, 

magma? Perhaps the new magma was initially depleted in 

excess volatiles, or it may have mixed with a large amount of 

leftover 1980–86 magma already greatly depleted in exsolved 

volatiles. Alternatively, the new magma may simply have lost 

most of its excess-volatile load by fluid escape while stored 

at shallow reservoir depths long before the events of September 2004. The degassing of CO2 measured on June 22, 1998, 

may represent the product of such fluid escape during a short 

period of seismicity accompanying recharge of the shallow 

reservoir. Although this is the strongest evidence of post-1986 

fluid escaping and reaching the surface, we cannot rule out the 

possibility that fluid escaped from the shallow reservoir at other 

times since 1986 and was captured by hydrothermal fluid at 

depth or scrubbed by surficial water (Shevenell and Goff, 1993); 

the large increase of HCO3

−

 in the waters of thermal springs at 

Loowit and Step canyons since 1994 (Bergfeld and others, this 

volume, chap. 25) may be the result of these processes.

The small load of excess volatiles in the 2004–5 dacite 

was an important factor in making it possible for scrubbing to shut down and strongly reduce gas emissions to the 

atmosphere in the period of unrest and the early stage of the 

eruption. The 2004–5 eruption affirms several implications of 

scrubbing for assessment of volcano hazards by volcanic gas 

monitoring: scrubbing can mask shallow magmatic degassing and thus severely restrict gas-emission monitoring as a 

volcano hazards assessment tool; scrubbing and its masking 

effects are most effective in the early stages of unrest; no great 

solace should be taken from low SO2 emission rates during 

periods of unrest; early monitoring of CO2 should accompany 

early SO2 monitoring; the monitoring of CO2, SO2, and H2S 

can track the drying out of a volcano and help distinguish the 

effects of gas scrubbing and permeability sealing on emissions; and emissions of CO2 and H2S in the early stages of 

unrest, when SO2 and emissions are absent or trivial, should 

not be regarded as merely the degassing of hydrothermal fluid, 

as they can reflect significant shallow magmatic degassing 

combined with the effects of scrubbing.

Acknowledgments

We are indebted to the following, all of whom made valuable contributions to this paper by sharing ideas, observations 

and data: Roger Denlinger, Dan Dzurisin, Marie Edmonds, 

Mike Lisowski, Larry Mastin, John Pallister, Kelly Russell, 

Steve Schilling, Carl Thornber, and Paul Wallace. Our special 

thanks to Jeff Sutton for help with instrumentation and early 

monitoring flights, and to Larry Mastin and John Pallister for 

technical reviews. Finally, we acknowledge the support of the 

U.S. Geological Survey Volcano Hazards Program.

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 27

1

U.S. Geological Survey, P.O. Box 51, Hawaii National Park, HI 96718; 

now at University of Cambridge, Downing Street, Cambridge CB2 3EQ, U.K.

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Chlorine Degassing During the Lava Dome-Building 

Eruption of Mount St. Helens, 2004–2005

By Marie Edmonds1, Kenneth A. McGee2, and Michael P. Doukas2

Abstract

Remote measurements of volcanic gases from the 

Mount St. Helens lava dome were carried out using OpenPath Fourier-Transform Infrared spectroscopy on August 31, 

2005. Measurements were performed at a site ~1 km from 

the lava dome, which was used as a source of IR radiation. 

On average, during the period of measurement, the volcanic 

gas contained 99 mol percent H2O, 0.78 percent CO2, 0.095 

percent HCl, 0.085 percent SO2, 0.027 percent HF, 4.8×10-4

percent CO, and 2.5×10-4 percent COS close to the active 

vent. The fluxes of these species, constrained by synchronous 

measurements of SO2 flux, were 7,200 t/d H2O, 140 t/d CO2, 

22 t/d SO2, 14 t/d HCl, 2.0 t/d HF, 54 kg/d CO, and 59 kg/d 

COS, ±20 percent. Observations of H2O/Cl in the vapor and 

melt are compared to models of closed- and open-system 

degassing and to models where a closed system dominates 

to depths as shallow as ~1 km, and gases are then allowed 

to escape through a permeable bubble network. Although 

several features are consistent with this model—for example, 

(1) H2

O/Cl in the gases emitted from stagnant parts of the 

lava dome, (2) the concentration of Cl in the matrix glass of 

erupted dacite, and (3) the glass H2O/Cl—the gases emitted 

from the active part of the lava dome have much higher H2O/

Cl than expected. These higher H2O/Cl levels result from 

a combination of two factors (1) the addition of substantial 

amounts of ground water or glacier-derived H2O to the gases 

at shallow depths, such that only ~10 mol percent of the measured H2

O is magmatic, and (or) (2) some Cl present as alkali 

chloride (NaCl and KCl) in the gas phase. The mean molar 

Cl/S is similar to gases measured at other silicic subductionzone volcanoes during effusive activity; this may be due to 

the influence of Cl in the vapor on S solubility in the melt, 

which produces a solubility maximum for S at vapor Cl/S ~1.

Introduction

Volcanic gases comprise water (H2O), carbon dioxide 

(CO2), sulfur dioxide (SO2), hydrogen sulfide (H2S), hydrogen chloride (HCl), and, to a lesser extent, hydrogen fluoride 

(HF), carbon monoxide (CO), and a host of trace species 

(for example, Symonds and others, 1994). Volatiles (mainly 

H2O and CO2) exsolve from magma during decompression, 

generating vesicularity and buoyancy relative to the surrounding rocks, thereby driving eruptions. The rate and style of 

degassing controls the rheological development of magma and 

governs eruptive style (for example, Sparks and others, 2000). 

Preeruptive dissolved H2O, CO2, S, and Cl contents in silicic 

magmas typically range from 1 to 6 weight percent and 0–400, 

30–200 and 900–3,000 ppm, respectively, after having undergone ~60–70 percent vapor-saturated fractional crystallization 

at 2–5 kb in the crust to evolve from mafic magma compositions (Wallace, 2005).

Chlorine is a significant volatile component of magmas 

associated with subduction zones and is derived from devolatilization of the subducting lithospheric slab (Manning, 2004). 

Recent analysis of volatile fluxes at subduction zones indicates 

that the flux of subducted Cl in the form of seawater and sediments approximately balances the output from arc volcanism 

(~4–7×1012 g/yr), consistent with the low abundance of Cl in 

mantle plume-derived basalt tapping the deep mantle (Wallace, 

2005). Along with other volatiles, Cl dissolves in the partial 

melt and, thereafter, behaves as an incompatible element during crystallization, although small amounts may be taken up 

by amphibole phases. Cl is finally removed from the melt by 

partitioning into a vapor phase during magma ascent through 

the crust. A high partition coefficient (vapor/melt) for Cl was 

first recognized by Kilinc and Burnham (1972). The release 
574 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

of Cl with H2O and other volatiles from magma is important 

for the transport and concentration of ore metals, such as gold, 

copper, and iron, which may eventually form mineral deposits 

in large magmatic systems (Lowenstern and others, 1991). 

The acidic nature of HCl-containing gases may leach and alter 

volcanic rocks and contribute to instability in volcanic edifices 

(Le Friant and others, 2002).

Hydrogen chloride is the dominant chloride species in 

volcanic gases (Webster and Holloway, 1988); significant 

amounts of NaCl and KCl have also been detected (Le Guern 

and others, 1975). Gaseous HCl enters the atmosphere during 

volcanic eruptions; tropospheric emissions can cause local 

acid rain, and stratospheric injection of HCl gas can lead to 

degradation of the ozone layer (Tabazedeh and Turco, 1993). 

The rate of emission of HCl from volcanoes, compared to that 

of SO2, has been linked to eruptive processes and has applications for volcano monitoring. Stoiber and Rose (1974) noted 

that gas Cl/S tends to increase throughout an eruptive episode, 

owing to the lower solubility of S. At Augustine Volcano, 

Symonds and others (1990) noted that older volcanic deposits 

exsolved gas with higher Cl/S than the main eruptive vents, 

suggesting that Cl continues to degas from melt after emplacement at the surface. At Soufrière Hills volcano3

, Montserrat, 

S-rich basaltic magma is supplied to the base of a porphyritic, 

Cl-rich andesite body. Volcanic gases with high levels of 

Cl/S are emitted during eruptive periods and gases with low 

levels of Cl/S are emitted during noneruptive periods. These 

emissions are caused by the degassing of negligible Cl but 

extensive S from the basalt to a vapor phase at 5–7 km; and 

the degassing of Cl but negligible S from the andesite during 

eruption (Edmonds and others, 2002).

During the current eruption of Mount St. Helens (ongoing in August 2006), CO2

, SO2, and H2S emission rates have 

been measured regularly and are well constrained (Gerlach 

and others, this volume, chap. 26); H2O, HCl, and other minor 

species had not been quantified until this study. HCl is an 

important constituent of volcanic gases, yet fewer measurements of HCl than of SO2

 are carried out during volcanic 

eruptions. This paper presents new data on the composition of 

the volcanic gas plume emitted from the lava dome at Mount 

St. Helens in August 2005. Remote spectroscopic measurements of the volcanic plume made by using Open-Path 

Fourier-Transform Infrared spectroscopy (OP–FTIR) yield 

the relative proportions of magmatic H2

O, CO2, HCl, SO2, HF, 

CO, and COS of the gas phase. Synchronous measurements 

of SO2 emission rates are used to calculate emission rates for 

the other six species.

The emission of Cl in the form of HCl gas is the focus of 

the analysis. In order to interpret the rate and mechanism of 

HCl emission from volcanoes, an understanding of how magmatic vapor evolves chemically with changes in pressure, tem3Capitalization of “Volcano” indicates adoption of the word as part of the 

formal geographic name by the host country, as listed in the Geographic 

Names Information System, a database maintained by the U.S. Board on Geographic Names. Noncapitalized “volcano” is applied informally—eds.

perature, and melt composition is necessary; this is reasonably 

constrained for H2O, CO2, and H2O-CO2 melt-vapor systems 

at shallow to intermediate pressures (for example, for basaltic 

melts, Dixon and others, 1995) but less well understood for 

systems containing S and Cl. The measured molar H2O/Cl of 

the gases is compared to models describing different degassing 

regimes in order to evaluate the evolution of the vapor phase 

with magma ascent. The gas composition is compared to the 

compositions of volcanic gases at other silicic volcanoes and 

to observations made during 1980–86 at Mount St. Helens. 

The Cl/S molar ratio is evaluated with reference to the solubility of S and Cl in silicic melts.

Method

Open-path Fourier-transform infrared spectroscopy 

(OP–FTIR) is a spectroscopic technique that utilizes the 

absorptivity of specific gas molecules in order to quantify 

them accurately (fig. 1A). A spectrometer acquires infrared 

(IR) radiation from a source that has passed through volcanic 

gas. Di-atomic and monatomic compounds (for example, 

H2, He) do not have an IR signature and, thus, cannot be 

measured by this technique. Absorption due to the gases is 

isolated and calibrated. OP–FTIR has been used successfully 

at several volcanoes, including Unzen, Japan (Mori and others, 1993); Vulcano, Italy (Francis and others, 1995); Etna, 

Italy (Burton and others, 2003); Soufrière Hills, Montserrat 

(Edmonds and others, 2002); and Kīlauea, USA (Edmonds 

and others, 2005). The OP–FTIR technique has advantages 

over direct methods of sampling, the most important of which 

is the remote aspect of the measurements and, hence, the 

higher level of safety, because it is not necessary to approach 

the fumaroles or vents closely. This technique is particularly 

useful for measuring gases at lava dome-building volcanoes, 

where explosions can occur with little warning; measurements can be collected throughout eruptive and noneruptive 

periods. For example, during the 1980–86 Mount St. Helens 

eruption, gas sampling was possible only from peripheral 

fumaroles, and when the lava dome was particularly active, 

no gas samples were collected at all (Gerlach and Casadevall, 

1986). OP–FTIR also allows rapid measurements; several 

hundred spectra may be collected every hour, each containing 

information about the amount of various major and minor gas 

species in the pathlength.

The OP–FTIR data presented in this paper were collected 

on August 31, 2005, using the 2004–5 Mount St. Helens lava 

dome as a source of IR radiation (fig. 1). A digital elevation 

model formulated from photographs taken on September 1, 

2005, (S.P. Schilling, written commun., 2008) is shown in figure 2. The massive parts of the dome were broken whalebackshaped spines, each surrounded by a thick layer of blocky 

talus material. The measurements were taken in three sets; 

pathlengths ranged from 1,040 to 1,150 m (fig. 2). Temperatures on the surface of the lava dome were highly variable, 
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 575

ranging from ~10°C to >200°C (fig. 3). For the OP–FTIR 

measurements, cracks in the spine surfaces were used as IR 

sources and had temperatures as high as 600°C (Schneider and 

others, this volume, chap. 17). Volcanic gas rose from the lava 

dome (fig. 1B) and passed through the spectrometer’s field of 

view.

A MIDAC IR spectrometer with a resolution of 0.5 cm-1

was mounted onto a 20-inch Newtonian telescope on a heavyweight tripod (fig. 1A). The spectrometer design incorporates a 

Michelson interferometer, laser optics, CaF2 optical windows, 

and an InSb Stirling-cycle cooled detector with a range of 

5,000–1,800 cm-1. The field of view of the instrument is 2–3 

m at 800 m. A laptop computer and PCMCIA interface enable 

real-time data acquisition and processing, using Autoquant4 

software (MIDAC Corp.). Analysis is based on a linear model 

of the sample absorbance known as Beer’s Law, which can be 

written as:

∑=

=

M

j

i aA j

1

Lij C i = 1, 2...N, (1)

where M is the number of compounds assumed to absorb 

in the spectral region analyzed; Ai is the observed sample 

absorbance at the ith frequency, Cj is the (unknown) concentration of the jth component of the mixture, L is the absorption 

pathlength used in recording the sample spectrum and aij is 

the absorptivity of the jth compound at the ith IR frequency. 

Ai, the absorbance spectrum, is obtained from the measured 

spectrum by normalizing the measured spectrum, Si, by a 

background spectrum B, which is free of, or contains less, 

absorption due to volcanic gases:





 



 −= B

S A i i log . (2)

Figure 1. Fourier-transform infrared (FTIR) spectrometer in field settings. A, At FTIR measurement site (Brutus), on east crater rim of 

Mount St. Helens. B, Collecting radiation from hot cracks in the dacitic lava dome of Mount St. Helens, 2005. Volcanic gas is passing 

into the path in front of the infrared source.

A

Aperture

IR from lava

dome

Spectrometer

Telescope

Crater rim

PC acquisition

B

Spectrometer

Telescope

2004–2005 lava dome

Gas plume

West crater rim

The absorbance spectra, therefore, contain information 

pertaining only to the volcanic gas phase; all other background 

gases are canceled out by this procedure. The absorbance 

spectra, Ai, are baseline corrected (in each fitting region) 

before analysis, using a second-order polynomial, and they 

also are corrected for their temperature and pressure relative 

to the reference spectra. A calibration set of spectra of pure 

components at known concentrations are used to calculate the 

absorptivities aij.

The uncertainty in the measurements is calculated from 

the residual spectrum, which should have a mean of zero in the 

fitted region and is proportional to the square root of the sum 

of the squares of the residual:

( )1

2 2

−

= ∑

n

R

i

 i , (3)

where σ2 is the error variance, R is the residual spectrum, and 

n is the number of observations. The standard error of the 

estimated concentration, ε, is equal to:

∑

=

i

Ai

C

2

. (4)

This technique allows the evaluation of precise concentration-pathlengths (LCj from equation 1) for each species, 

which have units of parts per million meters (ppm·m). The 

ratios of the concentration-pathlengths for different species 

in the plume typically are accurate to within a few percent, 

given appropriate analysis. The absolute values, however, are a 

minimum because the background spectrum, B, is not a clear-
576 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1

2 3

200

TEMP.

(

oC)

10

Figure 2. Shaded-relief digital elevation model of September 1, 2005, showing three paths used for OP–FTIR 

measurements in this study (courtesy of S.P. Schilling, USGS). Lengths of paths are labeled. Datum and projection are 

NAD83, UTM zone 10.

Figure 3. Thermal image, taken by 

forward-looking infrared radiometer (FLIR), 

showing lava dome from a helicopter near 

OP–FTIR measurement site (Brutus) on 

September 2, 2005 (courtesy of M. Logan, 

USGS). Numbers mark the parts of the 

lava dome used as infrared sources for 

measurement sets 1 to 3. Temperature 

scale on right.

0

-122°12'

46°12'30"

46°11'30"

46°12'

-122°11'30" -122°11' -122°10'30"

500

METERS

1980−1986 lava dome

2004−2006

lava dome

1,039 m

1,121 m

OP–FTS

“Brutus”

1,152 m
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 577

sky spectrum for these measurements; rather, it is the spectrum 

containing the fewest molecules of volcanic gas in the series 

of measured spectra, Si. A background spectrum under the 

exact conditions of the measurements containing no volcanic 

gas often is impossible to obtain in practice in this mode. 

Spectra were acquired once per second, and eight spectra were 

averaged, a procedure which typically took about 8 s.

In this study, retrievals for seven major volcanic gas species were carried out: H2

O, CO2, HCl, SO2, HF, CO, and COS. 

Retrievals for HCl and CO are relatively low in error, owing 

to their strong and fast absorption features; COS is detectable 

owing to its high absorptivity. Measured SO2 exceeds detection 

limits, but the errors are larger owing to its weak absorption 

signature (at around 2,500 cm-1). H2O has a moderately high 

error owing to its high and variable background concentration 

derived from nonmagmatic, highly variable local sources, such 

as evaporating ground water, ice, and meteoric water. At these 

long pathlengths, background CO2 is around 280,000 ppm·m, 

and, in order to be detected, volcanic CO2 must be present at 

levels >5 percent above background levels. The errors and 

detection limits on the measurements are shown in table 1.

The results of the measurements take the form of molar 

ratios and gas compositions. The molar ratios are calculated 

by plotting the abundance (in concentration-pathlength units) 

of one gas species against another. The gas species will plot 

in a straight line if they are derived from a common origin 

(the volcanic vent) and a gas of constant composition is being 

emitted. All background gases, including random amounts 

of nonvolcanic H2O, are eliminated from the results, owing 

to the normalization by background spectra. Any mixing of 

the volcanic gas with ambient air should, therefore, not affect 

the results unless chemical reactions are taking place, or 

unless there is significant heterogeneity in the atmospheric air 

composition in the pathlength (which is significant here for 

H2O). The mean slope of the plots of gas abundances is equal 

to the molar ratio between the two species. Changes in the 

concentration-pathlength of a species with time are controlled 

by source effects (variations in the emission rate of gases), 

combined with fluctuating wind speeds and directions, which 

Species Wavenumber window (cm-1) Detection limit (ppm·m) Mean error (%)

H2O 2134.80–2140.51 4,000 13.2

CO2 2264.26–2273.20 2,000 15.3

HCl 2744.18–2848.38 3 7.6

SO2 2447.36–2532.96 15 17.4

HF 4062.31–4089.16 20 14.1

CO 2089.29–2199.87 0.8 8.2

COS 2093.60–2020.87 0.7 7.6

Table 1. Summary of wavenumber windows used in analysis, detection limits, and mean errors for 

volcanic gas species detected by OP–FTIR in this study, Mount St. Helens, Washington, 2005.

blow the plume gases out of the pathlength. Pairs of species 

concentration-pathlengths are used to calculate molar ratios, 

thereby eliminating these source and wind effects, as they will 

affect all gases similarly.

Emission rates for SO2 were measured by absorption 

spectroscopy with a miniature spectrometer manufactured 

by Ocean Optics and an instrumental housing (incorporating calibration cells; FLYSPEC) and software developed by 

Horton and others (2006). Horizontal traverses were made 

beneath the plume, approximately 1.5 km downwind of the 

lava dome. Integrated concentration-pathlength across the 

plume width was multiplied by the plume speed (measured by 

using methodology developed by Doukas, 2002) to generate 

SO2 emission rates. Three measurements of SO2 emission rates 

were carried out, and errors are less than 10 percent.

Results

The spectra measured on August 31, 2005, record seven 

volcanic gas species: H2

O, CO2, HCl, SO2, HF, CO, and COS, 

in typical order of abundance for the three sets of measurements (fig. 2). The first set of measurements contained 109 

spectra. The IR source for the measurements was part of spine 

6, and the pathlength was 1,150 m (figs. 2, 3). These spectra 

recorded absorptions due to H2O, HCl, CO, and COS gases; 

CO2, SO2, and HF were below detection, possibly due to a 

relatively low source temperature and (or) the long pathlength, 

both of which tend to decrease the signal-to-noise ratio. Figure 

4 shows the characteristic absorption signature due to HCl gas 

in the measured spectra compared to a reference spectrum, the 

concentration-pathlength of H2

O and HCl, and the concentration-pathlengths of CO and COS with time for this set of measurements. Table 2 shows a summary of the data (labeled set 

1) and includes mean and maximum concentrations measured 

for each species and the mean H2O/Cl ratio. The relatively low 

abundances of all the detected gases (HCl, as much as 147 

ppm·m, and H2

O, as much as 21,000 ppm·m) suggest that the 
578 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

abundances of CO2 and SO2 were probably below detection for 

this set of measurements. The mean H2O/Cl molar ratio was 

120 (table 2).

The second set of measurements were taken using a 

shorter pathlength of 1,040 m and a remnant of spine 5 as an 

IR source (figs. 2, 3). Figure 5 shows that HCl, HF, CO, and 

COS are detectable during the entire time period, whereas 

H2O, CO2, and SO2 are detectable in only some of the spectra. 

Nondetection in some parts of the time series might be due 

A

B

C

2,850

1035 1040 1045 1050 1055

1035 1040 1045 1050 1055

30,000

25,000

20,000

15,000

10,000

5,000

0

7.0

180

150

120

90

60

30

0

2.5

2.0

1.5

1.0

0.5

0

6.0

5.0

4.0

3.0

2.0

1.0

0

2,800 2,750 2,700 2,650

Reference spectrum

Measured spectrum

WAVENUMBER, CM-1

 PACIFIC DAYLIGHT TIME (AUGUST 31, 2005)

H2O, IN PPM·M

HCl, IN PPM·M

CO, IN PPM·M

ABSORBANCE

COS, IN PPM·M

HCl

H2O

CO

COS

to a weakening of the IR source or a decrease in gas-column 

abundance. The abundance of HCl ranged from 0 to 624 

ppm·m over the pathlength, whereas SO2 reached 251 ppm·m 

(table 2, set 2). Carbon monoxide reached a maximum concentration-pathlength of 15 ppm·m, and COS reached 2.6 ppm·m. 

The mean Cl/S ratio was 2.1, and H2O/Cl was 85 (table 2).

The third set of measurements was taken using an IR 

source on spine 6, with a pathlength of 1,120 m (figs. 2, 3). 

This set of measurements imaged gases emitted from the 

Figure 4. Data for measurement set 1, Mount St. Helens, Washington, 2005. A, Measured 

absorbance spectrum and reference spectrum for HCl gas for spectra collected along path 1. B, 

HCl and H2

O concentration pathlengths. C, CO and COS concentration pathlengths with time for 

measurement set 1.
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 579

Set HCl SO2 CO CO2 COS H2O HF Cl/S Cl/F H2O/Cl SO2/COS CO/CO2

1

N=109

(LC)

max 147 bd 5.1 bd 2.0 21,000 bd

(LC)

mean 57.0 bd 1.7 bd 0.47 7,000 bd -- -- 120±25 -- --

2

N=479

(LC)max 624 251 15 10,300 2.6 51,000 310

(LC)

mean 279 145 5.0 5,870 1.2 28,900 74 2.1±0.4 8.8±2.0 85±15 120±15 0.00086±0.0002

3

Active vent

N=97

(LC)

max 261 155 1.3 4,400 0.84 310,000 86

(LC)

mean 110 121 0.42 1,300 0.35 64,000 40 1.1±0.3 3.5±2.0 1040±200 350±60 0.00032±0.00006

Table 2. Summary of OP–FTIR measurements (sets 1, 2, and 3), Mount St. Helens, Washington, 2005.

[N is number of spectra in each set. Maximum (LC)

max and mean (LC)mean are concentration pathlengths (in ppm·m) for the seven gas species detected; bd, 

below detection. Mean molar ratios are given with estimates of errors. Dashes show ratios not calculated where sulfur data are lacking.]

area closest to the active vent at the base of the most recently 

erupted spine. H2O, CO2, HCl, SO2, HF, CO, and COS were 

detected in the absorption spectra (fig. 6 and table 2). The 

abundances of the gases were generally less than for measurement set 2; HCl concentration pathlengths reached 261 

ppm·m, and SO2 reached 155 ppm·m (table 2, set 3). The mean 

molar H2O/Cl was 1,040 and Cl/S was 1.1. The molar ratios 

can be used to formulate a gas composition (table 3, column 

A). The gas was composed mainly of H2O (99 mol percent), 

with 0.78 mol percent CO2, 0.095 mol percent HCl, 0.085 

mol percent SO2, and minor amounts of HF, CO, and COS. 

For comparison, a gas sample collected in September 1981 

(Gerlach and Casadevall, 1986) is shown in column D. Columns B and E show the same gas compositions recalculated 

to H2

O-free compositions, and columns C and F show them 

recalculated to H2O- and CO2-free compositions. The gas composition measured in 2005 is similar to that measured in 1981, 

albeit with less CO and more COS.

The ultraviolet spectrometer traverses carried out at 

1400 PDT on August 31, 2005, yielded a mean SO2 flux of 

22 t/d. By converting the molar gas composition into the mass 

composition (table 4) for measurement set 3, the SO2 emission rate (bold, table 4) can be used to calculate emission 

rates for the six other species. On August 31, 2005, H2

O was 

emitted at rates of about 83 kg/s (7,200 t/d), CO2 at 1.6 kg/s 

(140 t/d), HCl at 0.16 kg/s (14 t/d), and HF at 0.024 kg/s (2.0 

t/d). CO and COS were emitted at rates of 0.63 and 0.69 g/s, 

respectively; these fluxes are associated with errors of ~±20 

percent. CO2

 emission rates were measured independently 

by using LICOR and were found to be 198 t/d (Gerlach and 

others, this volume, chap. 26). The excess can be attributed to 

a small amount of CO2

 degassing from the 1980–86 lava dome 

(Gerlach and others, this volume, chap. 26), which was not 

captured by the OP–FTIR measurements.

Discussion

Partitioning of Cl into Aqueous Vapor

Experimental work on Cl behavior in silicic magmas 

demonstrates the strong affinity that Cl has for a H2O-rich 

vapor phase (Kilinc and Burnham, 1972; Webster and Holloway, 1988). The fluid-melt partition coefficient increases with 

decreasing temperature and with increasing dissolved melt Cl 

content; and decreases with decreasing pressure (Shinohara 

and others, 1989), although the latter is not well constrained. 

The partition coefficient is high for subaluminous melts and 

decreases with increasing aluminous or peralkaline character 

(Webster, 1992). Increasing the Cl molality of the vapor phase 

to high levels causes the formation of an immiscible Cl-rich 

liquid, as observed in the H2

O-NaCl system (Shinohara and 

others, 1989). Signorelli and Carroll (2002) determined Cl 

solubility limits of 4,800–6,800 ppm for the hydrous rhyolite 

melt of recent Soufrière Hills volcano eruptions, which has a 

(Na+K)/Al ~0.6. These concentrations are much higher than 

those observed in melt-inclusion glasses (Edmonds and others, 
580 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

B

C

1050 1100 1110 1120 1130 1140 1150 1200 1210 1220

10,000

20,000

30,000

40,000

50,000

60,000

70,000

800

700

600

500

400

300

200

100

0

0

20.0

15,000

18,000

12,000

9,000

6,000

3,000

0

3.5

2.0

2.5

1.0

0.5

2.0

1.5

0

15.0

COS

CO

H2O

HCl

SO2

HF

CO2

10.0

5.0

0

1050 1100 1110 1120 1130 1140 1150 1200 1210 1220

1050 1100 1110 1120 1130 1140 1150 1200 1210 1220

PACIFIC DAYLIGHT TIME (AUGUST 31, 2005)

H2O, IN PPM·M HCl, SO2, HF IN PPM·M

CO2, IN PPM·M

CO, IN PPM·M

COS, IN PPM·M

Figure 5. Data for measurement set 2, Mount St. Helens, Washington, 2005, as concentration pathlengths 

with time. A, HCl, HF, and SO2. B, H2O and CO2. C, CO and COS.
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 581

HCl, SO2, IN PPM·M CO2, IN PPM·M

H2O, IN PPM·M

HCl, IN PPM·M

0

100

200

300

400

A

B

C

1210

0

1,000

2,000

3,000

4,000

5,000

6,000

0

200,000

100,000

HCl

SO2

H2O

CO2

HCl

HF

300,000

400,000

HF, IN PPM·M

0

60

40

20

80

100

0

50

250

200

150

100

300

350

1215 1220 1225 1230

1210 1215 1220 1225 1230

1210 1215 1220 1225 1230

PACIFIC DAYLIGHT TIME (AUGUST 31, 2005) 

Figure 6. Data for measurement set 3, Mount St. Helens, Washington, 2005, as concentration pathlengths with 

time. A, HCl and SO2. B, H2O and CO2. C, HCl and HF.
582 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

This study (set 3) September 17, 1981

A B C D E F

H2

O 99 99

CO2 0.78 79 0.89 83

HCl 0.095 9.7 46 0.076 7.2 43

SO2 0.085 8.6 41 0.067 6.3 38

HF 0.027 2.8 13 0.030 2.8 17

CO 0.00048 0.049 0.23 0.0023 0.22 1.3

COS 0.00025 0.025 0.12 1.8×10-5 0.0017 0.010

Table 3. Gas compositions at Mount St. Helens, 1981 and 2005.

[Compositions (in molar percent) derived from this study (column A) are compared with a representative gas composition derived from sampling of gases by using caustic soda in May 1981 (column D; Gerlach and Casadevall, 

1986). Columns B and E show the gas compositions recalculated to H2O-free compositions. Columns C and F 

show the gas compositions recalculated to H2O- and CO2-free compositions.]

Species Molar % Mass % Flux (kg/s) Flux (t/d)

H2

O 99 98 83 7,200

CO2 0.78 1.9 1.6 140

HCl 0.095 0.19 0.16 14

SO2 0.085 0.30 0.25 22

HF 0.027 0.028 0.024 2.0

CO 0.00048 0.00073 0.00063 0.054

COS 0.00025 0.00080 0.00069 0.059

Table 4. Gas composition derived from this study (from measurement 

set 3), Mount St. Helens, Washington, 2005, in molar percent and mass 

percent.

[Flux of each species is calculated using the flux of SO2 gas measured with ultraviolet 

spectrometers (22 t/d; in bold). Fluxes expressed both in kg/s and in t/d.]

2002) and suggest that the aqueous vapor phase was relatively 

Cl poor. In the case of Mount St. Helens, melt-inclusion data 

for the erupted products of 1980–86 indicate that the maximum Cl content of the preeruptive melt was around 1,000 ppm 

(Rutherford and others, 1985), which is within the range of Cl 

concentrations in melt inclusions from the 2004–5 

eruption of Mount St. Helens (Pallister and others, this volume, 

chap. 30) and is much lower than the solubility for Cl in a 

melt with (Na+K)/Al ~1 at 850–900°C and 200 MPa (around 

2,000–2,500 ppm; Metrich and Rutherford, 1992). If the 

2004–5 Mount St. Helens melt had an initial H2

O content of 

~4.6 weight percent, similar to that obtained from melt-inclusion analysis for the 1980s erupted products (Rutherford and 

others, 1985), and if Ptotal is equal to PH2

O, then the melt was in 

equilibrium with a H2O-rich vapor at a depth of around 6 km 

(assuming lithostatic pressure with depth). The preeruption 

dissolved Cl content of the melt at about 6 km is assumed to 

be close to 1,000 ppm, on the basis of the maximum Cl concentration measured in melt inclusions (Pallister and others, 

this volume, chap. 30), which is within range of typical silicic 

arc magmas (Wallace, 2005).

The H2O/HCl molar ratio of the vapor varies during melt 

ascent and eruption and is controlled by water solubility and 

degassing, partitioning of Cl into the vapor, the degassing 

regime (whether open- or closed-system degassing or some 

combination of the two is in operation), and microlite crystallization on ascent (which will tend to increase the concentration of volatiles in the melt and promote further degassing). 

These factors can be incorporated into models describing the 

evolution of the vapor phase on ascent and compared with the 
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 583

observations presented here. This analysis assumes there is 

no breakdown of hydrous phases (for example, amphibole), 

which would tend to increase the bulk H2O of the system on 

magma ascent.

Closed-system degassing occurs when volatiles exsolve 

and remain in contact and in equilibrium with the host melt. 

The vesicularity of the melt is related to the amount of volatiles exsolved and the total pressure. Closed-system degassing can occur (1) during rapid ascent of magma from depth, 

resulting in fragmentation when the overpressure inside the 

expanding bubbles exceeds the strength of the magma, (2) 

in static magma bodies as a result of second boiling, and (3) 

during slow ascent of magma at depths of greater than a few 

kilometers, where gas fractions will typically be too low for 

fragmentation to occur. The partitioning of Cl between melt 

and aqueous vapor can be used to describe closed-system 

degassing:

( )OHOH meltfluid X D

X XX  −



 



 =− −1 0 0

2 2 , Cl 

Cl 

Cl 

 (5)

where 0 X 2OHand X 0

Cl are the original H2

O and Cl mass fractions, X 2OH and XCl are the melt H2

O and Cl mass fractions, 

and DCl

fluid–melt­is the vapor-melt coefficient for Cl (assumed 

here to be constant). X 2OH varies with pressure according 

to its solubility, which is calculated using VolatileCalc for a 

rhyolite melt at 920°C and with a 0 X 2OH of 4.6 weight percent 

(Newman and Lowenstern, 2002).

Open-system degassing involves increments of vapor 

being removed from the host melt after exsolution. For 

volatiles that partition strongly into a vapor phase, this style 

of degassing results in efficient removal of the species from 

the melt. This style of degassing may occur (1) along conduit 

sidewalls, where exsolved volatiles can escape by migrating 

upward or through the conduit walls, or (2) during shallow, 

slow ascent of magma, when sufficient porosity develops to 

allow permeability and vapor escape through a three-dimensional bubble network (for example, Blower, 2001); significant 

gas loss through bubble ascent would be suppressed by the 

high viscosity of the magma and crystals. Open-system degassing can be modeled using a Rayleigh fractionation equation 

(for example, Villemant and Boudon, 1999):

fXX 0

Cl Cl = , (6)

where = ( ) −1 −meltfluid DCl , (7)

and ( ) 



 



 = 0

0

2

2 2 - -1 OH

OHOH

X

XX f , (8)

where f is the fraction of water remaining in the melt. XCl is 

calculated for 5-MPa increments in pressure and converted to 

depth (km) assuming lithostatic pressures and a mean crustal 

density of 2,500 kg/m3.

In order to account for microlite crystallization on ascent, 

which decreases the melt volume and causes increased exsolution of water, a factor q was introduced to the equation for the 

open-system degassing case, whereby

( ) 













 − −= 0

0

00

2

2

OH

OHOH

X

XX

qqq 2 , (9)

where 0 q is the initial melt fraction (taken to be 0.6). The 

expression for f, the fraction of H2O remaining in the melt, is 

modified thus:

( ) 



 



 − 













  −



 



 =

−

OHOH

n

n

OH

OHOH XX

q

q

X

XX f 2 2

2

2 2

1

0

0 - -1 , (10)

where n q is the melt fraction for the previous step and qn−1 is 

the melt fraction for the current step and qn−1 < n q . This has the 

effect of decreasing f to a slightly greater degree at each step 

to account for the effects of microphenocryst and microlite 

crystallization.

Figure 7 shows the results for closed- and open-system 

degassing for three partition coefficients (D). These results show 

the large dependence of the vapor composition at the surface 

on degassing regime. The plots show the depletion of Cl in the 

melt, the ratio of dissolved H2O/Cl, and the H2O/Cl ratio in the 

vapor phase, with depth. Three measured parameters are marked 

on the plots to compare to the model data: (1) the mass fraction 

of dissolved Cl, (2) the H2O/Cl in the matrix glass of the erupted 

products (for microcrystalline and flow-banded dacite; from 

Pallister and others, this volume, chap. 30; summarized in table 

5), and (3) the mass ratio H2O/Cl in the volcanic gases (converted from molar ratios given in table 2). The H2

O/Cl in gases 

from measurement sets 1 and 2 range from 35 to 70 and are 

labeled stagnant dome (SD), and those from measurement set 3 

range from 430 to 630 and are labeled active dome (AD).

The mass fraction of Cl in the melt, dissolved melt H2O/

Cl, and vapor H2O/Cl from the stagnant dome are consistent 

with closed-system degassing, with a large partition coefficient 

(D~50) for Cl, from depths of ~ 6 km. The active dome gases 

are richer in H2

O and are closer to the modeled H2O/Cl for an 

open-system degassing regime. Open-system degassing from 

depth would deplete the melt almost entirely of Cl, however, 

which is not the case (figs. 7A, B); open-system degassing 

from depth is therefore inconsistent with observation. Closedsystem degassing from depth with constant D is consistent 

with observation but unrealistic on two counts: (1) closed-system degassing could not physically lead to lava dome building, 

owing to the inevitably large gas fractions that would result 

at shallow depths, and (2) a constant partition coefficient for 

Cl is improbable, based on the experiments of Shinohara and 

others (1989), which suggests that D is strongly dependent on 

both dissolved melt Cl and on pressure.

Recent studies of degassing during the lava dome building of Soufrière Hills volcano, Montserrat, have shown that 

the degassing of H2

O in the upper conduit leads to cooling, 
584 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

MASS FRACTION Cl, IN MELT MELT H2O/Cl VAPOR H2O/Cl

DEPTH, IN KILOMETERS

0

1

2

3

4

5

6

0.00001 0.0001 0.001 1 10 100 1,000 10,000 100,000 10 100 1,000

A B C

50

10

10

3

3

3

10

10

10

5

5

5

2

2

2

50

10

50

MCD MCD

SD AD

FBD

FBD

FBD

MCD

SD

AD

OBSERVED

Flow banded dacite

Microcrystalline dacite

Stagnant dome

Active dome

CLOSED

OPEN

Matrix glass Cl 

(ppm)

H2O 

(wt%) H2

O/Cl

Microcrystalline dacite

 Mean concentration 278 1.40 41

 1σ deviation 53 0.98 24

Flow-banded dacite

 Mean concentration 853 2.40 28

 1σ deviation 47 0.29 4.5

Figure 7. Effect on evolution of gases with depth for various models of degassing regime. Solid lines indicate closed-system 

degassing models and are labeled with the fluid-melt partition coefficient, D. Dot-dash lines indicate open-system degassing 

models and are labeled with the fluid-melt partition coefficient. Blue boxes indicate observed melt Cl, melt H2

O/Cl, and vapor 

H2

O/Cl. A, Fraction of Cl in melt. B, Dissolved melt H2O/Cl. C, vapor H2O/Cl.

Table 5. Mean concentrations of Cl and H2

O in erupted 

matrix glasses of microcrystalline (from 5 measurements) 

and flow-banded (from 12 measurements) dacite, Mount 

St. Helens, Washington, 2004–2005.

[From Pallister and others, this volume, chap. 30.]

crystallization, and an increase in viscosity of the magma 

(for example, Sparks and others, 2000). The development of 

permeability during degassing in the top ~1 km of the conduit 

leads to gas loss and effusive eruption. In order to investigate 

whether this mechanism is consistent with the proportions 

of H2O and Cl in the melt and in volcanic gases at Mount St. 

Helens, the model is adjusted by using equation 5 in order to 

incorporate the effects of closed-system degassing at depths 

greater than 1 km with the same mean fluid-melt partition 

coefficients for Cl (3, 10, and 50), followed by open-system 

degassing along the top 1 km of the conduit (fig. 8), using 

equations 6–10 with fluid-melt partition coefficients of 1, 2, 

and 3 ( 0 X 2OH and X 0

Cl are set equal to the final X 2OH and XCl 

given by the closed-system degassing model at 1-km depth; 

fig. 8). A partition coefficient of 1 represents the case whereby 

no further degassing of Cl occurs. The plots in figures 7 and 8 

are, therefore, identical up to 1-km depth; at shallower depths, 

the effects of open-system degassing are shown in figure 8. 

Observations of the melt and gases are marked on figure 8 to 

compare to the models. The mass fraction of Cl remaining dissolved in the melt, the melt H2

O/Cl, and the H2O/Cl for gases 

emitted from the stagnant part of the lava dome are consistent 

with deep, closed-system degassing (D~10), followed by 

open-system degassing with D~2–3 (fig. 8). A decrease in D

to <5 was noted by Shinohara and others (1989) for lithostatic 

pressures corresponding to depths less than 2 km.

The H2

O/Cl of gases emitted from the active part of the 

lava dome, however, are not consistent with any of the theoretical models formulated here. This may be due to one or all 

of three factors: (1) a cessation of Cl degassing in the top 1 km 

of the conduit, (2) addition of substantial amounts of meteoric, 

ground-water, and glacier-derived H2O to the gases at shallow 
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 585

depths, or (3) some Cl present as alkali chloride (NaCl and 

KCl) in the gas phase, as observed by Le Guern and Shinohara 

(1985) and studied experimentally by Shinohara and others 

(1989). The proportions of dissolved H2O and Cl remaining 

in the erupted glasses suggests that there is no cessation of 

Cl degassing in the top 1 km of the conduit (fig. 8; compare 

observations to solid line presenting D=1), so this factor is discounted. Shevenell and Goff (1999) used oxygen- and hydrogen-isotope data from hot springs to show that only 30–70 mol 

percent of the H2O in the fumaroles around the lava dome in 

1981–86 was magmatic. In this case, the discrepancy between 

models and observations requires that only ~10 mol percent of 

the measured H2O be magmatic. Modification of the gas composition is shown in table 3, column A; using this factor yields 

91 mol percent H2O and 7 mol percent CO2, which might be 

more representative of the magmatic gas composition. The 

presence of alkali chlorides in the gas phase is consistent with 

the presence of chloride encrustations on the lava dome (possibly halite and (or) sylvite) observed during 2004–5 (C.R. 

Thornber, oral comm. 2005). Yellow encrustations on the 

September 1980 lava dome were noncrystalline and composed 

of Al, Cl, Fe, Ca, S, and H2

O. A sublimate film near fumaroles 

<250°C produced a yellow-brown stain on crack walls, which 

had a high Cl content (Keith and others, 1981).

MASS FRACTION Cl, IN MELT MELT H2O/Cl VAPOR H2O/Cl

DEPTH, IN KILOMETERS

0

1

1

2

3

2

3

4

5

6

0

1

2

3

4

5

6

0

1

2

SD

MCD MCD

FBD

FBD AD

3

4

5

6

0.00001 0.0001 0.001 1 10 100 1,000 10 100 1,000

A B C

Shallow open-system

 degassing, with D

Deep, closed-system

 degassing

Labeled D

OBSERVED

Flow-banded dacite

Microcrystalline

 dacite

Stagnant dome

Active dome

50

10

3

50

50

10

10

3

3

FBD

MCD

SD

AD

Figure 8. Effect of closed, then open-system (in the upper 1 km of the conduit) degassing on the composition of volcanic gases at 

the surface. The evolution of (A) the fraction of Cl in the melt, (B) dissolved melt H2

O/Cl, and (C) vapor H2O/Cl with depth (km) for 

various models of degassing regime are shown. Closed-system degassing regimes are indicated by solid lines up to 1 km depth, 

labeled with values of D. Above 1 km depth, open-system degassing regimes are shown as dashed-dotted lines for three different 

values of D (see key). Observed melt Cl, melt H2

O/Cl, and vapor H2O/Cl are shown as blue boxes.

Comparison and Significance of Molar Cl/S

The behavior of S in silicic melts differs significantly 

from both Cl and H2O. The partitioning of sulfur into a vapor 

phase is controlled by several factors, the most significant 

being the fugacities of oxygen and sulfur in the vapor phase 

and melt composition (Carroll and Webster, 1994). The solubility of sulfur is much lower than that of Cl for a given set 

of magmatic conditions. Table 6 shows molar Cl/S measured 

using a variety of techniques during eruptions at a range of 

dacitic and andesitic subduction-zone volcanoes, including 

the 1980–86 Mount St. Helens lava-dome eruption. The molar 

ratio measured during this study is similar (within error) to 

ratios measured at most other volcanoes and to ratios measured during the 1980–86 Mount St. Helens eruption, with 

the notable exception of Soufrière Hills volcano, which shows 

considerably higher Cl/S during lava dome-building activity. A molar Cl/S of ~1 appears to be a dominant feature of 

most reported gas compositions measured during eruptions in 

subduction-zone settings.

Botcharnikov and others (2004) reported solubility data 

for melts coexisting with mixed H2

O-Cl-S gases to illustrate 

the effect of multiple volatile components on their solubility. 

The total Cl content of the melt-vapor system can affect the 
586 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Volcano Type Method Cl/S

Usu-san, Japan1 Basaltic andesite OP–FTIR 0.8

Mount St. Helens, USA2 Dacite EBS 1.2

Mount St. Helens, USA3 Dacite OP–FTIR 1.1

Augustine Volcano, USA4 Dacite EBS 1.0

Unzen-dake, Japan5 Dacite OP–FTIR 0.9

Soufrière Hills, Montserrat6 Andesite OP–FTIR 2.0

1

Mori and others, 1997.

2

Gerlach and Casadevall, 1986.

3

This study.

4

Symonds and others, 1990.

5

Mori and others, 1993.

6

Edmonds and others, 2002.

Table 6. Molar Cl/S ratios in volcanic gases measured at other silicic subduction-zone 

volcanoes and at Mount St. Helens, Washington, 2005.

[Errors are estimated as ranging from 5 to 10 percent. OP–FTIR, open-path Fourier-transform IR spectroscopy; EBS, evacuated bottle sampling.]

solubility of S (Botcharnikov and others, 2004). The addition 

of 1 weight percent Cl increased S solubility by two times at 

850°C and 200 MPa for a dacite-composition melt, with all 

other conditions constant. For a given H2O vapor concentration, over a range of total S content and solubility, a maximum 

S solubility occurred corresponding to a vapor Cl/S of ~1. A 

vapor Cl/S of ~1 may, therefore, be a direct consequence of 

the presence of moderate amounts of Cl in the melt, which is 

typical for arc magmas; vapor Cl/S might further be buffered 

by the addition or depletion of small amounts of Cl to the 

system. The maximum in S solubility might be explained by 

two main factors (Botcharnikov and others, 2004): (1) S is 

dissolved as sulphate (SO4

2 – ) in hydrous melts, where oxygen 

fugacity is relatively high, and as sulfide ( 2− S ) at low oxygen 

fugacities (Carroll and Webster, 1994); the transition occurs 

at oxygen fugacities close to NNO. The addition of Cl to an 

S-bearing system might shift the redox conditions to slightly 

more oxidized conditions. The other possibility is (2) that 

S-Cl-bearing complexes may form in the melt, which would 

increase the solubilities of both components.

The Abundance of Minor Gaseous Components 

(HF, CO, COS)

The solubility of fluorine (F) in silicic melts is generally 

much higher than that of Cl and, hence, does not degas to the 

same extent (Carroll and Webster, 1994). The vapor-melt partition coefficient for F is usually less than unity (Webster, 1990). 

Table 2 shows that the molar ratio of Cl/F in the volcanic gases 

is around 3.5. A similar molar Cl/F has been reported at a 

number of other silicic volcanoes in subduction-zone settings 

(Symonds and others, 1994) and compares well with the abundance of HF emitted from fumaroles at Mount St. Helens in 

September 1981 (table 3; Gerlach and Casadevall, 1986).

The mean CO/CO2

 ratio of the gases of around 3.2×10-4

(table 2) implies that the gases equilibrated at a temperature of 

about 570°C, on the basis of the relation:

log CO

CO2

 

     =1.76 − 4417

T , (11)

where T is the temperature in Kelvin (Mori and Notsu, 1997). 

Collection temperatures derived from similar relations in gases 

sampled in 1981 ranged from 620°C to 762°C (Gerlach and 

Casadevall, 1986).

COS has been detected previously in volcanic plumes 

(for example, Le Guern and others, 1975), although measurements are few. Rasmussen and others (1982) reported 

abundances of 5 ppm by volume in the Mount St Helens gas 

plume of May 18, 1980, and suggested that injection of COS 

contributes to the stratospheric S burden during large silicic 

eruptions. Gerlach and Casadevall (1986) report molar SO2/

COS of 2,000 to 4,750 for gas samples collected at Mount St 

Helens from May to September 1981 from radial fumaroles 

around the lava dome. The mean molar ratios obtained August 

31, 2005, using OP–FTIR are much lower than they were in 

the earlier samples, 350 (table 2, set 3) and 120 (set 2). Mori 

and Notsu (1997) reported SO2 / COS of 400 for gases emitted 

from a fumarole at Aso, Japan, at a temperature of ~740°C.
27. Chlorine Degassing During the Lava Dome-Building Eruption of Mount St. Helens, 2004–2005 587

Conclusions

Several conclusions can be drawn from the results and 

interpretation presented here:

•	 The fluxes of magmatic volatiles from Mount St. 

Helens on August 31, 2005, were 7,200 t/d H2O, 

140 t/d CO2, 14 t/d HCl, 22 t/d SO2, 2.0 t/d HF, 54 

kg/d CO, and 59 kg/d COS.

•	 The H2O/Cl ratio in the gases emitted from 

stagnant parts of the Mount St Helens lava 

dome, and the amount of Cl and the H2O/Cl in 

the erupted glasses in the melt, are all consistent 

with a model of closed-system degassing up to ~ 

1-km depth (with a fluid-melt partition coefficient, D, for Cl of ~10), followed by open-system 

degassing to the surface (with a D of ~2–3). This 

is similar to models proposed for lava domebuilding eruptions elsewhere, where much of the 

gas loss occurs in the upper 1 km of the conduit, 

causing cooling, crystallization, and an increase 

in magma viscosity.

•	 The gases emitted from the active part of the 

Mount St Helens lava dome have an H2O/Cl that 

is about one order of magnitude higher than that 

predicted for closed- or open-system degassing or 

the two-stage model above. This is explained by 

as much as 90 mol percent of the measured H2O 

vapor being nonmagmatic (ground- or glacierderived) and by some of the Cl being present as 

alkali chlorides in the gas phase (NaCl and KCl).

•	 Molar Cl/S in the gases measured in August 2005 

is similar to that measured during 1980–86 at 

Mount St. Helens and at silicic subduction-zone 

volcanoes elsewhere during similar styles of eruption. This similarity may be caused by a maximum 

in S solubility when vapor Cl/S ~1, which is valid 

for the typical range in initial S and Cl (100–300 

and 1,000–3,000 ppm, respectively) at silicic 

subduction-zone volcanoes.

•	 The gas analyses presented here confirm that 

remote spectroscopic techniques are capable of 

measuring the composition of high-temperature 

gases remotely and safely during lava dome 

building at silicic volcanoes, provided the measurements can be made using a pathlength ≤1 km 

with sufficient gas molecules in the path to enable 

detection. There is scope for automated measurements with more compact instruments that 

consume less power, enabling the development of 

time series and further analysis of the degassing 

regime throughout an eruption.

Acknowledgments

This work was done with the support of the USGS 

Mendenhall Postdoctoral Fellowship Program and the Center 

for the Study of Active Volcanoes (CSAV) at the University 

of Hawai`i, Hilo. Michelle Coombs and Jacob Lowenstern 

provided detailed and thorough reviews, which improved the 

manuscript enormously.

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590 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006
Heading 1 591

Petrological and Geochemical 

Investigations of Eruptive Products

Magma—erupt it passively to produce lava flows or domes; erupt it 

explosively and it is deposited as pyroclastic flows and tephra falls. Either 

way the products can be sampled and studied in great detail. Locked in 

the glass and minerals of lava flows and tephra is a complex, sometimes 

fragmentary record of a magma’s temperature and pressure during ascent 

and eruption. Advances in analytical technology now allow the probing 

of single grains for their short-lived radioactive nuclides. The ratios of 

these isotopes vary at known rates, so that an absolute sense of time can be 

imposed on the petrologic processes. Thus texture, mineralogy, and chemistry provide wide-ranging insight into the conditions of magma formation, 

transport, storage, and eruption.

Is the 2004–6 lava the gas-depleted residuum of the 1980s eruption? 

Petrochemical details and short-lived isotopes suggest that it is—that the 

newly erupted dacite is mostly or entirely derived from the same reservoir 

as the 1980s magma.

Is new magma necessary to account for the compositional changes 

seen? That answer is less clear. Some component of hotter dacitic magma 

(possibly injected during earthquake swarms in the 1990s?) could have 

influenced the overall magma composition and shifted the abundance of 

certain indicator elements. Growth patterns and isotopic values tell us that 

a substantial proportion of the plagioclase phenocrysts grew at about the 

time of the 1980s eruption. The iron-titanium oxide geothermometer indicates magmatic temperatures of about 850°C; higher temperatures, above 

1,000°C in some cases, may be explained by recharge of hotter magma at 

depth (which left no telltale geochemical, mineralogical, or petrographic 

signature) or by latent heat released during the rapid and shallow groundmass crystallization. The geochemical fingerprints of the 2004–6 magma 

are too fuzzy to identify a single scenario for the eruption, but they point 

toward a timeline that began more than a decade before magma reached the 

surface in October 2004.
592 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

An unusual perspective, this oblique view west as taken from the International 

Space Station on October 28, 2008. Much of the rugged terrain in the 

amphitheater is the new dome, but the 1980s dome can be recognized. The 

Breach drains north to Spirit Lake, seen at right of center. NASA photo by the 

Expeditiion 18 crew. (http://earthobservatory.nasa.gov/IOTD/view.php?id=36298)
Heading 1 593

A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 28

1

 U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

The Pleistocene Eruptive History of Mount St. Helens, 

Washington, from 300,000 to 12,800 Years Before Present

By Michael A. Clynne1, Andrew T. Calvert1, Edward W. Wolfe2, Russell C. Evarts1, Robert J. Fleck1, and 

Marvin A. Lanphere1

Abstract

We report the results of recent geologic mapping and 

radiometric dating that add considerable detail to our understanding of the eruptive history of Mount St. Helens before 

its latest, or Spirit Lake, stage. New data and reevaluation of 

earlier work indicate at least two eruptive periods during the 

earliest, or Ape Canyon, stage, possibly separated by a long 

hiatus: one about 300–250 ka and a second about 160–35 ka. 

Volcanism during this stage included eruption of biotite- and 

quartz-bearing dacite domes and pyroclastic flows in the area 

west of and beneath the present-day edifice, accompanied by 

the deposition of set C tephras. Ape Canyon-stage rocks are 

compositionally similar to younger Mount St. Helens dacite.

The Cougar stage, about 28–18 ka, was probably the most 

active eruptive stage in Mount St. Helens’ history before the 

Spirit Lake stage. During the Cougar stage, a debris avalanche 

buried the area south of the present-day edifice, and voluminous pyroclastic flows, dacite domes, tephra, and a largevolume pyroxene andesite lava flow were erupted. Two tephra 

sets, M and K, were deposited midway through this stage.

Swift Creek-stage deposits were emplaced in two phases, 

beginning about 16 ka and ending about 12.8 ka. During the 

first phase, set S tephras and three large fans and at least one 

smaller fan of dacitic fragmental material were deposited on 

the northwest, west, south, and southeast flanks of Mount 

St. Helens. The fans are dominated by lithic pyroclastic-flow 

deposits associated with dome building but include both primary and reworked material from pumiceous pyroclastic flows 

and lahars. One Swift Creek-age dome on the west flank of the 

volcano has been located, and others must have been nearby. 

During the second phase, set J tephras were deposited, but no 

pyroclastic flows or domes are known to be associated with 

the andesitic set J tephras.

Preliminary petrographic analysis of these older rocks 

suggests that the volcano’s magmatic system was simpler during the Ape Canyon stage than during subsequent stages and 

that the magmatic system has evolved from relatively simple to 

more complex as the volcano matured. Compositional cycles 

as envisioned by C.A. Hopson and W.G. Melson for the Spirit 

Lake stage probably did not occur during the Ape Canyon stage 

but developed later during the Cougar and Swift Creek stages.

Introduction

The general eruptive history of Mount St. Helens was 

established through the work of U.S. Geological Survey 

(USGS) scientists D.R. Mullineaux and D.R. Crandell. 

Although the edifice was constructed primarily by eruptions 

of lava domes and flows, the volcano’s fallout tephra record 

is extensive; and away from the mountain and its fragmental apron, pumice and ash locally bury the older landscape, 

commonly to a depth of 3 m or more. Detailed study of these 

tephra deposits over a period of 30 years (see Mullineaux, 

1996, and references therein) has established a general 

geochronologic framework for understanding the volcano’s 

eruptive history. Four periods of intermittent volcanism, 

called stages, are recognized (fig. 1) on the basis of observed 

time-stratigraphic groups, or “sets,” of similar-age tephras, 

separated by weathered intervals. Tephra set C is the dominant tephra of the earliest, or Ape Canyon, stage; tephra sets 

M and K mark the Cougar stage, tephra sets S and J mark the 

Swift Creek stage, and tephra sets Y, P, B, and W (and several 

less extensive ash layers) were deposited during the latest, or 

Spirit Lake, stage. An extensive apron of fragmental deposits 

surrounds the volcano and fills valleys draining the mountain 
594 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

(fig. 2). Study of these deposits, which include sequences of 

pyroclastic flows, lahars, and debris avalanches, placed them 

into the general geochronologic framework (see Crandell, 

1987, and references therein).

Age control on the eruptive history of Mount St. Helens was established by radiometric (14C) dating of organic 

material in or directly beneath tephra, pyroclastic, and lahar 

deposits (Crandell and others, 1981; Crandell, 1987; Mullineaux, 1996); however, most dating focused on the youngest and best-preserved part of the volcano. Verhoogen (1937) 

published a crude geologic map, and Hopson (2008) a more 

detailed geologic map, of the edifice, both of which focus on 

the youngest part (Spirit Lake stage) of the volcano and provide little information on the older rocks.

New geologic mapping and radiometric dating allow us 

to refine and extend the general geochronologic framework—

especially in the older part of the record that predates construction of the modern volcanic edifice. Here we report new 

radiometric (primarily 40Ar/39Ar) ages of Ape Canyon- and 

Cougar-stage lavas and pyroclastic deposits and describe the 

older rocks erupted from Mount St. Helens and preserved in 

its debris apron and in river valleys draining the volcano. We 

Stage and age Period and age Tephra set

Spirit Lake

stage,

3.9–0 ka

Modern period,

1980–present 1980

Goat Rocks period,

1800–1857 C.E. layer T

Kalama period,

1479–1750 C.E.

set X

set W

Sugar Bowl period,

1,200–1,150 yr B.P. layer D

Castle Creek period,

2,200–1,895 yr B.P. set B

Pine Creek period,

3,000–2,500 yr B.P. set P

Smith Creek period,

3,900–3,300 yr B.P. set Y

Dormant interval, 12.8–3.9 ka

Swift Creek stage, 16–12.8 ka set J

set S

Dormant interval, 18–16 ka

Cougar stage, 28–18 ka set K

set M

Dormant interval, 35–28 ka

Ape Canyon stage, 300–35 ka set C

Figure 1. Mount St. Helens eruptive history.

integrate the tephrochronology of Mullineaux (1996) and the 

age information on lahars and pyroclastic flows reported by 

Crandell (1987) into a stratigraphy of the fragmental apron of 

the volcano. Although we modify the age ranges and intervals 

between eruptive stages at Mount St. Helens, we retain Mullineaux and Crandell’s terminology. The single major modification to their scheme is that we assign volcanism substantially earlier than about 40 ka and not recognized by them to 

their Ape Canyon stage, which thus lasted from about 300 to 

about 35 ka. In all, ranges of the four stages are Ape Canyon, 

300–35 ka; Cougar, 28–18 ka; Swift Creek, 16–12.8 ka; and 

Spirit Lake, 3.9–0 ka, as summarized in figure 1.

The eruptive products of Mount St. Helens are dominated 

by varied hypersthene-hornblende dacite but include olivine 

basalt, olivine and pyroxene andesite, and quartz- and biotitebearing dacite. Since the 1980 eruption, a tremendous volume 

of geochemical data has been generated. Most analyses are 

of Spirit Lake-stage rocks, obtained to determine the general 

chemistry of the youngest part of the volcano; some analyses 

were applied to determining the origin of Mount St. Helens 

magmas, but most analyses were not carefully correlated with 

stratigraphy. Unlike their well-studied younger counterparts, 

the products of the earlier eruptive stages are inadequately 

characterized. Here we combine major-element geochemical 

data from the literature with data generated by USGS laboratories and correlate them with the petrography of the rocks to 

characterize the Ape Canyon, Cougar, and Swift Creek stages 

of Mount St. Helens.

Analytical Techniques

K-Ar and 40Ar/39Ar Geochronology

Using 40Ar/39Ar incremental-heating techniques, we 

analyzed whole rock, groundmass, or plagioclase phenocrysts 

from 17 lava samples (see fig. 3 for locations). For whole rock 

experiments, samples were crushed, ultrasonicated, and sized 

to narrow ranges (typically 500–1,000 μm). For groundmass 

separates, samples were crushed, ultrasonicated, and sized 

to narrow ranges (typically 250–350 μm), depending on the 

distance between phenocrysts. Dense, clean groundmass was 

concentrated by using a Frantz magnetic separator and careful 

handpicking under a binocular microscope. For plagioclase 

separates, samples were crushed and sieved to narrow size 

ranges, depending on the sizes of inclusion-free plagioclase 

(typically 300–400 μm), and purified by using magnetic 

and density techniques, dilute HF etching, and handpicking. 

Samples were irradiated in six different packages between 

March 2000 and May 2005, all with identical technique. For 

irradiation, samples weighing 150–250 mg were packaged in 

Cu foil and placed in cylindrical quartz vials, together with 

27.87-Ma Taylor Creek sanidine fluence monitors and K-glass 

and fluorite to measure interfering K and Ca isotopes; quartz 

vials were wrapped in 0.5-mm-thick Cd foil to shield samples 
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 595

Figure 2. Mount St. Helens area, southwestern Washington, before the 1980 eruption. Dashed outline 

denotes approximate limit of Mount St. Helens debris apron. Fragmental deposits derived from Mount St. 

Helens extend farther down all major drainages heading on volcano. Swift Reservoir fills the Lewis River 

valley between Swift Creek and Pine Creek, and Yale Lake fills valley west of Rain Creek. Base is pre-1980 

topographic quadrangle map (U.S. Geological Survey, 1919) in North American datum (later renamed North 

American Datum 1927), enhanced by D.W. Ramsey to create hillslope shading as if illuminated from northwest.

Butte Camp

 dome

Crescent

 Ridge

 fan

Swift

Creek

 fan Swift Creek

South Fork Toutle River

Castle Cr

Lahar

Ape Canyon

Cedar Flats

 fan

Swift Reservoir

Yale

Lake

Rain

Creek

West Fork

Swift

Creek

 flow

Crescent

 Ridge

 dome

Goat

Mtn

Muddy

River

June

Lake

Sheep Camp

 Worm

Complex

 flows

 Timberline

Campground

Summit

 Dome

Goat

Rocks

Floating

 Island

Sugar

Bowl

Cave Basalt

Pine

Creek

Smith

Creek

Lewis River

Lewis River cliffs

Studebaker Cr

North Fork Toutle River

Kalama River

122°15' 122°10'

46°15'

46°10'

46°05'

122°05'

0 5 KILOMETERS

0 2 MILES
596 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

from thermal neutrons during irradiation. Samples were irradiated for 2 hours in the central thimble of the TRIGA reactor at 

the USGS laboratory in Denver, Colo. (Dalrymple and others, 

1981); reactor vessel was rotated continuously during irradiation to avoid lateral neutron-flux gradients. The reactor constants determined for these irradiations were indistinguishable 

from those in other recent irradiations, and a weighted mean 

of constants obtained over the past 5 years yields 40Ar/39ArK = 

0.000±0.0004, 39Ar/37ArCa = 0.000706±0.000051, and 36Ar/37ArCa

= 0.000281±0.000009. Sanidine TCR-2 from the Taylor Creek 

Rhyolite (Duffield and Dalrymple, 1990) is a secondary standard 

calibrated against the primary intralaboratory standard, SB-3, 

which has an age of 162.9±0.9 Ma (Lanphere and Dalrymple, 

2000). Fluence monitors were analyzed using a continuous laser 

system and a MAP 216 mass spectrometer, as described by 

Dalrymple (1989). Argon was extracted from groundmass and 

plagioclase separates using a Mo crucible in a custom resistance 

furnace modified from the design of Staudacher and others 

(1978) attached to the aforementioned mass spectrometer. Heating temperatures were monitored with an optical-fiber thermometer and controlled with an Accufiber model 10 controller. Gas 

was purified continuously during extraction using two SAES 

ST-172 getters operated at 2.5 and 4 A.

Mass-spectrometric discrimination and system blanks are 

important factors in the precision and accuracy of the 40Ar/39Ar 

age determinations of Pleistocene lavas because of low radiogenic yields. Discrimination was monitored by analyzing splits 

of atmospheric Ar from a reservoir attached to the extraction 

line. Typical system blanks, including mass-spectrometer 

backgrounds, were 1.5×10-18 mol of m/z 36, 9×10-17 mol of 

m/z 37, 3×10-18 mol of m/z 39, and 1.5×10-16 mol of m/z 40, 

where m/z is the mass/charge ratio.

WASHINGTON

AREA OF MAP

Miles

Kilometers

0 2 4 6

0 2 4 6 8

122°30' 122°15' 122°00'

46°00'

46°15'

Toutle

River

Toutle

River

Fork

South

North

Green

River

Fork

Spirit

Lake Bean

Creek

Clearwater

Creek

Pine

Creek

Smith

Creek

Muddy

River

River

Lewis

Swift

Reservoir

Swift

Creek

Yale

Lake

Clear

Creek

Coweeman River

Kalama

River

Merrill

McBride

Lake

Lake

Lake

Cape Horn Cr.

Speelyai

Creek

Lewis

River

Merwin

Studebaker

Cr.

Castle

Cr.

Fork

Cr.

Rain

West

Bear

Creek

Coldwater

Creek

Hoffstadt

Creek

Cougar

Mount

St. Helens

SC99-192

SC98-155

SC99-343

SC03-611D

SC03-612

SC03-

619A

P90-22A

SC03-620A

SC98-251A

SC98-193

SC97-143

SC03-642

99YC-P32

98LCQ219B, H2

W97-141

SC03-517A

 JEO

09/15/05-3/1

Kid

Valley

Yale

Green Mountain Dam

SC97-109

Swift Dam

Merwin

 Dam

SC98-166B

Figure 3. Sketch map of Mount St. Helens area, showing locations of newly dated samples reported in 

this chapter. Rectangle, area of figure 2.
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 597

Commonly accepted criteria (McDougall and Harrison, 

1999) for a meaningful incremental-heating age are (1) a welldefined plateau (horizontal age spectrum with no significant 

slope) comprising more than 50 percent of 39Ar released, (2) a 

well-defined isochron for plateau gas fractions, (3) concordant 

plateau and isochron ages, and (4) an 40Ar/36Ar isochron intercept not significantly different from 295.5. Because several 

of the samples violate one or more of these criteria, for those 

samples we favor the isochron age or report a plateau comprising less than 50 percent of 39Ar released.

For isochron plots, data are not corrected by using an 

atmospheric ratio. Reported isochron ages include plateau 

steps on well-behaved samples or a data subset that includes 

the most steps yielding a reasonable goodness of fit. We 

commonly exclude the highest and lowest temperature steps 

because they are most strongly affected by 39Ar recoil. We 

show normal isochron plots for these low-radiogenic-yield 

rocks because the data are easier to visualize. Isochron ages 

with a high probability-of-fit regression (a low mean square of 

weighted deviates, [MSWD ~1]; York, 1969) and an 40Ar/36Ar 

intercept not within the error of the present-day air ratio are 

interpreted to contain nonatmospheric initial Ar. For these 

samples, we interpret the isochron age as most meaningful. 

Isochron ages with an MSWD value greater than the critical 

value defined by Mahon (1996) are here reported with errors 

expanded by MSWD (Ludwig, 1999); errors are reported at 

the 1σ level.

The K-Ar ages for samples SC97-141 and SC97-143 

were determined in Menlo Park, Calif., using techniques delineated in Hildreth and Lanphere (1994).

Radiocarbon Dating

The samples with a laboratory number prefixed by 

“WW” were prepared at the USGS laboratory in Reston, 

Va., by John McGeehin and analyzed at the Center for 

Accelerator Mass Spectrometry of Lawrence Livermore 

National Laboratory in Livermore, Calif. The samples 

were pretreated with an acid-alkali-acid leaching process 

to remove contaminant C in the form of inorganic carbonates and organic soil acids before conversion to CO2

 and 

reduction with H2

 to pure C in the form of graphite over 

an Fe catalyst. Additional details of sample processing and 

analysis were reported by McGeehin and others (2001) and 

Roberts and others (1997), respectively.

All 14C ages were calculated using the Libby 5,568-year 

half-life with δ13C = -25 and given in years before present 

(B.P.), defined as 1950 C.E. Errors as stated include one 

standard deviation (1σ) of counting statistics and an additional uncertainty attributable to isotope fractionation and 

sample processing. Radiocarbon ages of 0–22,000 yr B.P. 

were calibrated to calendar years by using the INTCAL04 

terrestrial age calibration (Reimer and others, 2004); older 

14C ages were calibrated by using the method of Cutler and 

others (2004).

Geochemistry

Analytical methods for USGS major-element analyses 

by wavelength-dispersive X-ray spectroscopy were routine, 

as described by Taggart and others (1987). All analyses 

were recalculated to 100 percent anhydrous, with the Fe2O3/

FeO ratio set to 0.2 of the total Fe analyzed as Fe2O3. Analyses from the literature were recalculated by using the same 

scheme. A few analyses, however, lack either P2O5 or MnO 

content, and so those two components were estimated on the 

basis of analyses of similar rocks. John Pallister, Cynthia 

Gardner, and Rick Hoblitt provided additional USGS analyses.

Results

K-Ar and 40Ar/39Ar Analyses

The results of K-Ar and 40Ar/39Ar analyses of Mount 

St. Helens rocks are summarized in table 1, and the data for 

incremental-heating experiments are listed in appendix 1. 

Plateau and isochron ages are plotted in figure 4. The dated 

samples are briefly described below, from youngest to oldest 

in the same order as in table 1. Sample locations (fig. 3) are 

referable to North American Datum 1927 (NAD 27).

Sample SC98-155.––Dacite of Kalama Dome (Evarts and 

Ashley, 1990a), a small dome of porphyritic hornblende dacite 

with sparse resorbed biotite overlying till of Hayden Creek age 

in the Kalama River valley (lat 46o07.79' N., long 122o19.76' 

W.). Generally, only a single population of plagioclase is present, although a few phenocrysts are resorbed and all display 

oscillatory zoning with thick sodic rims. The groundmass 

is fresh and contains abundant microlites in clear glass. The 

whole-rock sample yielded all-negative apparent ages but 

has a well-constrained 16.3±12.9-ka isochron age comprising 94 percent of 39Ar released, with an 40Ar/36Ar intercept of 

289.2±0.9 and MSWD = 1.4. Higher-temperature steps contained little K-derived 39Ar and had older apparent ages (see 

chemical analysis in table 5).

Sample SC99-343.––Andesite of Swift Creek, a thick 

andesite lava flow of porphyritic hornblende-augite-hypersthene andesite to dacite overlying Cougar-stage pyroclastic 

flows on the south flank of the volcano (lat 46o08.60' N., long 

122o10.27' W.). Several populations of plagioclase phenocrysts 

are present, and most hornblende is converted to an aggregate 

of plagioclase, pyroxene, and Fe-Ti oxide. The groundmass is 

turbid, with tiny crystals set in cryptofelsite. The whole-rock 

sample yielded an excellent plateau age of 17.8±2.7 ka and an 

equivalent isochron age. The andesite of Swift Creek, named 

not for the stage but for its location along Swift Creek, was 

emplaced late in Cougar time.

Sample SC03-611D.––A clast of coarsely porphyritic 

hornblende dacite with sparse quartz from a Cougar-stage 

lahar deposit along the north shore of Lake Merwin in 

the Amboy quadrangle (lat 45o58.91' N., long 122o26.35' 
40Ar/39Ar data Plateau age Isochron age Sample Material Total-gas age (ka) %39Ar (steps) Age (ka) Error (1σ) Age (ka) Error (1σ) MSWD 

SC98-155 Whole rock -95.1±8.6 72 (9 of 27) -51.2 7.3 16.3 12.9 1.4 

SC99-343 Whole rock -52.4±8.7 79 (7 of 14) 17.8 2.7 18.7 4.5 0.94 

SC03-611D Groundmass 15.2±3.0 37 (4 of 14)* 21.7 1.8 21.2 2.9 1.8 

SC03-612 Plagioclase 363.8±7.3 37 (3 of 9) 24 11 84.5 13.6 2.7 

SC98-192 Whole rock 29.1±2.2 80 (6 of 14) 27.9 1.7 28.5 4.0 0.80 

SC97-109 Groundmass 7.8±6.2 88 (6 of 12) 34.9 4.6 39.5 5.4 0.41 

SC03-619A Whole rock 34.1±1.1 97 (12 of 14) 39.8 1.0 40.0 2.4 1.6 

P90-22A Plagioclase 302.7±8.4 36 (2 of 8) 54 10 SC03-620A Plagioclase 754.1±8.5 36 (3 of 9) 74 11 79.8 29.9 1.8 

SC98-251A Whole rock -361.9±12.4 83 (10 of 18) -184 10 91.5 24.2 1.3 

SC98-193 Groundmass -42±18 20 (2 of 9) 112 24 147 37 0.65 

SC98-193 Plagioclase 369.2±35.9 69 (5 of 9) 107.4 39.2 134 169 0.52 

SC97-143 Plagioclase 695±31 42 (2 of 8) 113.8 40.8 -71.4 145.4 0.61 

SC03-642 Plagioclase 530.4±8.8 52 (3 of 10) 160 8.8 166.6 46.0 2.9 

99YC-P32 Plagioclase 519±122 77 (7 of 13) 247 12 155 55 1.3 

98LC-Q219H2 Plagioclase 366±17 62 (3 of 9) 263 19

98LC-Q219B Plagioclase 404±28 85 (7 of 9) 269 13 276 19 0.81 

K-Ar data K2O Argon

Sample Material K-Ar age (ka) Weight (g) wt % No. S.D. 40Arrad (mol/g) 

SC97-143 Whole rock 109±24 25.421 1.456 2 0.005 2.277 x 10-13 

W97-141 Whole rock 296±7 26.198 1.856 2 0.011 7.891 x 10-13 

* Weighted mean age, because the plateau does not comprise the required 50 percent of 39Ar released.

Table 1. Summary of 40Ar/39Ar and K-Ar ages.[See text for discussion of preferred ages, shown in bold. Maximum ages in italics. K-Ar ages calculated by using 1976 International Union of Geological Scien(Steiger and Jäger, 1977): λβ = 4.962×10-10/yr, λε= 0.581×10-10/yr, 40K/Ktotal = 1.167×10-4 mol/mol. Errors are estimates of analytical precision at the 68-percent co
Figure 4. Argon plateau and isochron diagrams for dated samples from Mount St. Helens deposits (see fig. 3 for locations).

-10-50510

0.0 0.2 0.4 0.6 0.8 1.0

-400-300-200-100

0

100200300

Total gas age = -95.0 ± 8.6 ka

 WMPA = -51.2 ± 7.3 ka (71.9% gas released)

SC98-155 Dacite 

600 625650

675 700

725 750 775 800 825 850 875

900

925

950

975

1000

1025

1040

1055

1070

1085

1100

1125

1150

0.00.51.01.52.0

0 25 50 75 100 125 150

280300320

550

600

625

650

675

700

725

750

775

825 800

850

875

900

925

950

975

1000

1025

1040

1055

1070

1085 1100

1125

1150

1200

Air

SC98-155 Dacite Isochron

40/36i = 289.2 ± 0.9

Isochron age (600-1000oC) = 16.3 ± 12.9 ka; MSWD = 1.4

0.0 0.2 0.4 0.6 0.8 

-80-60-40-20

0

20406080

100

Total gas age = -52.4 ± 8.7 ka

WMPA = 17.8 ± 2.7 ka (78.9% gas released)

Apparent Age (ka)

SC99-343 Andesite 

700

750

800 850 900 950

1000

1

%Radiogenic

0.00.40.8

K/Ca

0 100 200 300 400 500 600 

275280285290295300305310

700

750

800

1000 950

1050

1100

1150

1200

Air

SC99-343 Andesite Isochron 

40/36i = 294.4 ± 1.4

Isochron age (750-1150°C) = 18.7 ± 4.5 ka; MSWD = 0.94

Cumulative 39Ar released 

39Ar/36Ar

40Ar/36Ar

- Apparent Age (ka) %Radiogenic K/Ca

Cumulative 39Ar released 

39Ar/36Ar

40Ar/36Ar
0.0 0.2 0.4 0.6 0.8 1.0 -100 -50

0

50

100150200

Apparent Age (ka)

SC03-611D Dacite 

550

600

650

650

730 760

790 820 8509009751050

1125

1200

-20020

%Radiogenic0.0

0.51.01.52.0

K/Ca

0 200 400 600 800 1000

240280320360

550

600

650

650

730 760

790

820

850 975 900 1050

1125

1200

Air

SC03-611D Dacite Isochron

0.0 0.2 0.4 0.6 0.8 1.0

0

500

10001500

Total gas age = 363.8 ± 7.3 ka

WMA = 24 ± 11 ka (36.5% gas released)

SC03-612 Plagioclase 

600

700 800 900 975 1050

1125

1225

1350

02040

0.00 600.050.100.150.20

0 25 50 75 100 125 150 175

300400500600

600

700 800 900 975 1050

1125

1225 1350

Air

SC03-612 Plagioclase Isochron

40/36i = 288.4 ± 1.6

Isochron age (700-1050oC) = 84.5 ± 13.6 kaMSWD = 2.7

0.0 0.2 0.4 

-100

-75-50-25

0

255075

100

Total gas age = 

WMPA = 27.9 ± 1.7 ka (79.6% 

SC98-192 Dac

600650700

750 800 850

900 

0.0 0.51.01.52.0

0 200 400 600 8

280300320340

550

600

650 700

750

800

1100

1150

1200

Air

SC98-192 Dac

40/36i = 295.5 ± 3.0

Isochron age (700-1100oCMSWD = 0.80

Total gas age = 15.2 ± 3.0 ka

WMA = 21.7 ± 1.8 ka (36.8% gas released)

40/36i = 298.7 ± 2.3

Isochron Age (790-1200oC) = 21.2 ± 2.9 kaMSWD = 1.8

39Ar/36Ar

Cumulative 39Ar released 

40Ar/36Ar

Apparent Age (ka) %Radiogenic K/Ca

39Ar/36Ar

Cumulative 39Ar released 

40Ar/36Ar

Apparent Age (ka) %Radiogenic K/Ca

39Ar/36A

Cumulative 39

40Ar/36Ar

Figure 4.—Continued.
0.0 0.2 0.4 0.6 0.8 1.0 -400 -300 -200 -100

0

100200

Total gas age = 7.8 ± 6.2 ka

WMPA = 34.9 ± 4.6 ka (87.8% gas released)

Apparent Age (ka)

SC97-109 Dacite 

550 600 650 700 750 800

925

1000

1100

1250

-5

05

1015

%Radiogenic

0.00.51.01.5

K/Ca

0 200 400 600 800

280290300310320330340350360

550

600 650

700

750

800

925

1000

1100

1250

Air

Isochron age = 39.5 ± 5.4 ka; MSWD = 0.41

SC97-109 Dacite Isochron

-100

0

100200

Total gas age = 34.1 ± 1.1 ka

WMPA = 39.8 ± 1.0 ka (96.9% gas released)

Apparent Age (ka)

SC03-619A Dacite

550

600 640 670 700 730 760 790 820

850

900

975

1050

1125

-20-1001020

%Radiogenic0.0

0.51.01.52.0

K/Ca

0 200 400 600 800 1000

280320360

550

600

640

670

700

730

760

790

820 850

900

975

1050

Air 1125

SC03-619A Dacite

Age = 40.0 ± 2.4 ka; MSWD = 1.6

0.0 0.2 0.4 

0

200400600800

100012001400

Total gas age = 

WMA = 54 ± 10 ka 

Apparent Age (ka)

P90-22A Plag

700

800

900

975 

%Radiogenic0.00

0.040.08

K/Ca

0 25 50 75 100 12

320360400 700

800

90

1

1

1350

Air

40Ar/36Ari = 393.6 ± 4.1

Isochron age = -208 ± 3

P90-22A Plagioclase I

40Ar/36Ari = 290.0 ± 0.8

40Ar/36Ari = 295.0 ± 2.2

39Ar/36Ar

Cumulative 39Ar released 

0.0 0.2 0.4 0.6 0.8 1.0

Cumulative 39Ar released Cumulative 39Ar 

40Ar/36Ar

39Ar/36Ar

40Ar/36Ar

39Ar/36A

40Ar/36Ar

Figure 4.—Continued.
-20020406080

0.0 0.2 0.4 0.6 0.8 1.0

-2000-1000

0

1000200030004000

Total gas age = 754.1 ± 8.5 ka

WMA = 74 ± 11 ka (36.0% gas released)

Apparent Age (ka)

SC03-620A Plagioclase 

600

700 800 800 975 1050 1150

1250

1350

%Radiogenic 0.00

0.040.08

K/Ca

0 25 50 75 100 125 150 175 200 225

400600

600

700

800

800

975

1050

1150

1250 1350

Air

SC03-620A Plagioclase Isochron

Isochron age = 79.8 ± 29.9 ka; MSWD = 1.8

0.0 0.2 0.4 0.6 0.8 1.0

-3000-2500-2000-1500-1000

-500

0

Total gas age = -361.9 ± 12.4 ka

WMPA = -184 ± 10 ka (83.3% gas released)

Apparent Age (ka)

SC98-251A Dacite

550

600 625 650 675 700 725 750 775 800 850

900

925

950

975

990

1000

1015

-15-10-505

%Radiogenic

0123

K/Ca

0 10 20 30 40 50 60 70 80

260280300

550

600

625

650

675

700

725

750

775

850 800

900

950 925

975

990

1000

1015

Air

SC98-251A Dacite Isochron

Isochron age (675-950oC) = 91.5 ± 24.2 kaMSWD = 1.3

0.0 0.2 0.4 

-1000

-800-600-400-200

0

200

Total gas age = -4

SC98-193 Da

550 600

650

-4-2

024

%Radiogenic K/Ca

Discontinuous plateau age 

0 10 20 30 4

280285290295300305310

550

600

6

900

1000

1250

Air

SC98-193 Dacit

Isochron Age (700-1000oC) = MSWD = 0.65

39Ar/36Ar 39Ar/36Ar

Cumulative 39Ar released Cumulative 39Ar released 

40Ar/36Ar

40Ar/36Ar40Ar/36Ari = 293.2 ± 5.1 40Ar/36Ari = 281.5 ± 1.1

Apparent Age (ka)

39Ar/36A

Cumulative 39A

40Ar/36Ar

40Ar/36Ari = 289.7 ± 1.8

Figure 4.—Continued.
0.0 0.2 0.4 0.6 0.8 1.0 -500

0

500

10001500

Age = 369.2 ± 35.9 ka

WMPA = 107.4 ± 39.2 ka (68.5% gas released)

Apparent Age (ka)

SC98-193 Plagioclase 

550

650

750 850 925 1000

1050

1150

1350

01020

%Radiogenic

0.030.040.05

K/Ca

0 5 10 15 20 25 30 35 40 45 50 280

290300310320330340350360

550 650

750

850

925

1000

1050

1150

1350

Air

Isochron Age (750-1050oC) = 134.3 ± 168.8 kaMSWD = 0.52

SC98-193 Plagioclase Isochron

0.0 0.2 0.4 0.6 0.8 1.0

0

100200300400500600700800900

1000

Total Gas Age = 695 ± 31 ka

Maximum age = 113.8 ± 40.8 ka(41.5% gas released)

Apparent Age (ka)

SC97-143 Plagioclase 

600

700

800

900 1000

1100

1200

1350

10203040

%Radiogenic

0

0.020.040.060.08

K/Ca

0 20 40 60 80 100

300350400450

600

700

800

900

1000

1100

1200

1350

Air

Isochron Age = -71.4 ± 145.4 ka; MSWD = 0.61

SC97-143 Plagioclase Isochron

0.0 0.2 0.4 0.6 

0

500

10001500200025003000

Total gas age = 530.4 ± 

WMPA = 160.0 ± 8.8 ka (51.

Apparent Age (ka)

SC03-642 Pla

625

725 825 925 1000 1

%Radiogenic0.00

0.040.08

K/Ca

0 25 50 75 100 1

300400500

600

625725

825

1075

1250

1350

Air

SC03-642 Plagioclas

Isochron Age (600-1075oC) = 16MSWD = 2.9

40Ar/36Ari = 294.1 ± 8.3

40Ar/36Ari = 322.6 ± 16.0

39Ar/36Ar

Cumulative 39Ar released 

40Ar/36Ar

39Ar/36Ar

Cumulative 39Ar released 

40Ar/36Ar

40Ar/36Ari = 288.8 ± 4.1

39Ar/36Ar

Cumulative 39Ar re

40Ar/36Ar

Figure 4.—Continued.
0 50 100 150 200

320360400

550

600

675

725

775

825

875

925

1000

1075

1200 1150

1400

Air

Isochron Age (825-1200oC steps) = 155.0 ± 55.2 kaMSWD = 1.33

99YC-P32 Plagioclase Isochron

0.0 0.2 0.4 0.6 0.8 1.0 0.0

Total Gas Age = 519.3 ± 122.4 ka

WMPA = 247.0 ± 11.9 ka (77.3% gas released)

Apparent Age (ka)

99YC-P32 Plagioclase 

775

825

875

925

1000 1075 1150 1200

1400

0510152025

%Radiogenic

0.0320.0340.036

0.040.060.080.10

0.0 0.2 0.4 0.6 0.8 1.0

0

200400600800

1000

Total Gas Age = 366 ± 17 ka

WMPA = 263 ± 19 ka (61.8% gas released)

Apparent Age (ka)

Apparent Age (ka)

98LC-Q219H2 Plagioclase 

550

625

750

850

950

1050 1150

1250

1400

-40-200204060

%Radiogenic K/Ca

0 100 200 300 400

300400500600700

550

625750

850

1050 950

1150

1250

1400

Air

98LC-Q219H2 Plagioclase Isochron

0.0 0.2 0.4 0.

0

200400600800

1000

Total Gas Age = 4

WMPA = 269 ± 13 ka (84.8% 

98LC-Q219B P

775

850

925

1000 1050 

%Radiogenic

0.0320.0360.0400.044

K/Ca

0 50 100 150 200 250 300 

300325350375400425450475500

550

625

700

775 850

925

1050

1225

1350

Air Isochron Age (850-1225oC ste

MSWD = 0.81

98LC-Q219B Plagioc

40Ar/36Ari = 317.7 ± 15.3

40Ar/36Ari = 

39Ar/36Ar

Cumulative 39Ar released 

40Ar/36Ar

39Ar/36Ar 39Ar/36Ar

Cumulative 39Ar released Cumulative 39Ar r

40Ar/36Ar

40Ar/36Ar

1,000

800

600400200

Figure 4.—Continued.
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 605

W.). The lahar deposit contains debris-avalanche dacite and, 

probably, andesite of Swift Creek, and so it must have been 

emplaced late in Cougar time. The amphibole is mostly converted to an aggregate of pyroxene, oxide, and plagioclase and 

originally may have been cummingtonite. Plagioclase displays 

various zoning patterns, but resorbed phenocrysts are relatively 

uncommon. The groundmass is mostly devitrified to a relatively coarse-grained intergrowth of plagioclase and quartz. 

The groundmass sample yielded a disturbed age spectrum 

containing only four concordant steps, with a weighted-mean 

age of 21.7±1.8 ka. A 21.2±2.9-ka isochron age comprising 

43 percent of the 39Ar released is indistinguishable from the 

weighted-mean age and has a higher-than-atmospheric intercept and a good fit. We interpret the isochron age as the most 

reliable (see chemical analysis in table 5).

Sample SC03-612.––A clast of porphyritic hornblende 

dacite with sparse quartz from a Cougar-stage lahar deposit 

along the north shore of Lake Merwin in the Amboy quadrangle (lat 45o58.93' N., long 122o26.45'W.). The lahar 

deposit contains debris-avalanche dacite and, probably, 

andesite of Swift Creek and is of late Cougar age. Plagioclase displays various zoning patterns, but resorbed phenocrysts are relatively uncommon. The groundmass is fresh 

and consists of crystallite-poor glass. A plagioclase separate 

from this rock yielded a climbing age spectrum from 27 ka 

to nearly 2 Ma. We interpret the weighted-mean age of three 

concordant low-temperature steps (24±11 ka) as a maximum 

age for the rock.

Sample SC98-192.––Dacite of June Lake, a thick, weakly 

welded lithic pyroclastic-flow deposit exposed at June Lake 

(lat 46o09.14' N., long 122o09.65' W.) on the south flank of the 

volcano. The deposit overlies mid-Tertiary arc volcanic rocks 

and is overlain by Cougar- and Swift Creek-stage pyroclasticflow deposits. The rock is a porphyritic hypersthene-hornblende dacite. Most amphibole is converted to an anhydrous 

assemblage of pyroxene, plagioclase, and Fe-Ti oxides and 

may have originally been cummingtonite. The groundmass 

contains abundant tiny microlites and is nearly completely 

devitrified to cryptocrystalline material. The whole-rock 

sample yielded an excellent 27.9±1.7-ka plateau age and an 

equivalent isochron age.

Sample SC97-109.––Dacite of McBride Lake, a coarsely 

porphyritic hornblende dacite with sparse quartz from a Cougarstage debris-avalanche or lahar deposit (loc. 55 of Crandell, 

1987) near McBride Lake, on the southwest flank of the volcano 

(lat 46o08.49' N., long 122o15.45'W.). The rock is a common 

component of the lahar and reworked lahar deposits along 

the north shore of Lake Merwin (fig. 3). Plagioclase displays 

various zoning patterns, and a small proportion of phenocrysts 

have resorbed zones. The groundmass is cryptocrystalline and 

fresh. The groundmass sample yielded an excellent 34.9±4.6-ka 

plateau age comprising 88 percent of 39Ar released. The highesttemperature steps have very low 40Ar/36Ar ratios, forcing the 40Ar/36Ar intercept to be low and the isochron age to be slightly 

older. We interpret the plateau age as the most reliable (see 

chemical analysis in table 4).

Sample SC03-619A.––Coarsely porphyritic hornblende 

dacite with sparse quartz from an Ape Canyon-stage lahar 

deposit that overlies Hayden Creek Drift and is exposed just 

above high water along the north shore of Lake Merwin in 

the Ariel quadrangle (lat 45o59.07' N., long 122o31.55' W.). 

The rock is a common component of Ape Canyon-stage lahar 

deposits in the Lewis River valley (fig. 2). The amphibole is 

mostly converted to an aggregate of pyroxene, Fe-Ti oxide, 

and plagioclase. Plagioclase displays various zoning patterns, 

and a small proportion of the phenocrysts have resorbed 

zones. The groundmass is mostly devitrified to a relatively 

coarse grained intergrowth of plagioclase and quartz. The 

groundmass sample yielded an excellent plateau age of 

39.8±1.0 ka and an equivalent isochron age (see chemical 

analysis in table 4).

Sample P90-22A.––Porphyritic quartz-biotite-hornblende 

dacite clast from a pumiceous pyroclastic-flow deposit overlying Hayden Creek Drift. John Pallister collected this sample 

from locality 16 of Crandell (1987, probably unit P1, his fig. 

7) in Smith Creek valley (lat 46o12.35' N., long 122o05.05' 

W.). At this locality, the pyroclastic-flow deposits are interbedded with set C tephras, and the abundance of hornblende and 

paucity of cummingtonite suggest that these pyroclastic-flow 

deposits may be equivalent to tephra layer Cy. Crandell and 

others (1981) reported an age older than 46.4 ka cal yr B.P. for 

this deposit. The rock is pumiceous and consists of plagioclase, abundant hornblende, biotite, quartz, and sparse cummingtonite in a glassy groundmass with very sparse microlites. 

Plagioclase displays various zoning patterns but phenocrysts 

are generally fresh and lack sieve textures. A plagioclase separate from this rock yielded a U-shaped age spectrum; no good 

fit to the isochron data is evident. We interpret the weighted 

mean of the youngest apparent ages (54±10 ka) as a maximum 

age for the rock.

Sample SC03-620A.––Coarsely porphyritic quartz-biotitehornblende dacite from a lahar deposit exposed along the north 

shore of Lake Merwin in the Ariel quadrangle (lat 45o59.16' N., 

long 122o31.22' W.). The rock is a common component of Ape 

Canyon-stage lahar and derivative sedimentary deposits in the 

Lewis River from Yale Dam to Woodland, 32 km west. Smaller 

hornblende phenocrysts are converted to an aggregate of pyroxene, Fe-Ti oxide, and plagioclase, and larger ones have opacite 

rims. Plagioclase displays various zoning patterns, but resorbed 

phenocrysts are relatively uncommon. The groundmass is fresh 

and composed of crystallite-rich glass. A plagioclase separate 

from this rock yielded a climbing age spectrum from 72 ka to 

older than 3 Ma. We interpret the weighted-mean age of three 

concordant low-temperature steps (74±11 ka) as a maximum 

age for the rock.

Sample SC98-251A.––Light-colored, porphyritic 

hornblende-hypersthene dacite pumice from the two-pumice 

pyroclastic-flow deposit exposed in the cliffs west of Swift 

Creek (lat 46o05.87' N., long 122o11.95' W.). Most amphibole 

is converted to an anhydrous assemblage of pyroxene, plagioclase, and Fe-Ti oxide. Plagioclase is the most abundant 

phenocryst phase and displays several textures but is generally 
606 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

fresh and unresorbed. The groundmass consists of fresh glass 

choked with abundant tiny crystallites. The whole-rock sample 

yielded negative ages. An isochron comprising intermediate 

steps of the incremental-heating experiment yielded a good fit 

at 91.5±24.2 ka; however, we interpret the 24,380-yr age for 

laboratory No. W-2540 (table 2) as more reliable.

Sample SC98-193.––Andesite of Butte Camp, a porphyritic hypersthene-hornblende andesite from a glaciated lava 

dome at Butte Camp, collected from talus in an unnamed 

creek on the northwest side of the dome (lat 46o10.5' N., long 

122o14.3' W.). The rock consists of porphyritic hypersthenehornblende dacite without quartz and biotite. This rock is a 

component of Hayden Creek-age glacial deposits and Ape 

Canyon-age lahars in the Lewis River and North Fork of the 

Toutle River (fig. 3). Resorbed phenocrysts of plagioclase 

and hornblende are sparse; some phenocrysts are probably 

derived from disaggregation of cumulate-textured inclusions. 

The groundmass glass is clear and fresh but crowded with 

microlites. Age experiments on plagioclase and groundmass 

separates were attempted for this rock. Plagioclase yielded a 

good plateau age of 107.4±39.2 ka comprising 69 percent of 

39Ar released; isochron age is concordant but has a large error 

because of low radiogenic yields. The groundmass experiment was disturbed but yielded a 147.1±36.5-ka isochron, 

within the error of the plagioclase result. We interpret the 

plateau age as the most reliable (see chemical analysis in 

table 4).

Sample SC97-143.––Unnamed hypersthene-hornblende 

dacite from a Cougar-stage lahar deposit (lat 46o03.78' N., 

long 122o13.48' W.) along (Forest Service) Road 90, approximately 1 km west of Crandell’s (1987) locality 45. The rock 

consists of coarsely porphyritic hypersthene-hornblende 

dacite with sparse quartz. This rock is a minor component 

of Cougar- and Ape Canyon-age lahar deposits in the Lewis 

River from Swift Dam to Woodland, 46 km west-southwest. 

Amphibole phenocrysts are fresh but have thin opacite rims; 

plagioclase phenocrysts are fresh and are present mostly in 

a single population. The groundmass consists of microlitechoked clear glass. A plagioclase separate yielded a U-shaped 

age spectrum; no good fit to the isochron data is evident. We 

interpret the weighted mean of the youngest apparent ages 

(113.8±40.8 ka) as a maximum age for the rock. This age is 

equivalent to a 109±24-ka K-Ar age obtained in a pilot study, 

which we interpret as the most reliable (see chemical analysis 

in table 4).

Sample SC03-642.––Unnamed porphyritic hypersthenehornblende dacite from a megablock in the Cougar-stage debrisavalanche deposit exposed in Swift Creek (lat 46o05.25' N., long 

122o12.40' W.). The rock consists of porphyritic hypersthenehornblende dacite with no quartz or biotite. This rock is a minor 

component of the Cougar-stage debris-avalanche deposit (see 

"Discussion" section). Plagioclase displays various zoning patterns, including a few phenocrysts with resorption zones. The 

groundmass is fresh and consists of crystallite-choked glass. 

A plagioclase separate from this rock yielded a climbing age 

spectrum. Three intermediate incremental-heating steps yielded 

a plateau age of 160±8.8 ka, and an isochron age comprising 

the first 71 percent of 39Ar released yielded a similar age of 

167±46 ka. We interpret the plateau age as the most reliable 

(see chemical analysis in table 4).

Sample W97-141.––Dacite of Goat Mountain (Evarts 

and Ashley, 1990b), a coarsely porphyritic hornblende-biotite 

dacite with abundant quartz from the southeast flank of Goat 

Mountain (lat 46o08.85' N., long 122o16.96' W.). The rock is a 

common component of Hayden Creek-age glacial deposits and 

Ape Canyon-age lahars in the Lewis River from Yale Dam 32 

km west to Woodland and in the Kalama River west of (Forest 

Service) Road 81 (fig. 3). Our K-Ar age of 296±7 ka (table 

1) is significantly younger than the discordant K-Ar ages on 

hornblende (3.18±0.3 Ma) and biotite (1.06±0.6 and 0.76±0.6 

Ma) reported by Engels and others (1976).

Sample 99YC-P32.––Bed of set C tephra beneath Amboy 

Drift (equivalent to Hayden Creek Drift) on the crest of 

Green Mountain, about 33 km southwest of the volcano in the 

Amboy quadrangle (lat 45o57.84' N., long 122o29.04' W.). The 

bed, which is about 2 m thick, is composed of coarse ash and 

scattered lapilli of weathered dacitic glass containing crystals 

of plagioclase, quartz, biotite, green hornblende, and cummingtonite, contains charred woodchips, and overlies a soil 

developed on Tertiary bedrock. Plagioclase yielded a plateau 

age of 247±12 ka comprising 78 percent of 39Ar released. Isochron results suggest a high 40Ar/36Ar intercept but within the 

error of air. We interpret the plateau age as the most reliable.

Sample 98LC-Q219H2.––Block of slightly vesicular, coarsely and densely porphyritic dacite from a debrisflow(?) bed in the sequence of ice-contact deposits at Cape 

Horn Creek in the Ariel quadrangle (lat 45o59.16' N., long 

122o33.48' W.). Phenocrysts of plagioclase, quartz, biotite, 

green hornblende, and cummingtonite occur in a groundmass 

of hydrated glass. Plagioclase exhibits complex zoning and 

some resorption but is fresh and free of inclusions. Some hornblende grains are zoned to cummingtonite rims. Plagioclase 

yielded a three-step plateau age of 263±19 ka comprising 62 

percent of 39Ar released but no usable isochron age.

Sample 98LC-Q219B.––Pumice-lapilli bed in a 

9-m-thick sequence of ice-contact deposits exposed in a 

roadcut on the south valley wall of Cape Horn Creek, about 

37 km southwest of the volcano in the Ariel quadrangle 

(lat 45o59.16' N., long 122o33.48' W.). The deposits consist 

almost entirely of Mount St. Helens-derived dacite debris and 

overlie Amboy Drift (Hayden Creek Drift). The sample, from 

a 0.5-m-thick bed of slightly reworked pumice lapilli, consists 

of phenocrysts of plagioclase, quartz, biotite, green-brown 

hornblende, and cummingtonite within a matrix of totally 

weathered glass. Plagioclase yielded an excellent plateau age 

of 269±13 ka comprising 85 percent of 39Ar released and a 

comparable isochron age of 276±19 ka.

Radiocarbon Ages

Absolute time control for Crandell and Mullineaux’s 

stratigraphy of the Ape Canyon and Cougar stages was based 
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 607

Laboratory 

No. Sample No. Radiometric age (

14C yr B.P.) Calibrated age

Cougar-stage deposits

W-2413 — 18,560±180 22,170 (+190/−240)

W-4531 — 19,160±250 22,620 (+520/−210

WW-4543 SC03-517A 19,670±60 23,590 (+120/−150)

— Scott BC 19,700±550 23,650 (+485/−965)

W-2540 — 20,350±350 24,380 (+510/−480)

— MS-5 22,720±1,400 28,850 ±1,400

Ape Canyon-stage deposits

WW-3481 SC98-166B 42,950±560 47,430±600

WW-5561 JEO 09/15/05-3/1 44,960±930 49,540±1,000

Table 2. Radiocarbon and calibrated ages of Ape Canyon- and Cougar-stage 

deposits.

[All values in years ± 1σ. Samples with laboratory numbers prefixed by “WW” were prepared 

at the U.S. Geological Survey laboratory in Reston, Va., and analyzed by John McGeehin at 

Lawrence Livermore National Laboratory in Livermore, Calif. by accelerator mass spectrometry. 

These samples were pretreated with standard acid-alkali-acid wash, and ages were calculated with 

a 14C half-life of 5,568 yr and δ13C = –25 (see text for sample descriptions and locations). Samples 

with laboratory numbers prefixed by “W” are from Crandell and others (1981) and were not 

pretreated (see Crandell and others, 1981, for sample descriptions and locations). Sample MS-5 is 

from Major and Scott (1988), and sample Scott BC from Scott (1989). Samples younger than 22 

ka were calibrated by using the curve of Reimer and others (2004); samples older than 22 ka were 

calibrated by using the curve of Cutler and others (2004). Calibrated ages are reported as calendar 

years before 1950 C.E. Calibration curve of Cutler and others (2004) does not give uncertainties, 

which are rounded upward from the uncalibrated age.]

on several uncalibrated ages of samples that were not pretreated before analysis (Crandell and others, 1981). Radiocarbon samples, when not pretreated to reduce contamination 

by modern organic material, commonly yield inconsistent or 

minimum ages. Furthermore, the calibrated ages, especially 

those in the range 20–50 ka, are 2–6 ka older than the uncalibrated ages. These factors lead to differences between the 

ages reported here and those reported by Crandell and others 

(1981), Crandell (1987), and Mullineaux (1996).

Brief descriptions of three new samples from the Ape 

Canyon and Cougar stages and two samples from the report 

by Crandell and others (1981) are given below and listed in 

table 2. Sample locations, which are referable to NAD 27, are 

shown in figure 3.

Sample WW-4543 (SC03-517A).––Charcoal from a 

Cougar-stage hypersthene-hornblende dacite pyroclastic-flow 

deposit (“Cougar-stage white pumice”) exposed at lake level 

on the north shore of Swift Reservoir (figs. 2, 3), overlying 

set K tephra (lat 46o04.22' N., long 122o09.67' W.), in the 

Mount Mitchell quadrangle. The sample is stratigraphically 

equivalent to laboratory No. W-2413 (table 2, Crandell and 

others, 1981) from nearby locality 46 of Crandell (1987) and 

approximately stratigraphically equivalent to laboratory No. 

W-4531 (table 2; Crandell and others, 1981) on a hypersthenehornblende dacite pyroclastic-flow deposit at locality 36 of 

Crandell (1987) in Pine Creek (figs. 2, 3).

Sample Scott BC.––Wood from sediment interbedded with Cougar-stage lahar deposits in the South Fork of 

the Toutle River (fig. 1) about 0.5 km upstream from the 

confluence with Bear Creek (approx. lat 46o14.0' N., long 

122o12.65' W.) in the Elk Mountain quadrangle. The sample 

was described by Scott (1989, p. B32) as from the Bear Creek 

section but was unnumbered.

Sample W-2540.––Charcoal from a pumiceous pyroclasticflow deposit (“two-pumice pyroclastic flow”) exposed near Rain 

Creek, west of the Swift Reservoir (loc. 46 of Crandell, 1987), 

and underlying sets M and K tephras (approx. lat 46o03.25' N., 

long 122o12.9' W.) in the Mount Mitchell quadrangle.

Sample MS-5.––Charcoal from alluvial sediment below 

a lahar deposit (Major and Scott, 1988) at the intersection of 

Washington Highway 503 and Baker Road, near Speelyai Bay, 

Lake Merwin (lat 45o59.78' N., long 122o23.95' W.) in the 

Amboy quadrangle. The deposit contains Cougar-stage dacite 

but lacks debris-avalanche dacite, and so the deposit is Cougar 

age but predates the Cougar-stage debris-avalanche deposit 

(see next section).

Sample WW-3481 (SC98-166B).––Charcoal from unit 2, 

measured section C-2, upper zone of layer Cb, in the Muddy 
608 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

River quarry (lat 46o11.06' N., long 122o03.09' W.) at the 

confluence of the Muddy River and Smith Creek (Mullineaux, 

1996, p. 21) in the Smith Creek Butte quadrangle. The sample 

is equivalent to laboratory No. W-2661, the age of which is 

37,600±1,300 14C yr B.P. (Crandell and others, 1981).

Sample WW-5561 (JEO 09/15/05-3/1).––Charcoal collected from a sandy lahar deposit along the north shore of 

Lake Merwin near Woodland Park (lat 45o59.48' N., long 

122o29.11' W.) in the Amboy quadrangle. The clast population 

of the deposit is dominated by quartz-biotite dacite pumice, 

possibly equivalent to layer Cw (Vogel, 2005). Age courtesy 

of Jim O’Connor (U.S. Geological Survey).

Other Ages

Crandell and others (1981) reported 17 uncalibrated 

ages for the Swift Creek stage. Although these ages span a 

substantial period, some are internally inconsistent, and many 

are inconsistent with the ages reported by other workers from 

places where set S tephras are preserved. Here we present the 

most consistent raw ages and calibrated equivalents for the 

Swift Creek stage from Crandell and others (1981) and other 

literature (table 3).

Discussion

Ape Canyon Stage

Before the Ape Canyon stage, the Mount St. Helens area 

was an eroded terrain of moderate relief carved into mid-Tertiary (32–23 Ma) arc volcanic rocks (Evarts and others, 1987). 

The Cascade Range in southern Washington was repeatedly 

glaciated during the middle and late Pleistocene, and glaciers 

extended from the vicinity of the volcano westward to as low 

as about 30-m altitude in the Lewis River valley as recently as 

60 ka (Grigg and Whitlock, 2002; Evarts, 2004b). Thus, the 

early record of Mount St. Helens volcanism is poorly preserved and incomplete. A few set C tephras and lava domes 

are locally preserved, but the most extensive record of early 

volcanism in the area is preserved as clasts in the Cougar-stage 

debris-avalanche deposit (see discussion below) and in glacial 

deposits and lahars in the lower Lewis River valley (Evarts, 

2004a, b, 2005) and the valleys of the North and South Forks 

of the Toutle River (Scott, 1989; Evarts, 2001). Clasts of 

similar rock types in many younger Mount St. Helens pyroclastic deposits indicate that these older rocks also underlie 

the present-day edifice. New and existing radiometric ages of 

rocks and deposits emplaced during the Ape Canyon stage are 

plotted in figure 5.

Domes in the Goat Mountain and Butte Camp Area

Recent geologic mapping has established the extent of 

lava domes and remnants of lava domes that project above 

younger deposits in the area between Goat Mountain and 

the dome at Butte Camp (fig. 2), and additional rock types 

have been recognized in the Cougar-stage debris-avalanche 

deposit (see discussion below). The quartz-biotite dacite 

of Goat Mountain contains 68–69 weight percent SiO2

 and 

yielded a K-Ar age of 296±7 ka (sample W97-141, table 1). 

The hypersthene-hornblende andesite of Butte Camp contains 

62–63 percent SiO2 and yielded an 40Ar/39Ar plateau age of 

107.4±39.2 ka (sample SC98-193, table 1). A small remnant 

of another dome east of Butte Camp is a quartz-biotite dacite 

containing 68 percent SiO2. All these rocks occur in younger 

glacial deposits and lahars in the Toutle, Kalama, and Lewis 

River valleys. Similar rock types have been recognized as 

lithic clasts in Spirit Lake-stage pyroclastic deposits (for 

example, the Sugar Bowl blast deposit).

Debris-Avalanche Dacite and Rocks Preserved 

in the Cougar-Stage Debris-Avalanche Deposit

A substantial part of the Ape Canyon-stage edifice was 

removed by the Cougar-stage debris avalanche (see discussion 

below). A quartz-biotite dacite, called debris-avalanche dacite, 

which occurs as a few dome remnants and composes about 

90 percent of the debris-avalanche deposit, is lithologically 

similar to other Ape Canyon-stage quartz-biotite dacites but 

is hydrothermally altered. Preliminary ion microprobe U-Th 

dating of zircons in debris-avalanche clasts indicates an age of 

about 71±9 ka for the dacite (J.S. Pallister, written commun., 

2004). The debris-avalanche dacite and dome remnants vary 

little in composition and contain 64–65 percent SiO2.

A few other dacite types constitute the remaining 10 percent of the Cougar-stage debris-avalanche deposit. A hornblende 

dacite containing 64 percent SiO2 yielded an 40Ar/39Ar plateau 

age of 160±8.8 ka (sample SC03-642, table 1). A sample of 

the dacite of McBride Lake from the deposit at McBride Lake 

(sample SC97-109, fig. 3; loc. 44 of Crandell, 1987) yielded an 

age of 34.9±4.6 ka (table 1), one of the youngest ages from Ape 

Canyon-stage deposits.

Deposits Associated with Hayden Creek Drift in 

the Lewis River Valley

Near Mount St. Helens, the oldest preserved volcanic 

deposits overlie till that Crandell (1987) correlated with the 

Hayden Creek Drift of the Mount Rainier region (Crandell and 

Miller, 1974) on the basis of similar weathering characteristics. Crandell concluded that activity at the Mount St. Helens 

volcanic center entirely postdated this glaciation, which he 

inferred was age equivalent to marine isotope stage 4 (~60–75 

ka)—too young, as we show below.

Equivalent deposits in the lower Lewis River valley 

mapped as Amboy Drift by Mundorff (1964) and Evarts 

(2004a, b, 2005), contain clasts of quartz- and biotite-bearing 

dacite compositionally indistinguishable from Ape Canyonstage rocks, recording the occurrence of preglacial or syn-
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 609

Laboratory 

No. Description Radiometric age(

14C yr B.P.) Calibrated age

W-3548 Wood below layer Jg, South Fork of the Toutle River 10,710±150 12,790 (+80/−180)

W-5724 Peat below layer Jg, Fargher Lake 10,980±250 12,900 (+300/−150)

USGS-2780 Peat directly above and below layer Jyn 12,020±60 13,855 (+90/−50)

W-5719 Peat below layer Jy(?), Fargher Lake 11,580±250 13,400 (+300/−200)

W-2832 Base set J, east of Lahar 11,700±90 13,570 (+90/−130)

W-2441 Base set J, Smith Creek 11,880±110 13,750 (+120/−140)

W-2655 Pre-set J lahar, Muddy River, Cedar Flats fan 11,800±90 13,730 (+160/−260)

W-2866 Post set S, pre-set J lithic pyroclastic-flow deposit, Smith 

Creek, Cedar Flats fan 11,900±190 13,775 (+200/−225)

W-2870 Post set S, pre-set J lithic pyroclastic-flow deposit, Smith 

Creek, Cedar Flats fan 11,550±230 13,400 (+250/−200)

W-2868 Post set S, pre-set J lithic pyroclastic-flow deposit, Smith 

Creek, Cedar Flats fan 12,110±110 13,970 (+100/−140)

W-3145 Pre-set J lahar, South Fork of the Toutle River, Crescent Ridge fan 12,270±90 14,130 (+190/−140)

W-3133 Peat above set S, Tower Peak quadrangle 12,120±100 13,990 (+100/−140)

Porter 1 Sediment above set S, Skykomish Valley, Wash. 13,460±70 15,970 (+110/−180)

Porter 2 Sediment above set S, Skykomish Valley, Wash. 13,560±70 16,110 (+240/−160)

Porter 3 Peat above set S, Ohop Valley, Wash. 13,600±70* 16,150 (+230/−170)

QL-1436 Peat above set S, Davis Lake, Wash. 13,800±210 16,425 (+335/−320)

W-2983 Pyroclastic-flow deposit between layers Sg and So, mouth of 

Swift Creek 13,130±350 15,525 (+540/−470)

W-3141 Pumiceous pyroclastic-flow deposit, early set S, Forest Road 

90 at Swift Creek 12,910±160 15,225 (+270/−210)

Clague 1 Estimated layer Sg age, Mono-Fish Lake paleomagnetic record 13,350±100* 15,850 (+200/−230)

WSU-2714 Shells below layer Sg, Mabton, Wash. 13,325±185 15,800 (+300/−300)

USGS-684 Shells below layer Sg, Touchet, Wash. 14,060±450 16,750 (+800/−600)

*

 Uncertainty not given and arbitrarily assumed to be 70 or 100 yr, respectively.

Table 3. Radiocarbon and calibrated ages of Swift Creek-stage deposits.

[All values in years ± 1σ. Calibrated ages are reported as calendar years before 1950 C.E. All samples were pretreated except those with a laboratory number prefixed by “W.” Most of those are from Crandell and others (1981) except samples W-5719 and W-5724, which are from Mullineaux 

(1996), and sample W-2983, which is from Hyde (1975) (see Crandell and others, 1981, for sample descriptions and locations). Samples Porter 1 

and Porter 2 are from S.C. Porter and T.W. Swanson (written commun., 2006); sample Porter 3 is from Porter and others (1983); sample USGS-684 

is from Waitt (1985); sample USGS-2780 is from Carrara and Trimble (1992); sample QL-1436 is from Davis and others (1982); sample WSU2714 is from Baker and Bunker (1985); and sample Clague 1 is from Clague and others (2003). All samples were calibrated by using the curve of 

Reimer and others (2004).]

glacial eruptions at the Mount St. Helens volcanic center. 

Furthermore, small remnants of alluvial and possible debrisflow deposits composed mostly of Mount St. Helens-derived 

quartz-biotite±cummingtonite dacite are interbedded with drift 

at scattered sites in the Lewis River valley (Evarts and others, 

2003; Evarts, 2004a, b, 2005). The largest such deposit, in the 

valley of Cape Horn Creek (fig. 3) about 37 km southwest of 

the modern volcano, is composed of well-bedded silt, sand, 

pumiceous pebble gravel, and probable debris-flow beds that 

contain prismatically jointed dacite clasts as large as 35 cm 

across. These beds, as thick as 9 m, consist almost entirely 

of quartz-biotite dacite debris and overlie till that contains 

sparse clasts of similar lithology. Several quartz-biotite dacite 

clasts from the Cape Horn Creek locality contain 65–67 

percent SiO2

. The paucity of Tertiary rock fragments indicates 

that these beds are not typical fluvial sedimentary deposits 

of the Lewis River system, and their position—nearly 460 

m above the valley floor—is much too high for them to be 

an erosional remnant of a valleywide fill. The sedimentary 

deposits are interpreted as freshly erupted dacitic debris that 
610 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Ape Canyon stage

300325 275 250 225 200 175 150 125 0255075100

Cougar, Swift Creek, 

and Spirit Lake stages

Dacite of

Goat Mountain 

Calibrated radiocarbon age

U/Th age

Thermoluminescence age

K-Ar or 40Ar/39Ar age. Open

 symbol, maximum age

Apparent

dormant

interval

Andesite of 

Butte Camp

Dacite of McBride Lake

Set C

tephra

Lahar

Dacite clasts incorporated into

Cougar-stage debris avalanche,

clasts in lahars in the Lewis River,

and ash in central Washington

 Ash and blocks

recovered from glacial deposits

Dacite clast in lahar

AGE, IN THOUSANDS OF YEARS

APPROXIMATE

STRATIGRAPHIC POSITION

Upper

Lower

EXPLANATION

Figure 5. Dated samples for the Ape Canyon stage of Mount St. Helens’ eruptive history, showing ages 

and analytical (1σ) errors. Samples without error bars have uncertainties smaller than symbol, except for 

oldest thermoluminescence age, which lacks an estimated error. Open symbols, maximum ages. Arranged in 

approximate stratigraphic order; direct stratigraphic relations do not exist or are unknown for some samples and 

deposits. See tables 1 and 2 and text for data and references. Radiocarbon ages in calibrated years before 1950 

C.E. The Ape Canyon stage lasted from nearly 300 to about 35 ka, with a long hiatus from about 250 to about 160 

ka. A second, smaller hiatus may have occurred from about 70 to about 50 ka; however, additional, as-yet-undated 

samples from glacial and lahar deposits in the Lewis River may partly or completely fill these gaps. Cougar, Swift 

Creek, and Spirit Lake stages of Mount St. Helens eruptive history fit into the small box at the upper right.

was transported across the surface of a glacier that occupied 

the Lewis River valley at that time and that accumulated in a 

small lake at the glacier margin. Plagioclase separated from a 

pumice-lapilli bed in the Cape Horn Creek section yielded an 

40Ar/39Ar plateau age of 263±19 ka (sample 98LC-Q219H2, 

table 1), and plagioclase separated from a lithic dacite block 

in a debris-flow(?) bed yielded an analytically similar plateau 

age of 269±13 ka (sample 98LC-Q219B, table 1). Although 

both these ages are somewhat younger than the conventional 

K-Ar age of 296±7 ka for the dacite of Goat Mountain, they 

all indicate that the earliest activity at the Mount St. Helens 

volcanic center was considerably earlier than inferred by Crandell (1987) and Mullineaux (1996), and that the Hayden Creek 

or Amboy Drift is, at least partly, considerably older than 

inferred by Crandell and represents a glacial advance during 

marine isotope stage 8 (~300–250 ka).

Lahar Deposits in the Lewis River Valley

The most diverse known record of Ape Canyon-stage 

volcanism is preserved in lahar and related sedimentary 

deposits of post-Hayden Creek Drift age in the Lewis River 

valley. The occurrence of various similar rock types suggests the presence of several lava domes in the vicinity of 

the volcano and Goat Mountain before construction of the 

present-day edifice. Rocks from these domes are preserved in 

till and lahar deposits around Speelyai Bay, along the shores 

of Lake Merwin (fig. 3; Evarts, 2004a, b), and as far as 15 

km farther downstream near Woodland (Evarts, 2005; Vogel, 

2005). The lahar and lahar-runout deposits are dominated by 

Ape Canyon-stage dacite and contain little Tertiary bedrock 

material. Most of these deposits consist of dense dacite lava 

clasts that are well rounded, suggesting that they originated 

as stream alluvium derived from near the volcano and were 

incorporated into and transported by lahars. Typically, a few 

rock types dominate each lahar deposit, but several of the 

deposits are monolithologic, for example, the deposit exposed 

near the Speelyai Bay boat ramp (loc. E of Major and Scott, 

1988) composed of dacite of McBride Lake. The dacite of 

McBride Lake from the deposit near McBride Lake (loc. 

44 of Crandell, 1987) yielded a plateau age of 34.9±4.6 ka 

(sample SC97-109, table 1). Three other clasts of common, 

quartz-bearing hornblende dacite containing 65–66 percent 

SiO2

 collected from Ape Canyon- and Cougar-stage lahar 

deposits in the Lewis River valley yielded a plateau age of 
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 611

39.8±1.0 ka (sample SC03-619A, table 1), a K-Ar age of 

109±24 ka (sample SC97-143, table 1), and a maximum age 

of 74±11 ka (sample SC03-620A, table 1). Some deposits 

that consist almost exclusively of well-rounded quartz-biotite 

dacite pumice clasts within a sandy matrix are interpreted as 

reworked set C tephra emplaced as lahars or reworked from 

lahar deposits, as discussed below.

Terraces of laharic gravel and dacite-bearing alluvium 

occupy several levels along the shore of Lake Merwin (fig. 3). 

These deposits, as high as 40 m above the pre-reservoir river 

level, demonstrate that during the Ape Canyon and Cougar 

stages and after the Hayden Creek-age glacier had retreated, 

the lower Lewis River valley (fig. 2) contained an extensive 

fill of volcaniclastic debris from Mount St. Helens. Alternatively, these deposits may represent remnants of lahars and till 

deposited along the margins of the valley at times when it was 

filled with glacial ice. The rock types in similar deposits in the 

North Fork of the Toutle River and lower Cowlitz River valleys (Evarts, 2001) are poorly studied.

Set C Tephra

The oldest deposits from Mount St. Helens recognized 

by Crandell and others (1981) are set C tephras, with age 

about 40 ka. A new calibrated age of 47,430±600 yr B.P. 

(laboratory No. WW-3481, table 2) for tephra layer Cb in 

the Muddy River quarry is from the same locality as laboratory No. W-2661 (uncalibrated age 37,600±1,300 yr B.P.) 

of Crandell and others. Charcoal from a monolithologic and 

probably syneruptive pumiceous sandy lahar on the north 

shore of Lake Merwin (fig. 3, loc. H of Major and Scott, 

1988) yielded a calibrated age of 49,540±1,000 yr B.P. 

(laboratory No. WW-5561, table 2). Near the volcano, set C 

deposits that overlie Hayden Creek till are all younger than 

about 50 ka because any older tephras have been removed or 

obliterated by glaciation.

Deposits of quartz- and biotite-bearing tephra mineralogically equivalent to set C of Crandell (1987) and Mullineaux (1996) are widespread southwest of the volcano, 

where they are locally more than 2 m thick (Mundorff, 1964; 

Evarts and Ashley, 1990a, b; Evarts, 2004a, b, 2005). These 

deposits, which generally overlie Hayden Creek or Amboy 

Drift, were assumed to be coeval with the proximal C tephra 

and are no older than about 50 ka. However, at a few sites in 

the Lewis River valley (fig. 2), fallout beds of set C tephra 

underlie till (Evarts, 2005). Plagioclase from a tephra bed 

beneath till on Green Mountain, about 33 km southwest of 

Mount St. Helens (fig. 3; Evarts, 2005), yielded an 40Ar/39Ar 

plateau age of 247±12 ka (sample 99YC-P32, table 1), 

confirming that much-older set C tephras are preserved 

locally. Furthermore, two other tephras in central Washington that are correlated chemically with Mount St. Helens 

yielded thermoluminescence ages of 83.2±8.3 and about 115 

ka (Busacca and others, 1992; Berger and Busacca, 1995). 

A thermoluminescence age of 46±5 ka for tephra layer Cy 

(Berger and Busacca, 1995) agrees closely with the calibrated ages for the adjacent layer Cb and the reworked set C 

pumice emplaced as lahars, as discussed above.

Set C tephras and related pyroclastic-flow deposits 

consist of coarsely porphyritic quartz-biotite-cummingtonitehornblende dacite, mineralogically similar to dense quartzbiotite dacite of the Ape Canyon stage. Fresh set C pumice 

is rare because most deposits are deeply weathered. Two 

samples of set C tephra and pumice from pyroclastic-flow 

deposits at locality 16 of Crandell (1987) have compositions 

consistent with Ape Canyon-stage dacite. They contain 65–68 

percent SiO2

, but most analyses have a loss on ignition of 

more than 5 percent and plot off the trends defined by dense 

dacites for most components, suggesting that their compositions have been affected by hydration and alkali and silica 

loss (fig. 6).

Figure 6. Silica-variation diagrams for Ape Canyon-stage 

rocks from the Mount St. Helens area (see figs. 2, 3). Samples 

are mostly U.S. Geological Survey analyses but include some 

analyses compiled from published reports (Crandell, 1987; 

Smith and Leeman, 1987; Mullineaux, 1996). A, K2

O versus SiO2. 

B, MgO versus SiO2

.

0.6

0.8

1.0

1.2

1.4

1.6

1.8

2.0

2.2

60 61 62 63 64 65 66 67 68 69 70 71 72

SiO2, IN WEIGHT PERCENT

K2O, IN WEIGHT PERCENT

A Ape Canyon stage

0.5

1.0

1.5

2.0

2.5

3.0

60 61 62 63 64 65 66 67 68 69 70 71

SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT

B

Dacite of McBride Lake

EXPLANATION

Tephra set C and pyro-

 clastic-flow deposits

Debris-avalanche dacite

Andesite of Butte Camp

Quartz-biotite dacite

Dacite of Goat Mountain
612 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Ape Canyon-Stage Petrography and Chemistry

Most Ape Canyon-stage rocks consist of coarsely 

porphyritic hornblende dacite characterized by the presence of quartz and, commonly, biotite. Phenocrysts typically 

constitute 30 to 40 volume percent of the rock. Plagioclase 

dominates the phenocryst assemblage and is commonly 0.5 

to 1 cm across; complex zoning patterns are present, but 

resorption textures are sparse. Rounded and embayed quartz 

is generally present and locally abundant; in a few rocks, 

quartz is euhedral. Biotite is sparsely present in about half 

of Ape Canyon-stage dacites, and abundant in a few. Biotite 

phenocrysts are generally small but can be as large as 5 mm 

across―for example, in the dacite of Goat Mountain. Small 

phenocrysts of hypersthene are typically present, especially 

in rocks lacking biotite, but are generally sparse. Hornblende, 

commonly 0.5 to 1 cm long, is present in most Ape Canyonstage dacites. Cummingtonite is sparsely present in a few 

samples. Complex zoning patterns in hornblende are sparse. 

Coarse-grained cumulate-textured inclusions and small 

quench-textured andesitic inclusions are rare in Ape Canyonstage dacites (except in the andesite of Butte Camp).

Representative analyses of samples of Ape Canyon-stage 

rocks are listed in table 4, and a larger dataset is plotted in figure 6. The samples range in SiO2

 content from about 62 to 70 

weight percent and, despite mineralogic differences, are compositionally similar to dacite from the other volcanic stages at 

Mount St. Helens (fig. 7). About half of the samples of Ape 

Canyon stage-rocks contain about 64–66 percent SiO2, similar 

to the dacite from later volcanic stages, especially the Cougar 

and Swift Creek stages. Their contents of all other oxides are 

also similar to those of younger rocks, including high Na2O, 

which distinguishes Mount St. Helens from other Cascade 

volcanoes. A few Ape Canyon-stage dacites have higher SiO2

contents, as much as 68–69 percent. Andesite is the primary 

composition of the andesite of Butte Camp, and some weathered set C tephras may be andesitic. In summary, although 

little is known in detail about Ape Canyon-stage rocks, they 

are clearly related to younger Mount St. Helens dacites and are 

derived from the Mount St. Helens magmatic system.

Summary of Ape Canyon-Stage Volcanism

Radiometric dating of Ape Canyon-stage deposits indicates at least two periods of volcanism at Mount St. Helens 

during Ape Canyon time—one about 300–250 ka and another 

about 160–35 ka—possibly separated by a long hiatus. Alternatively, the evidence for eruptive events between 250 and 160 

ka has been buried by younger volcanic deposits or simply not 

yet recognized. The earlier period is represented by the dacite 

dome of Goat Mountain (fig. 2), just west of the volcano, and 

by tephra and sediment associated with glacial till in the Lewis 

River valley and elsewhere (fig. 2). The later period is represented by megablocks in the Cougar-age debris-avalanche 

deposit, dacite clasts in lahar deposits in the river valleys 

0.2

0.4

0.6

0.8

1.0

1.2

1.4

1.6

1.8

2.0

2.2

48 50 52 54 56 58 60 62 64 66 68 70 72

SiO2, IN WEIGHT PERCENT

K2O, IN WEIGHT PERCENT

A

Spirit Lake stage

Goat Rocks period

Kalama period

Sugar Bowl period

Castle Creek period

Pine Creek period

Smith Creek period

Swift Creek stage

Cougar stage

Ape Canyon stage

0

1

2

3

4

5

6

7

8

48 50 52 54 56 58 60 62 64 66 68 70 72

SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT

B

EXPLANATION

Figure 7. Silica-variation diagrams for 700 Mount St. Helens 

rocks plotted by eruptive stage or, for Spirit Lake-stage rocks, 

by eruptive period. Not shown are 1980–86 or 2004–6 rock 

compositions. Samples are mostly U.S. Geological Survey 

analyses but include some analyses compiled from published 

reports (Halliday and others, 1983; Crandell, 1987; Smith and 

Leeman, 1987, 1993; Gardner and others, 1995; Mullineaux, 1996; 

Hausback, 2000). A, K2

O versus SiO2. B, MgO versus SiO2.

draining the mountain, the andesite of Butte Camp and a few 

nearby eroded remnants of unnamed lava domes, set C tephra 

near the volcano, and older tephra in central Washington. A 

second, shorter hiatus may have occurred during the later 

period between about 70 and 50 ka. Some eruptions occurred 

while the Lewis River valley was occupied by ice.

An extensive hydrothermal system existed at Mount St. 

Helens late in Ape Canyon time. Most of the debris-avalanche 

dacite in the Cougar-stage debris-avalanche deposit and some 
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 613

Sample -------- SC98-193 SC03-642 SC03-619C SC02-539A SC97-143 SC03-603A SC03-619A

Unit------------- abc hbd qbp qbp hbd qbp* hbd

Latitude N----- 46°10.5′ 46°05.25′ 45°59.07′ 45°59.19′ 46°03.78′ 46°14.05′ 45°59.07′

Longitude W -- 122°14.3′ 122°12.40′ 122°31.55′ 122°33.49′ 122°13.48′ 122°09.08′ 122°31.55′

Major-element analyses, weight percent

SiO2 62.00 63.74 64.63 65.12 65.13 65.31 65.34

Al2

O3 17.74 17.67 18.07 18.78 17.11 16.91 17.60

Fe2

O3 1.08 0.93 0.89 0.82 0.94 0.93 0.82

FeO 3.90 3.33 3.22 2.93 3.38 3.35 2.96

MgO 2.70 2.08 1.80 1.56 1.93 2.04 1.70

CaO 5.68 5.38 4.74 4.18 4.64 4.56 4.75

Na2

O 4.53 4.48 4.53 4.62 4.48 4.48 4.62

K2

O 1.29 1.39 1.22 1.25 1.40 1.50 1.37

TiO2 0.80 0.69 0.61 0.56 0.66 0.65 0.59

P2

O5 0.20 0.22 0.23 0.11 0.26 0.21 0.18

MnO 0.08 0.07 0.07 0.07 0.08 0.07 0.07

LOI 0.02 1.02 1.93 3.04 0.84 0.14 0.20

FeO*/MgO 1.80 2.02 2.24 2.35 2.19 2.06 2.18

Sample -------- SC03-620A SC97-109 SC02-484 SC02-539B P90-22A W97-141 SC03-620B

Unit------------- qbp mcb dad qbp C pf gm qbp

Latitude N----- 45°59.16′ 46°08.49′ 46°04.03′ 45°59.19′ 46°12.36′ 46°08.85′ 45°59.16′

Longitude W-- 122°31.70′ 122°15.45′ 122°11.85′ 122°33.49′ 122°05.04′ 122°16.96′ 122°31.70′

Major-element analyses, weight percent

SiO2 65.54 65.77 65.82 66.61 67.42 67.76 70.23

Al2

O3 17.24 16.91 17.12 18.20 16.73 16.78 17.00

Fe2

O3 0.86 0.93 0.86 0.85 0.78 0.66 0.49

FeO 3.09 3.34 3.11 3.05 2.80 2.39 1.75

MgO 1.79 1.91 1.76 1.35 1.40 1.27 0.67

CaO 4.62 4.44 4.55 3.54 4.03 3.92 2.73

Na2

O 4.56 4.32 4.43 4.01 4.52 4.63 4.81

K2

O 1.43 1.49 1.48 1.58 1.57 1.79 1.81

TiO2 0.59 0.65 0.60 0.60 0.53 0.48 0.34

P2

O5 0.20 0.17 0.20 0.12 0.17 0.25 0.13

MnO 0.07 0.08 0.07 0.08 0.06 0.07 0.04

LOI 0.12 0.29 2.19 2.94 2.32 0.29 1.15

FeO*/MgO 2.15 2.19 2.21 2.81 2.49 2.36 3.26

Table 4. Composition of representative Ape Canyon-stage rocks.

[All samples analyzed at the U.S. Geological Survey laboratory in Denver, Colo.; analysts, David Siems and Joseph E. Taggart, Jr. All values in weight 

percent, recalculated to 100 weight percent on an anhydrous basis, with Fe2O3 = 0.2 total Fe, analyzed as Fe2O3. LOI, loss on ignition (in percent). 

Samples are coded to stratigraphic units as follows: abc, andesite of Butte Camp; dad, debris avalanche dacite; gm, dacite of Goat Mountain; hbd, quartzbearing hornblende dacite; mcb, dacite of McBride Lake; qbp, quartz-biotite dacite; qbp*, lithic clast in the Sugar Bowl blast deposit.]
614 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

remnants of similar rocks in the area between Goat Mountain 

and Butte Camp (fig. 2) are propylitically altered. Calcite is 

abundant and pyrite sparse within the matrix; mafic minerals 

are altered and replaced by chlorite and, rarely, epidote; and 

the groundmass of these rocks is thoroughly recrystallized, 

indicating alteration by hot water. Although the timing of this 

alteration is unclear, the debris-avalanche dacite is about 70 ka 

in age, and the absence of older altered rocks suggests that the 

hydrothermal system was active in the interval after about 70 

ka and before about 30 ka.

The long period of intermittent volcanism during the Ape 

Canyon stage produced a cluster of dacite domes at an altitude 

as high as about 1,800 m slightly west of and at the present site 

of Mount St. Helens. In tephra sections, the interval between 

sets C and M (Cougar stage) is occupied by ash deposits, as 

thick as 1 m, that contain at least three weathering profiles 

(Mullineaux, 1996). Thus, the nature and duration of the 

hiatus between the Ape Canyon and Cougar stages is poorly 

constrained, but no lava flows or domes related to Mount St. 

Helens have ages corresponding to the period 35–28 ka.

Cougar Stage

The Cougar stage, between about 28 and 18 ka, was probably the most active stage in the volcano’s history before its 

30,000 28,000 26,000 24,000 22,000 20,000 18,000 16,000

Debris avalanche

Tephra sets M and K

Lahar, Speelyai Bay

Dacite of June Lake

Dacite clasts in lahar

Cougar-stage 

white pumice 

pyroclastic flow(s)

Andesite of Swift Creek

Dacite of Kalama Dome

Two-pumice pyroclastic flow

Sediment in lahar section,

 South Fork of the Toutle River

AGE, IN CALIBRATED YEARS BEFORE PRESENT

APPROXIMATE

STRATIGRAPHIC POSITION

Upper

Lower

Calibrated

14C age

Explanation

K-Ar or

40Ar/39Ar age

most recent, or Spirit Lake, stage. During the Cougar stage, 

voluminous pyroclastic flows, dacite domes, tephra, and a 

large-volume pyroxene andesite lava flow accompanied a 

debris avalanche. A few events have recently been recognized 

that preceded previous reconstructions of the Cougar stage. 

In combination with calibrated radiocarbon ages, these events 

reveal that the Cougar stage began somewhat before the 21–18 

ka range proposed by Crandell (1987) and Mullineaux (1996). 

Cougar-stage stratigraphy is summarized in figure 8, and the 

ages of Cougar-stage rocks and deposits are listed in tables 1 

and 2. These new ages and those presented in the subsection 

below entitled “Swift Creek stage” suggest that the Cougar 

and Swift Creek stages may have been separated by only a 

short hiatus; nevertheless, herein we retain the terminology 

of Crandell (1987) and Mullineaux (1996). The eruptions and 

deposits related to tephra sets M and K are designated “Cougar 

stage,” whereas those related to tephra sets S and J are designated “Swift Creek stage.”

Early Domes and Lahar Deposits

Early in the Cougar stage, an unknown number of hornblende dacite domes were erupted under the area covered by 

the present-day volcano (fig. 2). Though buried by younger 

volcanic deposits, these domes are represented by blocks in 

Figure 8. Dated samples for the Cougar stage of Mount St. Helens’ eruptive history, showing ages and 

analytical (1σ) errors. Open symbols, maximum ages. Arranged in approximate stratigraphic order; direct 

stratigraphic relations do not exist or are unknown for some of these samples and deposits. Radiocarbon ages 

in calibrated years before 1950 C.E. The Cougar stage lasted from about 28 ka to about 18 ka. See tables 1 and 

2 and text for data and references. Stratigraphic position of dacite clasts in Cougar-stage lahar is shown at 

approximate position of parent domes, which may be related to dome building that accompanied Cougar-stage 

white pumice eruptions. Lahar containing dacite clasts also contains andesite of Swift Creek and slightly 

postdates it. Same symbols and error bars as in figure 5.
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 615

lithic pyroclastic-flow deposits, blocks in the Cougarstage debris-avalanche deposit, and clasts in Fraser-age glacial 

till and lahar deposits in the Lewis River valley. Lahar deposits 

containing various hypersthene-hornblende dacite clasts occur 

in the area around the east end of Swift Reservoir and in the 

lower Lewis River valley overlying Ape Canyon-stage lahar 

deposits. Some of these lahars probably traversed the Pine 

Creek drainage from a source in the vicinity of the present-day 

volcano. Lahar deposits and their reworked equivalents that 

lack debris-avalanche dacite are extensive in the Lewis River 

valley from Swift Dam to the east end of Lake Merwin and 

probably predate the Cougar-stage debris-avalanche deposit 

(see below). Charcoal in sediment from beneath a lahar deposit 

near Speelyai Bay yielded a calibrated age of 28,850±1,400 yr 

B.P. (sample MS-5, table 2) that is interpreted as a maximum 

age for the deposit (Major and Scott, 1988). A dacite clast from 

a similar nearby lahar deposit yielded an 40Ar/39Ar isochron 

age of 21.2±2.9 ka (sample SC03-611D, table 1). Lahar 

deposits of similar age and lithology also occur in the South 

Fork of the Toutle River and Cowlitz River drainages (fig. 3; 

Crandell, 1987; Scott, 1989; Evarts and Ashley, 1990b). Wood 

from sediment interbedded with these lahar deposits yielded 

a radiocarbon age of 23,650 (+485/−965) cal yr B.P. (sample 

Scott BC, table 2; Scott, 1989).

A small dome of hypersthene-hornblende dacite containing 63.5 percent SiO2

 in the Kalama River valley 13 km 

southwest of Mount St. Helens (figs. 1, 2) was informally 

designated the “dacite of Kalama Dome” by Evarts and 

Ashley (1990a). The rock yielded an 40Ar/39Ar isochron age of 

16.3±12.9 ka (sample SC98-155, table 1), but the large uncertainty of the age and absence of stratigraphic context for this 

dome prevents interpreting its relation to other Cougar-stage 

events. The dacite of Kalama Dome probably was erupted during the Cougar stage, but a Swift Creek age is also possible. A 

small exposure of a cummingtonite-hypersthene-hornblende 

dacite lithic pyroclastic-flow deposit, called the dacite of June 

Lake, underlies Swift Creek-stage deposits near June Lake. 

The rock, which contains nearly 69 percent SiO2, yielded an 40Ar/39Ar plateau age of 27.9±1.7 ka (sample SC98-192, table 

1), indicating deposition during the Cougar stage. Fragments 

of this rock type also occur as lithic clasts in the two-pumice 

pyroclastic-flow deposit (see below). Early Cougar-stage rocks 

in the Cougar-stage debris-avalanche and lahar deposits near 

the east end of Swift Reservoir consist mostly of porphyritic 

hypersthene-hornblende dacite containing 64–65 percent SiO2

that was emplaced as domes and lithic pyroclastic flows.

Cougar-Stage Debris-Avalanche Deposit and 

Two-Pumice Pyroclastic-Flow Deposit

The most devastating Cougar-stage event was a debris 

avalanche composed primarily of Ape Canyon-stage rocks. 

First noted by Mullineaux and Crandell (1981) and described 

briefly by Newhall (1982), the debris avalanche emplaced a 

200- to 300-m-thick sheet of mostly hydrothermally altered 

debris in the drainage of ancestral Swift Creek as far south as 

the Lewis River (fig. 2). Crandell suggested that this debris 

avalanche originated in the area between Butte Camp and 

Goat Mountain. The approximately 17-km runout of the 

debris avalanche, however, suggests an origin from an altitude 

of about 2,000 to 2,200 m, estimated by using the length/

distance relation of Siebert (1984). Together with remnants 

of the debris-avalanche deposit that are buried under younger 

deposits in the Cedar Flats area, the evidence suggests a 

source in the area of the present-day edifice. The volume of 

the Cougar-stage debris-avalanche deposit was at least 1 km3

, 

probably nearly 2 km3, similar in volume to the 1980 debris 

avalanche. The Cougar-stage debris avalanche blocked the 

Lewis River at the present site of Swift Creek, and incision of 

the temporary dam generated flood-breakout lahars downstream in the Lewis River valley as far as the Columbia River 

and filled the lower Lewis River valley with lahar deposits to 

depths of at least 75 m. Filling of Yale Lake and Lake Merwin 

flooded many exposures of these lahar deposits, but they crop 

out sporadically along the shores of both reservoirs (Evarts, 

2004a, b) and are well exposed in the vicinity of the community of Cougar. These lahar deposits are interpreted to be of 

Cougar age because they contain abundant clasts of debrisavalanche dacite and pumice from the overlying two-pumice 

pyroclastic-flow deposit. Lahar deposits and material eroded 

from the Cougar-stage debris-avalanche deposit also filled the 

east end of the valley now covered by Swift Reservoir to a 

depth of about 75 m.

Emplacement of the Cougar-age debris-avalanche deposit 

was followed immediately by a large dacitic pyroclastic eruption that produced the two-pumice pyroclastic-flow deposit, 

charcoal from which has a calibrated age of 24,380 (+510/–480) 

yr B.P. (laboratory No. W-2540, table 2). The contact between 

these two deposits displays no evidence of erosion or soil formation; thus, onset of the eruption that produced the two-pumice pyroclastic flow probably initiated the Cougar-stage debris 

avalanche, which in the Swift Creek drainage is buried under as 

much as 100 m of pumice and ash.

The two-pumice pyroclastic-flow deposit consists of 

many individual flow units and has a volume of at least 

1 km3

. These flow units contain pumice of two distinct 

lithologic and compositional types: light-colored pumice 

of hypersthene-hornblende dacite containing 66–68 percent 

SiO2

, and homogeneous light- to dark-brown pumice of 

augite-hypersthene-hornblende dacite containing 63.5–64.5 

percent SiO2 (fig. 9; table 5). Blocks of banded pumice of 

both compositions are rare. The two-pumice pyroclastic-flow 

deposit can be recognized by abundant blocks of pumice, 

several meters across, as far as 12 km from the volcano. 

Dense, prismatically jointed blocks of dacite near the top of 

the two-pumice pyroclastic-flow deposit demonstrate that 

late in the eruption, dacite domes grew in the vent. Small 

amounts of the two-pumice pyroclastic-flow deposit or its 

reworked equivalent also occur in the Cedar Flats area and 

in some lahar deposits in the lower Lewis River and in the 

South Fork of the Toutle River west of Sheep Camp (fig. 3; 
616 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

loc. 71 of Crandell, 1987). The distribution of the two-pumice pyroclastic-flow deposit is similar to that of the underlying Cougar-stage debris-avalanche deposit, suggesting that 

the vent was in the area of the present-day edifice.

Tephra Sets M and K and Cougar-Stage White 

Pumice

Sets M and K tephra were erupted in quick succession 

after the two-pumice pyroclastic flow and were followed by a 

second large-volume dacitic pyroclastic eruption that produced 

the herein-named Cougar-stage white pumice. Early-emplaced 

set M tephras consist of cummingtonite-hornblende dacite and 

are overlain by hypersthene-hornblende-cummingtonite dacite 

tephra, also in set M (Mullineaux, 1996). The composition of 

set M tephra is poorly known, because only a few analyses 

of weathered pumice that yielded andesitic compositions are 

available. Set K tephra consists of hypersthene-hornblende 

dacite (Mullineaux, 1996), and a single analyzed sample contains 67 percent SiO2

.

Extensive pyroclastic-flow deposits of the Cougar-stage 

white pumice overlie sets M and K tephras and the twopumice pyroclastic-flow deposit on the south flank of Mount 

St. Helens from the Lewis River cliffs eastward to Marble 

Mountain (fig. 2). Smaller exposures occur in the Smith and 

Pine Creek drainages. Three calibrated ages for the Cougarstage white pumice (samples W-2413, W-4531, WW-4543, 

table 2) range from 22.2 to 23.6 ka. The Cougar-stage white 

pumice consists of porphyritic hypersthene-hornblende dacite, 

with subtle variations in phenocryst content and abundance. 

Six analyses fall into two groups: one containing about 63 

percent SiO2 and a second containing about 65 percent SiO2. 

The variations in lithology, composition, and age suggest that 

the Cougar-stage white pumice may represent more than one 

eruption, although no stratigraphic breaks within the unit have 

been observed.

Andesite of Swift Creek

The Cougar stage culminated with emplacement of the 

andesite of Swift Creek, a lava flow on the south flank of 

Mount St. Helens (fig. 2). The vent for the lava flow was near 

an altitude of about 1,830 m, probably at or near the summit of 

the volcano at that time. The andesite of Swift Creek consists 

of porphyritic hornblende-augite-hypersthene andesite to dacite 

containing 61–63.5 percent SiO2

 with a complex phenocryst 

assemblage. About 0.75 km3 in volume, this lava flow was 

probably the largest in the history of Mount St. Helens. Its 

varying composition and multiple populations of plagioclase 

and mafic phenocrysts indicate a magma-mixing origin.

The andesite of Swift Creek, as thick as 200 m, flowed 

nearly 6 km down Swift Creek, where it presently forms 

the divide between the West Fork and main stem of Swift 

Creek. In contrast to the interpretation by Crandell (1987), 

the lava flow appears to overlie the two-pumice pyroclasticflow deposit and probably overlies the Cougar-stage white 

pumice. Our radiometric dating of this lava flow indicates 

an 40Ar/39Ar plateau age of 17.8±2.7 ka (sample SC99-343, 

table 1). Fraser-age till (for example, locs. 39, 40, 47, and 49 

of Crandell, 1987) contains abundant clasts derived from the 

andesite of Swift Creek and so must at least partly postdate 

its emplacement.

Petrography of Cougar-Stage Rocks

Most Cougar-stage rocks are less coarsely porphyritic 

than Ape Canyon-stage rocks. Crystal content commonly 

approaches 50 volume percent, although the crystals in 

Cougar-stage rocks are generally smaller. Cougar-stage rocks 

are dominated by plagioclase, with hornblende and hypersthene as the major mafic phenocryst phases. Cummingtonite 

and augite are present in a few units, especially set M tephra, 

but quartz and biotite are generally absent, even in high-silica 

dacite. Plagioclase phenocryst populations vary, but plagioclase in most Cougar-stage dacite displays complex zoning 

0.8

1.0

1.2

1.4

1.6

1.8

2.0

60 61 62 63 64 65 66 67 68 69 70

Cougar stage

SiO2, IN WEIGHT PERCENT

K2O, IN WEIGHT PERCENT

A

1.0

1.5

2.0

2.5

3.0

60 61 62 63 64 65 66 67 68 69 70

SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT

B

Andesite of Swift Creek

Cougar-stage white pumice

Tephra sets M and K

Two-pumice pyroclastic flow

Domes and lithic

 pyroclastic flows

EXPLANATION

Figure 9. Silica-variation diagrams for Cougar-stage rocks 

from the Mount St. Helens area (figs. 2, 3). Samples are mostly 

U.S. Geological Survey analyses but include some analyses 

compiled from published reports (Crandell, 1987; Smith and 

Leeman, 1987, 1993; Evarts and Ashley, 1990a, b; Mullineaux, 

1996). A, K2

O versus SiO2. B, MgO versus SiO2.
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 617

Sample ------- SC97-114 SC97-141 SC98-251A SC98-250 SC01-440 SC02-487 SC98-155

Unit------------ asw two-pum dk two-pum dk two-pum lt white white dk

Latitude N---- 46°07.24′ 46°04.19′ 46°05.87′ 46°05.16′ 46°03.25′ 46°06.62′ 46°07.79′

Longitude W-- 122°12.17′ 122°11.84′ 122°11.95′ 122°12.25′ 122°12.88′ 122°10.69′ 122°19.76′

Major-element analyses, weight percent

SiO2 61.23 63.72 64.28 66.12 64.88 63.26 63.50

Al2

O3 17.66 17.29 17.24 17.16 17.56 17.87 17.55

Fe2

O3 1.15 1.05 1.00 0.83 0.96 1.04 0.97

FeO 4.13 3.80 3.61 2.97 3.44 3.75 3.51

MgO 2.67 2.21 2.09 1.52 1.68 2.13 2.10

CaO 5.97 4.80 4.70 4.18 4.23 4.82 4.90

Na2

O 4.50 4.47 4.50 4.57 4.67 4.53 4.63

K2

O 1.40 1.44 1.52 1.76 1.55 1.46 1.76

TiO2 0.84 0.77 0.72 0.58 0.63 0.76 0.73

P2

O5 0.36 0.35 0.26 0.24 0.33 0.31 0.27

MnO 0.09 0.09 0.08 0.07 0.07 0.08 0.08

LOI 0.08 0.96 1.17 0.95 2.29 1.85 0.06

FeO*/MgO 1.93 2.14 2.17 2.44 2.56 2.20 2.09

Sample ------- SC98-192 SC03-611D SC03-580 SC03-581 SC04-696 SC04-700C SC03-612

Unit------------ djl hbd hbd hbd hbd hbd hbd

Latitude N---- 46°09.14′ 45°58.91′ 46°06.29′ 46°06.30′ 46°02.43′ 46°02.66′ 45°58.93′

Longitude W-- 122°09.65′ 122°26.35′ 122°00.40′ 122°00.39′ 122°04.53′ 122°05.55′ 122°26.45′

Major-element analyses, weight percent

SiO2 68.92 66.06 65.01 65.06 63.70 64.41 69.24

Al2

O3 16.46 17.05 17.32 17.20 18.11 17.54 16.34

Fe2

O3 0.62 0.86 0.88 0.89 1.01 0.92 0.60

FeO 2.23 3.09 3.16 3.21 3.64 3.32 2.18

MgO 1.13 1.79 1.93 1.88 1.80 2.00 1.08

CaO 3.51 4.37 4.87 4.78 4.62 5.04 3.43

Na2

O 4.72 4.45 4.46 4.66 4.50 4.58 4.68

K2

O 1.79 1.46 1.48 1.41 1.50 1.27 1.80

TiO2 0.44 0.60 0.63 0.62 0.71 0.64 0.43

P2

O5 0.14 0.19 0.18 0.21 0.32 0.21 0.17

MnO 0.06 0.07 0.07 0.08 0.08 0.08 0.05

LOI 0.09 0.46 1.28 0.04 2.58 0.01 1.30

FeO*/MgO 2.47 2.16 2.05 2.14 2.53 2.08 2.51

Table 5. Composition of representative Cougar-stage rocks.

[All samples analyzed at the U.S. Geological Survey laboratory in Denver, Colo.; analysts, David Siems and Joseph E. Taggart, Jr. All values in 

weight percent, recalculated to 100 weight percent on an anhydrous basis, with Fe2O3 = 0.2 total Fe, analyzed as Fe2O3. LOI, loss on ignition (in 

percent). Samples are coded to stratigraphic units as follows: asw, andesite of Swift Creek; djl, dacite of June Lake; dk, dacite of Kalama Dome; 

hbd, hornblende dacite clasts in debris-avalanche and lahar deposits at the east end of Swift Reservoir and in lahar deposits in the lower Lewis 

River valley; two-pum dk or two-pum lt, dark or light-colored pumice in two-pumice pyroclastic flow; white, Cougar-stage white pumice.]
618 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

patterns; the zoning in larger crystals preserves an extended 

history of repeated crystallization and resorption, whereas 

smaller crystals are generally weakly zoned. Hornblende phenocrysts can be complexly zoned and commonly have thick 

to thin breakdown rims—features typical of those observed in 

younger Mount St. Helens dacites (Rutherford and Hill, 1993; 

Rutherford and Devine, this volume, chap. 31; Thornber and 

others, this volume, chap. 32).

Coarse-grained inclusions with cumulate or plutonic 

textures are present in Cougar-stage rocks but generally are 

smaller and less conspicuous than in Swift Creek- or Spirit 

Lake-stage rocks. Inclusions with quenched textures have been 

observed only in the dacite of June Lake and the two-pumice 

pyroclastic-flow deposit. Features indicating strong disequilibrium are generally absent in Cougar-stage dacite; however, 

Mullineaux (1996) reported the presence of olivine in separates of tephra layer Mo that suggests a basaltic component 

in the magmatic system at that time. Despite the petrographic 

differences, Cougar-stage rocks are compositionally similar 

to Ape Canyon-stage dacite (figs. 6, 9) and to younger Swift 

Creek- and Spirit Lake-stage dacites (fig. 7). The petrography and chemistry of Cougar-stage rocks indicate production 

from various magma batches of similar composition that were 

fractionated under similar conditions. Andesitic composition, 

shown by some set M tephras and the andesite of Swift Creek, 

suggest the presence of a more mafic component during 

the Cougar stage than earlier in Mount St. Helens’ history. 

The magmatic system of the Spirit Lake stage probably first 

evolved during the Cougar stage.

Summary of Cougar Stage Volcanism

The earliest known Cougar-stage eruptive events were 

lahars in the South Fork of the Toutle River, Pine Creek, and 

the Lewis River (fig. 2) and lithic pyroclastic flows on the 

south flank of Mount St. Helens, which are about 28 ka in age. 

These deposits probably originated from lava domes erupted 

in the area of the present-day edifice. Many Ape Canyon-stage 

deposits in that area were removed and redeposited in ancestral Swift and Pine Creeks by the Cougar-stage debris avalanche, which was followed immediately by the two-pumice 

pyroclastic flow (about 24.4 ka) and then by sets M and K 

tephras and the Cougar-stage white pumice (about 23.6–22.2 

ka). The debris flow and pyroclastic eruptions formed a thick 

sequence of fragmental deposits on the south flank of the volcano that blocked the Lewis River at Swift Creek. Subsequent 

failure of this blockage generated flood-breakout lahars that 

traversed the lower Lewis River. Emplacement of the andesite 

of Swift Creek about 18 ka divided the south flank into the 

West Fork of Swift Creek and the Swift Creek drainages.

At the end of the Cougar stage, Mount St. Helens consisted of a cluster of dacite domes at the site of the present-day 

edifice. No eruptions of the volcano are known between about 

18 and about 16 ka, corresponding to the demise of the Cougar 

stage and the onset of the Swift Creek stage. The hiatus is 

probably shorter than that envisioned by Crandell (1987) 

and Mullineaux (1996), because additional eruptions, unrecognized by them, occurred between the Cougar-stage white 

pumice and the basal set S tephra of the Swift Creek stage.

Swift Creek Stage

During the Swift Creek stage, the widespread sets S and 

J tephras were erupted, and three extensive fans of fragmental 

debris were emplaced on the southeast, south, west, and northwest flanks of Mount St. Helens from dacite domes. Crandell 

(1987) speculated that these fans were emplaced by disruption 

of a single dome approximately in the center of the presentday edifice. However, though broadly lithologically similar, 

the dacitic clasts that occur within the Swift Creek, Crescent 

Ridge, and Cedar Flats fans are lithologically distinct and must 

each have had separate sources.

Swift Creek Fan

The Swift Creek fan covers much of the south flank of 

Mount St. Helens in the drainage of the West Fork of Swift 

Creek (fig. 2), where it overlies Cougar-stage deposits and 

Fraser-age till. This fan includes some of the deposits described 

by Hyde (1975) as the Swift Creek assemblage. The lower part 

of the fan, which predates late set S tephra (Crandell, 1987), 

consists of pumiceous and lithic pyroclastic-flow deposits of 

cummingtonite-bearing hypersthene-hornblende dacite containing about 67 percent SiO2

. The upper part of the fan, which 

postdates set S tephra, consists of lithic pyroclastic-flow deposits of hypersthene-hornblende dacite containing 64–65 percent 

SiO2. The total thickness of the fan is about 30 m. Lahar deposits dominated by clasts of Swift Creek age cap the fan. Similar 

deposits in the small canyon just south of the old Timberline 

Campground at 1,340 m altitude on the northeast flank of the 

volcano are correlated with the Swift Creek fan. The site of the 

dome(s) that erupted the Swift Creek fan is poorly known, but it 

must have been in the vicinity of the present-day edifice. Hausback (2000) suggested that the Loowit and Archybacter domes, 

which are exposed in the walls of the 1980 eruption crater, are 

of Swift Creek age. If so, their petrography and lithology are 

most similar to Swift Creek fan rocks. The Swift Creek fan is 

smaller in volume than the Crescent Ridge or Cedar Flats fan. A 

lahar deposit exposed below Merwin Dam on the Lewis River 

(fig. 3) is probably related to the Swift Creek fan, and some 

lahar deposits in the vicinity of Cougar may be of Swift Creek 

age (Crandell, 1987).

Crescent Ridge Dome and Crescent Ridge Fan

The Crescent Ridge fan is a wedge of fragmental material that stretches from the Studebaker Creek area southward 

to Butte Camp on the northwest and west flanks of Mount St. 

Helens (fig. 2). This fan is at least 200 m thick, approximately 

centered below the Crescent Ridge dome, and consists domi-
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 619

nantly of lithic pyroclastic-flow deposits. Pumiceous pyroclastic-flow deposits are also common in the lower part of the 

fan. Lahar deposits from the Crescent Ridge fan are preserved 

locally in the South Fork of the Toutle River (for example, loc. 

72 of Crandell, 1987; Evarts and Ashley, 1990b) and in the 

Toutle River north of Silver Lake (Evarts, 2001).

Rocks of the Crescent Ridge fan are well exposed in the 

headwaters of the South Fork of the Toutle River (loc. 76 of 

Crandell, 1987) and in Studebaker Creek (fig. 2). The rocks fall 

into two lithologic groups: an early, primarily pumiceous group 

of cummingtonite-hornblende dacite (sparse small hypersthene) 

and a later, dominantly lithic group of augite-hypersthenehornblende dacite (abundant hypersthene and sparse augite). 

Lithology varies somewhat in each group, especially in the 

abundances of different size ranges of plagioclase, but the textural characteristics of plagioclase are similar. The abundance 

of relatively large plagioclase crystals derived from gabbroic 

inclusions also varies. Both groups contain 64–65 percent SiO2

, 

although a few samples contain less. The composition and 

phenocryst mineralogy of the Crescent Ridge dome are similar 

to those of the lower, pumiceous part of the fan.

Most, and possibly all, of the Crescent Ridge fan probably postdates set S tephra (Crandell, 1987). No set S tephras 

have been observed within the fan, and layer Jg tephra overlies 

the fan. Preliminary paleomagnetic data (D.E. Champion, 

written commun., 2005) suggest that the early and late parts of 

the fan differ in age, possibly by decades or centuries. The distribution and mineralogy of the Crescent Ridge pyroclastic-fan 

deposits suggest that they originated partly from the currently 

exposed part of the Crescent Ridge dome, as well as from a 

now-buried uphill extension of the Crescent Ridge dome or 

from nearby domes no longer exposed. On the west flank 

of Mount St. Helens, drainage radial to the Crescent Ridge 

dome incised the Crescent Ridge pyroclastic fan; some of that 

drainage is still active, but part was later buried by voluminous 

pyroclastic-fan deposits of Spirit Lake (Kalama) age.

Cedar Flats Fan

The Cedar Flats fan completely filled the valley of Pine 

Creek and spilled into the Lewis River valley in the vicinity 

of Cedar Flats (fig. 2) during the period between the eruption 

of sets S and J tephra. This fan is at least 100 m thick in Pine 

Creek and 50 to 100 m thick in the Cedar Flats area. Remnants 

of lahar and lithic pyroclastic-flow deposits in the Smith Creek 

and Ape Canyon drainages indicate that a small component of 

the Cedar Flats fan originated on the east flank of Mount St. 

Helens and traveled down the Muddy River. The fan is dominated by lahar deposits but also contains lithic pyroclasticflow deposits of dense hypersthene-hornblende dacite similar 

to set S tephra that were produced by hot collapses of growing 

domes which must have been situated near the present-day 

edifice. The Cedar Flats fan filled the Lewis River valley to a 

depth of at least 50 m in the area around the west end of Swift 

Reservoir and flowed an unknown distance down the Lewis 

River. Most clasts in the fan contain 64–65 percent SiO2

.

Age of Swift Creek-Stage Volcanism

The age of set S tephras (specifically the late set S layers Sg and So) is disputed. Radiocarbon ages (fig. 10; table 

3) cited by Mullineaux (1996) seem to establish the onset of 

set S volcanism at about 15.5 ka, on the basis of the age of a 

pyroclastic flow interbedded in the set. However, additional 

radiocarbon ages from localities where layers Sg and So are 

interbedded with Missoula Flood deposits in eastern Washington and Vashon Drift in the Puget Lowland are as old as 16.5 

ka (Davis and others, 1982; Porter and others, 1983; Baker and 

Bunker, 1985; Waitt, 1985; S.C. Porter and T.W. Swanson, written commun., 2006). These radiocarbon ages were determined 

on shells, peat, or sediment, and the results are neither internally 

nor stratigraphically consistent. Furthermore, in at least one 

sample, it is unclear which set S layer was being dated. The 

samples from set S pyroclastic flows were not pretreated, and 

so their ages may be slightly too young. Correlation using the 

Mono Lake-Fish Lake paleosecular-variation curve suggests an 

age for layers Sg and So of about 15.8 ka (Clague and others, 

2003), depending on the assumptions used. Consideration of all 

these ages suggests the most reliable age for layers Sg and So is 

about 16 ka. The older set S tephras are probably no more than 

a few hundred years older, because they lack weathering zones 

(Mullineaux, 1996).

The precise ages of the Swift Creek-stage pyroclastic fans 

are unknown. Accumulation of the Swift Creek fan probably 

began around 16 ka and continued for an unknown period. The 

Crescent Ridge fan at least partly overlies set S tephras and is 

overlain by set J tephras. These fans may have accumulated 

quickly; a dated lahar deposit overlying the Crescent Ridge fan 

is younger than about 14 ka. The Cedar Flats fan postdates set S 

tephra and predates set J tephra; determinations on unpretreated 

samples from within the fan indicate an age of about 14 ka if 

the age for tephra layer Jyn is accurate. The apparent hiatus 

shown in figure 10 in the period of about 15–14 ka separates 

volcanism related to the older set, S, and the younger set, J, and 

is corroborated by the weak soil zone at the top of set S tephra 

(Mullineaux, 1996). However, nonexplosive volcanism probably continued during at least part of this period and formed the 

Crescent Ridge and Cedar Flats fans.

All three fans of Swift Creek stage are overlain by set J 

tephra. Calibrated ages are slightly contradictory, probably 

owing to a lack of pretreatment among the samples processed 

before 1990. The best single age may be that of Carrara and 

Trimble (1992) for layer Jyn in northeastern Washington. 

The sample was pretreated, and its calibrated age, 13,855 

(+90/−50) yr B.P. (laboratory No. USGS-2780, table 3), suggests that set J tephra probably began accumulating at least as 

early as about 13.86 ka. Other ages suggest that the later set J 

tephras may be as young as about 12.8 ka. Additional geochronologic data are needed to refine the period represented 

by Swift Creek-stage volcanism.
620 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

18,000 17,000 16,000 15,000 14,000 13,000 12,000

Below layer Jg

Below layer Jy

Base set J

Lahar, pre-J, Crescent Ridge fan

Pre-set J, Cedar Flats fan

Lithic pfs,

post set S, pre-J

Cedar Flats fan

Peat, post-set S

Shells below

layer Sg

Paleomagnetic correlation with layer Sg

Sediment and

peat above set S

pf between layers Sg and So

pf, early set S, Swift Creek fan

Layer Jyn

Crescent Ridge and

Studebaker fans

AGE, IN CALIBRATED YEARS BEFORE PRESENT

APPROXIMATE

STRATIGRAPHIC POSITION

Upper

Lower

Figure 10. Dated samples for the Swift Creek stage of Mount St. Helens’ eruptive history, showing ages and analytical 

(1σ) errors. Arranged in approximate stratigraphic order; direct stratigraphic relations do not exist or are unknown for 

some samples and deposits. Radiocarbon ages in calibrated years before 1950 C.E. The Swift Creek stage lasted from 

about 16 ka to possibly as late as about 12.8 ka. See table 3 and text for data and references. Same symbols and error 

bars as in figure 5, except open symbol, which represents estimated age for tephra layer Sg on the basis of correlation of 

radiocarbon ages with Fish Lake and Mono Lake paleomagnetic records.

Swift Creek Stage Petrography

Set S tephras consist of cummingtonite- and hornblendebearing silicic andesite and dacite (fig. 11; table 6), and all but 

layer Sw also contain hypersthene (Mullineaux, 1986). Analyses indicate that layer So contains about 62–63 percent SiO2

and layer Sg about 64 percent SiO2. Varyingly porphyritic 

dacite, containing 63 to 67 percent SiO2 and containing combinations of hornblende, hypersthene, and augite, were erupted 

from Mount St. Helens during the period equivalent to and 

shortly after set S tephras and deposited as the fans described 

above. Though similar macroscopically, the pyroclastic fans 

are distinguishable on the basis of their relative proportions of 

phenocrysts, textural characteristics of phenocryst minerals, 

and systematic differences in bulk chemistry. Plagioclase and 

hornblende display complex zoning patterns and evidence of 

reheating, resorption, and reaction.

Early set J tephras, layers Js and Jy, consist of hypersthenehornblende silicic andesite containing 62–63 percent SiO2

; and 

later set J tephras, layers Jb and Jg, of hypersthene-hornblende 

andesite containing 58–61 percent SiO2

 (Mullineaux, 1996; 

Smith and Leeman, 1987). The youngest and most widespread 

set J tephra, layer Jg, also contains augite. No lava domes or 

flows are known to be associated with set J tephra; however, a 

dacite lava flow adjacent to, but clearly younger than, the Crescent Ridge dome may be similar in age to set J.

Some Swift Creek-stage dacite contains reacted cummingtonite or augite. Complexly sieve-textured plagioclase 

with thin overgrowth rims is common in tephra layer Jg, much 

more so than in any set S-related Swift Creek rocks described 

to date. Complexly zoned hornblende is also common. 

Coarse-grained inclusions with cumulate or plutonic textures 

are locally common in rocks of the Crescent Ridge and Cedar 

Flats fans but generally smaller and less conspicuous than in 

Spirit Lake-stage rocks. Inclusions exhibiting quench texture 

are absent in Swift Creek-stage rocks. Although no temperature data are available, the presence of cummingtonite in early 

Swift Creek-stage rocks suggests that the magmas may have 

evolved from relatively cool and wet to relatively hot and dry 

(Geschwind and Rutherford, 1992).
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 621

Figure 11. Silica-variation diagrams for Swift Creek-stage 

rocks from the Mount St. Helens area (figs. 2, 3). Samples 

are mostly U.S. Geological Survey analyses but include some 

analyses compiled from published reports (Halliday and others, 

1983; Crandell, 1987; Smith and Leeman, 1987; Mullineaux, 1996; 

Hausback, 2000). A, K2

O versus SiO2. B, MgO versus SiO2.

Summary of Swift Creek-Stage Volcanism

Volcanism was intense during the relatively short Swift 

Creek stage, from about 16.0 to about 12.8 ka. Eruptions 

during this stage were dominated by the extrusion of dacite 

domes and the generation of lithic pyroclastic flows derived 

from them. Lithic pyroclastic-flow deposits were extensively 

reworked in the Pine Creek drainage, and two widespread 

tephra sets were erupted: set S before most of the dome 

building and set J after the dome building. The composition 

and lithology of Swift Creek-stage dacite vary little, but set 

J tephras are andesitic. At the end of the Swift Creek stage, 

Mount St. Helens consisted of a cluster of dacite domes with a 

summit altitude as high as about 2,100 m. Thick sequences of 

fragmental deposits filled the drainages of the mountain from 

Studebaker Creek south around to Pine Creek. Additional fans 

were probably emplaced on the north flank but, if so, have 

been buried by younger deposits. No eruptions are known 

between about 12.8 ka and the beginning of the Spirit Lake 

stage at about 3.9 ka (Crandell, 1987; Mullineaux, 1996).

Spirit Lake Stage

The Spirit Lake stage (3.9–0 ka), which includes the 

youngest and most complex part of the eruptive history of 

Mount St. Helens, has been described in several reports, 

particularly those by Hoblitt and others (1980) and Mullineaux and Crandell (1981), and so only a brief summary is 

given here. The excellent preservation of deposits, abundance 

of radiocarbon ages (Crandell and others, 1981; Hausback 

and Swanson, 1990), tree-ring chronology (Yamaguchi, 

1983, 1985; Yamaguchi and Lawrence, 1993; Yamaguchi 

and Hoblitt, 1995; Yamaguchi and others, 1990) and paleomagnetic work (Hagstrum and others, 2002) have elucidated 

considerable detail for this period. Although volcanism was 

intermittent, with hiatuses of a few to about 600 years, at 

the scale of our knowledge of other stages the Spirit Lake 

stage appears continuous. The Spirit Lake stage is divided 

into seven eruptive periods (fig. 1): the Smith Creek (3.9–3.3 

ka), Pine Creek (3–2.5 ka), Castle Creek (2.2–1.9 ka), Sugar 

Bowl (1,200–1,150 yr B.P.), Kalama (1479–1750 C.E.), Goat 

Rocks (1800–1857 C.E.), and Modern (1980–present) (data 

and terminology of Mullineaux, 1996; slightly modified by 

Clynne and others, 2004; Clynne and others, 2005). Volcanism during the Spirit Lake stage is dominated by dacite, as 

were earlier stages, but various compositions from basalt to 

high-silica dacite have also been erupted (fig. 7). Basalt and 

basaltic andesite dominated during the Castle Creek period, 

andesite was erupted during the Pine Creek and Kalama periods, and basaltic to andesitic magmatic inclusions sparsely 

occur in some dacites of the Pine Creek, Sugar Bowl, and 

Kalama periods.

The bulk of the pre-1980 edifice of Mount St. Helens 

above about 1,800-m altitude was constructed during the Spirit 

Lake stage. During the Smith Creek period, Mount St. Helens 

was highly explosive and erupted mostly tephra and pyroclastic flows, but at least one lava dome was also emplaced, 

and large lahars swept down the North Fork of the Toutle 

River to at least the Columbia River. During the Pine Creek 

period, Mount St. Helens erupted at least three dacite domes, 

and two large fans of lithic dacite were deposited by lithic 

pyroclastic flows and lahars, one in the Pine Creek and Swift 

Creek drainages and another in the Studebaker and Castle 

Creek drainages. During the late Pine Creek period, Mount St. 

Helens erupted andesite and basaltic andesite lava flows; then, 

after a short hiatus, volcanism resumed. During the Castle 

Creek period, Mount St. Helens erupted dacite, andesite, and 

basalt as tephra, pyroclastic flows, and lava domes and flows. 

During the Pine Creek and Castle Creek periods, Mount St. 

Helens gained much of its cone shape, and by their end, it 

had attained an altitude of about 2,450 m. During the Sugar 

Bowl period, three dacitic lava domes were emplaced, and a 

small lateral blast and tephra layer were erupted. The early 

Kalama period was initiated by two large tephra eruptions and 

followed by the growth of dacite lava domes and pumiceous 

and lithic pyroclastic flows; the middle Kalama period was 

dominated by eruption of andesite lava flows; and during the 

0.6

0.8

1.0

1.2

1.4

1.6

1.8

58 59 60 61 62 63 64 65 66 67 68

Swift Creek stage

SiO2, IN WEIGHT PERCENT

K2O, IN WEIGHT PERCENT

A

1.0

1.5

2.0

2.5

3.0

3.5

4.0

58 59 60 61 62 63 64 65 66 67 68

SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT

B

Tephra set J

Cedar Flats fan

Crescent Ridge fan

Swift Creek fan

Tephra set S

EXPLANATION
622 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample ------- SC98-201 SC03-573 SC03-583 SC03-587A W95-84 W94-32

Unit------------ Jg cff cff cff cd csf

Latitude N---- 46°11.4′ 46°04.53′ 46°04.45′ 46°09.92′ 46°12.04′ 46°11.47′

Longitude W-- 122°14.1′ 122°00.25′ 122°00.07′ 122°05.51′ 122°12.73′ 122°13.85′

Major-element analyses, weight percent

SiO2 59.40 64.63 64.44 63.96 64.41 64.75

Al2

O3 19.02 17.32 17.59 17.33 17.46 17.39

Fe2

O3 1.26 0.90 0.91 0.97 0.95 0.91

FeO 4.54 3.23 3.27 3.50 3.40 3.29

MgO 3.16 2.09 1.96 2.21 1.95 1.91

CaO 5.94 4.98 4.91 5.13 4.97 4.78

Na2

O 4.29 4.54 4.64 4.55 4.58 4.65

K2

O 1.12 1.41 1.34 1.35 1.36 1.40

TiO2 0.93 0.63 0.63 0.71 0.67 0.63

P2

O5 0.25 0.19 0.22 0.22 0.18 0.20

MnO 0.10 0.07 0.08 0.08 0.08 0.08

LOI 2.73 0.20 0.01 0.15 0.09 0.14

FeO*/MgO 1.79 1.93 2.09 1.98 2.18 2.16

Sample ------- W94-72B W01-211 SC98-270A SC98-266 SC98-269 SC03-587B

Unit------------ csf swf swf swf swf S

Latitude N---- 46°12.81′ 46°07.79′ 46°05.64′ 46°07.21′ 46°07.23′ 46°09.92′

Longitude W-- 122°14.13′ 122°12.02′ 122°13.36′ 122°12.08′ 122°12.35′ 122°05.51′

Major-element analyses, weight percent

SiO2 65.02 64.46 64.69 64.97 66.99 65.76

Al2

O3 17.41 17.48 17.65 17.37 17.14 18.12

Fe2

O3 0.87 0.90 0.91 0.90 0.75 0.81

FeO 3.15 3.24 3.28 3.23 2.72 2.93

MgO 1.84 2.14 1.89 1.88 1.36 1.50

CaO 4.90 4.98 4.80 4.73 4.09 4.17

Na2

O 4.57 4.54 4.52 4.63 4.74 4.52

K2

O 1.40 1.37 1.39 1.42 1.47 1.35

TiO2 0.60 0.63 0.64 0.62 0.50 0.57

P2

O5 0.17 0.19 0.16 0.17 0.15 0.21

MnO 0.07 0.07 0.08 0.08 0.07 0.07

LOI 0.96 0.55 0.20 0.06 0.15 3.58

FeO*/MgO 2.14 1.89 2.17 2.14 2.49 2.44

Table 6. Composition of representative of Swift Creek-stage rocks.

[All samples analyzed at the U.S. Geological Survey laboratory in Denver, Colo.; analysts, David Siems and Joseph E. Taggart, 

Jr. All values in weight percent, recalculated to 100 weight percent on an anhydrous basis, with Fe2O3 = 0.2 total Fe, analyzed 

as Fe2O3. LOI, loss on ignition (in percent). Samples are coded to stratigraphic units as follows: cd, Crescent Ridge dome; cff, 

Cedar Flats fan; csf, Crescent Ridge fan; J, layer Jg pumice; S, set S pumice; swf, Swift Creek fan.]
28. The Pleistocene Eruptive History of Mount St. Helens, Washington, from 300,000 to 12,800 Years Before Present 623

late Kalama period, a large dome was emplaced at the summit 

of the volcano. Overall, during the Kalama period, lava flows 

and a summit dome added about 500 m height to the volcano 

and gave Mount St. Helens its pre-1980 form. During the Goat 

Rocks period, a tephra layer, lava flow, and a dome and small 

lithic pyroclastic fan were emplaced.

Implications of Older Rocks for the 

Mount St. Helens Magmatic System

Although only limited data are available to characterize the magmatic system of Mount St. Helens before the 

Spirit Lake stage, a few generalizations are possible. On the 

basis of the data discussed above, we infer that the general 

interpretations of Smith and Leeman (1987, 1993) that were 

derived primarily from their study of the Spirit Lake stage 

are generally applicable to the earlier stages. Dacite at Mount 

St. Helens probably originates by melting of mafic rocks in 

the lower crust (Smith and Leeman, 1987), and andesite is 

the result of mixing between dacite and basalt derived from 

the mantle (Smith and Leeman, 1993). The more detailed 

mixing model of Pallister and others (1992) for the Kalama 

period is probably applicable to other periods of the Spirit 

Lake stage. Disaggregation of the abundant coarse-grained 

inclusions in many Mount St. Helens dacites and andesites 

are a complicating factor for interpreting these rocks (Cooper 

and Reid, 2003).

Hopson and Melson (1990) recognized similarities in 

the composition of Mount St. Helens rocks that were repeated 

over time, and they introduced the concept of compositional 

cycles. They noted a tendency for eruptions early in a cycle to 

be highly explosive and the magma to contain cummingtonite 

and (or) quartz plus biotite. They inferred that the magma 

is relatively cool and water rich at the beginning of a cycle. 

As the cycle proceeds, the magma becomes hotter and less 

water rich, eruptions less explosive, and the mafic phenocrysts 

dominated by hornblende and hypersthene. The volcano tends 

to produce dacite that varies only slightly lithologically over a 

period of decades to centuries, probably indicating replenishment of the shallow magma chamber with hotter and dryer 

magma from deeper in the magmatic system. Cycles end with 

effusive eruptions of relatively mafic composition containing 

an anhydrous mineral assemblage of hypersthene, augite, and, 

sporadically, olivine. Hopson and Melson (1990) attributed 

the cycles to establishment of a compositional gradient in the 

magmatic system during repose of the volcano, eruption of 

part of the system, and reestablishment of the compositional 

gradient during the following repose.

Hopson and Melson’s compositional-cycle concept, 

which was developed before any comprehensive understanding of the eruptive history of Mount St. Helens, fails to explain 

many details of this history. These details are probably better 

explained by the occasional interaction of a fundamentally 

dacitic magmatic system with basalt from the mantle. Nevertheless, evidence suggests that volatile enrichment of the shallow magmatic system from below plays a role in the evolution 

of Mount St. Helens magmas.

Compositional cycles as envisioned by Hopson and 

Melson (1990) for the Spirit Lake stage were probably absent 

during the Ape Canyon stage. The common presence of biotite 

and quartz during the Ape Canyon stage indicates that early 

Mount St. Helens magmas were generally cooler and wetter 

than those later in its eruptive history. Successive eruptive 

units vary considerably petrographically, and equilibrium 

phenocryst textures suggest probably much less interaction 

between magma batches, possibly because recharge and eruptions were more intermittent. The volume erupted per unit 

of time was smaller and the magmatic system less integrated 

during the Ape Canyon stage, especially before about 50 ka. 

However, the detailed stratigraphy of the Ape Canyon stage 

is poorly known, and the incomplete preservation of Ape 

Canyon-stage deposits makes recognition of compositional 

cycles difficult.

By the Cougar stage, the magmatic system was beginning to resemble that of the present day. At times during the 

Cougar stage, Mount St. Helen erupted cool and wet magmas 

containing cummingtonite, although most Cougar-stage rocks 

lack cummingtonite, biotite, or quartz, are less crystalline than 

most Ape Canyon-stage rocks, and probably represent hotter, 

less evolved magma. Although the magmatic system was still 

primarily dacitic, more interaction occurred between magma 

batches, and the volcano was probably more continually active 

than during the Ape Canyon stage. Sequences of pumiceous 

pyroclastic-flow deposits as thick and widespread as those 

emplaced in Cougar time have not been erupted since. The 

Cougar stage probably culminated with eruption of the voluminous andesite of Swift Creek, a mixed-magma lava flow.

The Swift Creek stage may be the first well-developed 

magmatic cycle at Mount St. Helens. The thin weathering zone 

between the Cougar and Swift Creek stages and the return 

to cummingtonite-bearing magma suggests a short hiatus 

between the stages. After an initial explosive phase, the Swift 

Creek stage was primarily a period of dacite dome building. 

During this period, the volcano erupted dacite with decreasing chemical variation and an increasingly anhydrous mafic 

phenocryst assemblage. Complex zoning of plagioclase and 

hornblende phenocrysts and evidence of disequilibrium conditions indicate that interaction between dacitic magma batches 

was extensive. These features suggest cryptic interaction of 

the Swift Creek dacitic magma with a hotter, probably more 

mafic magma as the Swift Creek stage progressed. The Swift 

Creek stage ended with the eruption of set J tephras, which 

are sparsely porphyritic and andesitic and contain resorbed 

phenocrysts, indicating interaction with hot or mafic magma.

During the earlier stages, the Mount St. Helens magmatic 

system erupted proportionately more dacite than during the 

Spirit Lake stage. The mafic component is subtler, possibly 

because it represents a smaller proportion of the total magmatic system. Alternatively, interaction with mafic magma 

may have occurred at a deeper level during the earlier stages 
624 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

than during the Spirit Lake stage. Much additional study of 

the petrology of rocks from the Ape Canyon, Cougar, and 

Swift Creek stages is needed to further evaluate the magmatic 

system during the early history of Mount St. Helens.

Summary

The stratigraphy and radiometric dating described herein 

provide more detail on the early history of Mount St. Helens 

than was previously available, and the petrographic and bulkrock chemical data are useful for a preliminary evaluation of 

the long-term magmatic history of the volcano. Radiometric 

dating of rocks preserved in debris-avalanche, lahar, and 

glacial deposits demonstrates that Mount St. Helens has a 

much longer history than previously appreciated. Rocks as 

old as 300 ka are recognized and assigned to the Ape Canyon 

stage of Crandell (1987) and Mullineaux (1996). Many Ape 

Canyon-stage rocks are recognized by the presence of quartz 

and (or) biotite but are chemically similar to younger Mount 

St. Helens dacite.

During the Cougar and Swift Creek stages, the eruptive 

style of Mount St. Helens changed from intermittent dacitic 

activity of the Ape Canyon stage to the episodic and compositionally varying activity of the Spirit Lake stage. Significant eruptive events of the Cougar stage included dacite 

dome building, removal of part of the edifice by a debris 

avalanche, and emplacement of large-volume pyroclastic 

flows and an andesite lava flow. The Swift Creek stage was 

relatively short lived and dominated by the construction of 

dacite domes on the edifice and fans of fragmental material on the flanks of the volcano. Preliminary petrographic 

analysis of Cougar- and Swift Creek-stage rocks suggests 

that Cougar-stage rocks resemble Ape Canyon-stage rocks, 

whereas Swift Creek-stage rocks are more like those of the 

Spirit Lake stage. These characteristics indicate that the volcano’s magmatic system has evolved from relatively simple 

to more complex as the volcano matured and that interaction 

between dacitic magma batches with more mafic magma 

increased from the Ape Canyon to Spirit Lake stages. Further 

work may allow us to subdivide the history of the Ape Canyon stage and will clarify the magma processes active during 

the Ape Canyon, Cougar, and Swift Creek stages.

Acknowledgments

We thank Jim O’Connor for discussions on deposits in 

the Lewis River and age of the Swift Creek stage. Chemical 

analyses of USGS samples were expertly performed by Dave 

Siems, Joe Taggart, and Tammy Hannah. John Pallister, Cynthia Gardner, and Rick Hoblitt graciously contributed chemical 

analyses of some Spirit Lake-stage samples to our compilation. John McGeehin performed the radiocarbon analyses 

reported for the samples in table 2 with a laboratory number 

prefixed by “WW.” John Pallister provided preliminary data 

on the debris-avalanche dacite zircon age. Michelle Coombs, 

Duane Champion, and Scott Starratt provided technical 

reviews that improved the manuscript.

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Appendix 1. 40Ar/39Ar Analyses of Mount St. Helens Rocks

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

Complete incremental-heating, gas extraction, and radiometric age data from 17 experiments on 16 rock samples are 

tabulated in nine worksheets of a Microsoft Excel file. 
628 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 29

1

 Department of Geosciences, 104 Wilkinson Hall, Oregon State University, 

Corvallis, OR 97331; now at Department of Geoscience, 121 Trowbridge Hall, 

University of Iowa, Iowa City, IA 52242

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

3

 Department of Geosciences, 104 Wilkinson Hall, Oregon State University, 

Corvallis, OR 97331

Identification and Evolution of the Juvenile Component in 

2004–2005 Mount St. Helens Ash

By Michael C. Rowe1, Carl R. Thornber2, and Adam J.R. Kent3

Abstract

Petrologic studies of volcanic ash are commonly used 

to identify juvenile volcanic material and observe changes in 

the composition and style of volcanic eruptions. During the 

2004–5 eruption of Mount St. Helens, recognition of the juvenile component in ash produced by early phreatic explosions 

was complicated by the presence of a substantial proportion 

of 1980–86 lava-dome fragments and glassy tephra, in addition to older volcanic fragments possibly derived from crater 

debris. In this report, we correlate groundmass textures and 

compositions of glass, mafic phases, and feldspar from 2004–5 

ash in an attempt to identify juvenile material in early phreatic 

explosions and to distinguish among the various processes that 

generate and distribute ash. We conclude that clean glass in the 

ash is derived mostly from nonjuvenile sources and is not particularly useful for identifying the proportion of juvenile material in ash samples. High Li contents (>30 μg/g) in feldspars 

provide a useful tracer for juvenile material and suggest an 

increase in the proportion of the juvenile component between 

October 1 and October 4, 2004, before the emergence of hot 

dacite on the surface of the crater on October 11, 2004. The 

presence of Li-rich feldspar out of equilibrium (based on Liplagioclase/melt partitioning) with groundmass and bulk dacite early in the eruption also suggests vapor enrichment in the 

initially erupted dacite. If an excess vapor phase was, indeed, 

present, it may have provided a catalyst to initiate the eruption. 

Textural and compositional comparisons between dome fault 

gouge and the ash produced by rockfalls, rock avalanches, and 

vent explosions indicate that the fault gouge is a likely source 

of ash particles for both types of events. Comparison of the 

ash from vent explosions and rockfalls suggests that the fault 

gouge and new dome were initially heterogeneous, containing 

a mixture of conduit and crater debris and juvenile material, 

but became increasingly homogeneous, dominated by juvenile 

material, by early January 2005.

Introduction

The phreatic explosion at Mount St. Helens, Wash., on 

October 1, 2004, was the first of four such events that occurred 

between October 1 and 5, 2004, preceding the emergence of 

hot dacitic lava on the crater floor on October 11, 2004. From 

October 5, 2004, through December 2005, two additional 

explosions and numerous rockfalls sent volcanic ash over the 

crater rim (Scott and others, this volume, chap. 1).

Petrologic characterization of volcanic ash provides a 

means to monitor volcanic activity and to assess precursory 

evidence of changes in eruptive behavior (Taddeucci and others, 2002). Volcanic monitoring by way of ash characterization 

is commonly conducted because ash may be collected easily 

at relatively low cost. Also, because the volcanic edifice is not 

always accessible, volcanic ash may provide the only petrologic evidence for changes in eruptive behavior.

Juvenile magmatic glass has been observed in the 

tephra produced before extrusion of lava flows and domes or 

before large magmatic eruptions (Watanabe and others, 1999; 

Cashman and Hoblitt, 2004). “Juvenile,” as defined in this 

study, refers to ash and dome lava erupted hot, with textural 

and geochemical characteristics similar to the earliest dome 

material (sample SH304, collected Nov. 4, 2004) and later 

dome samples. (Dome samples dredged from the crater floor 

before sample SH304 was collected consist of a heterogeneous mixture of crater and conduit debris and minor juvenile 

material; Pallister and others, this volume, chap. 30). Monitor-
630 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ing of glass compositions in volcanic ash has also identified 

multiple magmatic components in eruptions and recorded 

compositional change over the course of a single eruption 

(for example, Pallister and others, 1992; Swanson and others, 

1994, 1995; Schiavi and others, 2006).

At Mount St. Helens, previous eruptive products complicate the ash story, making petrologic monitoring of volcanic 

ash more difficult. Lava-dome growth from 1980 to 1986 

generated a thick (~250 m) cap of dacitic lava over the preexisting conduit. In addition, crater-filling breccia and tephra 

from 1980 and earlier eruptions may underlie the 1980–86 

lava dome for 500 m or more (Friedman and others, 1981). 

Ash produced by explosions is therefore likely to contain a 

substantial proportion of “older” Mount St. Helens material. 

Analysis and comparison of ash is further complicated by 

mechanical sorting during transport, requiring consideration 

of such variables as sample location relative to the vent and 

windspeed and wind direction (for example, Sparks and others, 1997; Houghton and others, 2000).

Petrologic studies of older Mount St. Helens tephra 

provide a basis for comparing the products of the 2004 eruption (ongoing at time of writing, early 2007). Glass analyses 

of 1980–82 tephra and dome samples demonstrated trends 

of increasing crystallinity and decreasing water content, with 

lower water content likely associated with progressively lower 

volume and less intense post-May 18 explosive eruptions 

(Sarna-Wojcicki and others, 1981a; Melson, 1983). Textural 

comparisons of the May 18, 1980, blast material with that 

from precursory eruptions indicate that juvenile material 

appeared as early as 2 months before the climactic eruption 

and that the juvenile component was reflected by ash particles 

with either glassy or microcrystalline matrices, characteristic of shallow crystallization (Cashman and Hoblitt, 2004). 

Cashman and Hoblitt’s study is significant in that it identifies 

the need to consider partially crystalline material in addition 

to glassy fragments as potentially juvenile. Also, it provides a 

textural comparison for current eruptive products.

Several previous studies of trace elements, capable of 

transport within volatile phases, in tephras were undertaken 

to identify evidence for magmatic degassing associated with 

explosive events before and after May 18, 1980 (Thomas 

and others, 1982; Berlo and others, 2004). Whole-rock 210Pb 

excess and Li enrichment (max ~23 μg/g) in plagioclase feldspar in the 1980 cryptodome, followed by significantly lower 

Li contents in feldspar within the May 18 fallout tephra and 

pyroclastic-flow deposits and subsequent eruptive events, led 

Berlo and others (2004) to propose that the anomalously high 

Li contents resulted from vapor transfer from a deeper magma 

source to shallow stored/stalled magma.

The goals of this study of 2004–5 Mount St. Helens 

ash were to (1) identify and track changes over time in the 

composition and proportion of juvenile eruptive material in 

explosive events associated with the reactivation of Mount St. 

Helens; and (2) combine textural and geochemical observations to distinguish between the processes and products of ash 

generation, including a comparison of ash from rockfalls and 

vent explosions. Accomplishing these goals required a detailed 

examination of the textures and geochemical characteristics of 

erupted ash samples in comparison with dome petrology and 

geochemistry, discussed elsewhere in this volume (Cashman 

and others, this volume, chap. 19; Pallister and others, this 

volume, chap. 30; Thornber and others, this volume, chap. 32; 

Kent and others, this volume, chap. 35).

Methods

Sample Collection

After the October 1, 2004, phreatic eruption of Mount 

St. Helens, 27 ash-collection stations were established around 

the perimeter of the volcano (fig. 1). Ash samples collected 

before the placement of stations were from relatively clean flat 

surfaces (fig. 2A). Collection stations were distributed radially 

along line-of-sight to the volcano, ranging in distance from 

2.4 to 10 km from the vent (fig. 1). Each station consisted of 

a rebar-suspended double bucket (fig. 2B); the inner bucket, 

with drainage slits approximately a third up from its base, was 

suspended within the outer bucket to allow excess water to 

drain without significant loss of ash. Ash was collected at stations over periods ranging in length from less than 3 to more 

than 15 days until November 29, 2004 (fig. 3). Ultimately, 

ash collection at the established stations was a function of 

station accessibility, whereas the presence of ash depended 

on predominant wind directions during the period preceding 

A23

A20

A25

A19 A18

A12

A24 A11 JL1

A10 A9

A8 A14

A7

A2

A4 A3

A5

A27

A6

A26

A22 A21

A13

A15

A16

A17

0 2 4 KILOMETERS

46°15'

122°15'

WASHINGTON

46°7'30"

122°7'30"

AREA OF MAP

Figure 1. Mount St. Helens 

area, Washington, showing 

locations of ash-collection 

stations.
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 631

A B

C D

Inner

bucket

Drainage slits

Water and 

suspended 

ash

Figure 2. Examples of ash-collection sites. A, Clean, flat surfaces. B, At an ash collection station. Inset is a schematic drawing 

of the ash-collection buckets, showing drainage slits in upper bucket and drainage holes in base of lower bucket; bolts in lower 

bucket provide support and suspension of upper bucket. C, On snow-covered surfaces. D, In petrology spiders (see text for 

description). Inset is top view of petrology spider, with nylon netting for identification and collection of hot ballistic material on 

left and baffled compartments for ash and lithic fragments on right.

collection. After November 29, 2004, most ash samples were 

collected from discrete ash-producing events (either explosions or rockfalls and rock avalanches) on snow-covered (fig. 

2C) or otherwise-clean surfaces. Snow-covered surfaces provided easy identification of new ash and allowed for tracking 

of deposited ash to its source (fig. 2C). Additional ash samples 

were collected in August 2005 from collection devices placed 

in the crater (“petrology spiders,” fig. 2D), similar to those 

deployed for seismic and deformation monitoring (LaHusen 

and others, this volume, chap. 16).

A total of 15 samples collected between October 1, 2004, 

and March 9, 2005, were analyzed in this study (table 1), of 

which 12 are of ash spanning the period October 1, 2004, 

through March 9, 2005, two are of dome fault gouge (collected 

Nov. 4, 2004, and Feb. 22, 2005), and one is from a crater 

debris flow (collected Oct. 20, 2004) associated with collapse 

of the initial spine (table 1). (See Pallister and others this volume, chap. 30, for a detailed description of dome fault gouge 

and crater debris flow collection.)

Of the four early explosive events, only the three 

included in this study (12:02 p.m. Oct. 1, 9:43 a.m. Oct. 

4, and 9:05 a.m. Oct. 5) produced downwind ash fallout 

between October 1 and October 5, 2004 (Major and others, 

2005). Samples from the suite of Mount St. Helens 2004–5 

ash deposits (Rowe and others, 2008) were selected for our 

purposes on the basis of emplacement date, wind directions, 

and volume of material collected.

Analytical Methods

Major-element compositions of groundmass glass, melt 

inclusions (where possible), feldspar, and mafic minerals 

(amphibole, clinopyroxene, and hypersthene) were measured 
632 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample Type Date collected Eruption date

MSH04E1DZ_1 ash 10/1/2004 10/1/2004

MSH04E2A03_A1 ash 10/4/2004 10/4/2004

MSH04E3RANDLE_2 ash 10/5/2004 10/5/2004

MSH04A20_10_11 ash 10/11/2004 10/5/04–10/11/04

MSH04A09_10_12 ash 10/12/2004 10/4/04–10/12/04

MSH04A20_10_16 ash 10/16/2004 10/11/04–10/16/04

MSH04A21_10_20 ash 10/20/2004 10/15/04–10/20/04

MSH04A04_11_2 ash 11/2/2004 10/16/04–11/2/04

MSH04MR_11_4 ash 11/4/2004 11/4/2004

MSH05JP_1_14A ash 1/14/2005 1/13/2005

MSH05JV_1_19 ash 1/19/2005 1/16/2005

MSH05DRS_3_9_4 ash 3/9/2005 3/8/2005

SH303-1 gouge 11/4/2004 10/18/04

SH307-1 gouge 2/22/2005 2/12/05

SH300-1 crater debris1 10/20/2004 10/15/04

1

Only fine material collected from the crater debris flow is included in this study; includes 2004 dacite as well as

older crater-floor debris.

Table 1. Ash, gouge, and crater debris samples analyzed in this study.

[Eruption dates for crater debris and gouge samples estimated by J.S. Pallister. Further details of ash samples are 

available in Rowe and others (2008).]

for all 12 ash samples by electron microprobe analysis. To 

reduce potential sampling biases, 25 to 50 feldspar phenocrysts, ~20 mafic phenocrysts, and 10 to 20 glass specimens 

were analyzed from each sample, approximately proportional 

to their relative abundances in ash samples, as estimated visually. Backscattered electron images were taken of all the ash 

and dome gouge samples to document the heterogeneity of ash 

particles within and between samples. Trace-element contents 

in feldspar from 14 samples were measured by laser-ablation 

inductively coupled plasma mass spectrometry (LA-ICP-MS). 

In addition, major- and trace-element contents in feldspar from 

the two dome fault-gouge samples (SH303-1 and SH307-1, 

table 1) and trace-element contents in feldspar from the crater 

debris sample (sample SH300-1, table 1) were determined.

Electron microprobe analyses were conducted on a 

Cameca SX–100 instrument at Oregon State University. 

Analyzed samples were not sieved, and phenocrysts and 

glasses of all sizes were analyzed to reduce the possible bias 

created by mechanical sorting during transport (Houghton 

and others, 2000). Wherever possible, electron microprobe 

analyses of phenocryst phases were made within 15 μm of 

the grain boundary. Electron microprobe analyses of glass 

were conducted according a procedure modified from that of 

Morgan and London (1996). Na and K were counted for 60 s, 

using a 2-nA beam current, a 15-keV accelerating voltage, and 

a 10-μm beam diameter. Al, Ca, Cl, Fe, Mg, Mn, P, S, Si, and 

Ti were analyzed by using a 30-nA beam current, a 15-keV 

accelerating voltage, and 10-μm beam diameter, with count 

times ranging from 10 to 50 s. The narrow beam diameter 

(~20 μm is optimal) was required by the small size of glassy 

fragments in the ash (Morgan and London, 1996). Feldspar 

(30 nA) and mafic phases (50 nA) were measured by using 

a 15-keV accelerating voltage and 5- and 1-μm beam diameters, respectively. A rhyolite glass standard (USNM 72854 

VG–568), a feldspar standard (Labradorite USNM 115900), 

and a pyroxene standard (Kakanui augite USNM 122142) 

were analyzed before each analytical session. Glass standard 

statistics are presented in Rowe and others (2008).

Trace-element (Ba, Ce, Eu, La, Li, Nd, Pb, Pr, Sr, and 

Ti) contents in feldspar were determined by LA-ICP-MS 

analysis in the W.M. Keck Collaboratory for Plasma Spectrometry at Oregon State University, using a 193-nm ArF 

Excimer laser. Analyses were performed by using a stationary 

laser (70-μm spot size) to ablate a progressively deepening 

crater in the sample materials, requiring targeted feldspar 

phenocrysts to be larger than ~80 μm in diameter. Owing 

to the large spot size required and the fine grain size of the 

feldspar in the ash, analyses were made close to the center 

of inclusion-free grains. Each individual analysis represents 

40 s of data acquisition during ablation, with background 
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 633

JL1

A02

A03

A04

A05

A06

A07

A08

A09

A10

A11

A12

A13

A14

A15

A16

A17

A18

A19

A20

A21

A22

A23

A24

A25

A26

A27

0 7 14 21 28 35 42 49 56 63

DAYS AFTER INITIAL EXPLOSION (OCT. 1, 2004)

ASH-COLLECTION STATION

10/08/04

10/15/04

10/22/04

10/29/04

11/05/04

11/12/04

11/19/04

11/26/04

Figure 3. Collection dates for ash-collection stations (see fig. 1) 

from October 1 (day 0) through November 29 (day 60), 2004.

count rates measured for 30 s before ablation. A 4-Hz pulse 

frequency resulted in an ablation crater 15 to 20 μm deep. 

Trace-element abundances were calculated relative to the 

NIST 612 glass standard, which was analyzed under identical 

conditions throughout the analytical session. U.S. Geological Survey glass BCR-2G was also analyzed as a secondary 

standard. Counts were normalized to 29Si, also measured 

during ablation, and contents were determined according 

to the method of Kent and others (this volume, chap. 35). 

Precision of trace-element analysis is presented by Kent and 

others (this volume, chap. 35). Li contents, most relevant to 

this study, have a precision of 7 to 8 percent (1σ), although 

for one session, involving measurement of samples SH300-

1, MSH04A20_10_16, and MSH05DRS_3_9_4 (table 1), 

analytical uncertainties calculated from repeated analyses of 

standard BCR-2G may be as high as 25 percent (1σ). Overall, 

uncertainties are similar to those in the analyses by Berlo and 

others (2004) and Kent and others (this volume, chap. 35), 

who reported analytical uncertainties of ~10–15 percent.

Results

Glass and Groundmass

The groundmass in ash samples is dominated by crystalline to microcrystalline textures, with groundmass crystallization resulting in the formation of crystalline silica and feldspar 

microlites (fig. 4). Matrix textures of the juvenile dacite are 

described elsewhere in this volume (Cashman and others, this 

volume, chap. 19; Pallister and others, this volume, chap. 30). 

Late crystallization of the groundmass significantly influenced 

glass analyses. Al2O3 and SiO2 contents were used to screen 

glass analyses for microlite crystallization (fig. 4), resulting 

in 81 reliable analyses of glass, herein referred to as “clean 

glass.” Most clean-glass analyses were obtained from pumice 

fragments and glass adhering to phenocrysts. Major-element 

contents in clean glass vary widely, likely owing to localized 

effects of melt crystallization and the inherent heterogeneity of 

the ash, especially in early explosions. SiO2 contents in clean 

glass range from 52 to 79 weight percent but dominantly are 

from 72 to 78 percent (table 2). No temporal trend is apparent 

to suggest a systematic change in the melt composition over 

the course of 2004–5 sampling (fig. 5).

Mafic Minerals

Mafic minerals analyzed in the ash samples include 

hypersthene, amphibole, clinopyroxene, and olivine. Hypersthene, which is the dominant mafic mineral in the ash, varies 

widely in composition (En47.1–73.7). Hypersthene in dome 

sample SH304-2 (En53.3–65.8), as well as in older 1980s dome 

samples (En44.3–69.8), similarly varies widely in composition 

(fig. 6). Amphibole compositions in the ash, new dome dacite, 

and the 1980–86 dome dacite are essentially identical, most 

easily observed in Al2O3 and FeO* contents. Al2O3 content 

ranges from ~7 to 14 weight percent in the ash and 1980–86 

dome material and from ~6 to 15 weight percent in the 2004–6 

dome dacite, although in all three samples Al2O3 contents 

cluster between 10 and 13 weight percent (fig. 7). Similarly, 

total Fe contents range from ~10 to 18 weight percent FeO* in 

the ash, 1980–86 dome, and 2004–6 dome (fig. 7; Rutherford 

and Devine, this volume, chap. 31; Thornber and others, this 

volume, chap. 32).

The thickness of the outermost disequilibrium-reaction 

rims on amphibole phenocrysts in the ash varies widely, with 

thick reaction rims (>15–20 μm) present on some grains. 

Disequilibrium-reaction rims on amphiboles from new dome 

dacite typically are ~5 μm thick, sporadically 6 to 10 μm thick, 

and rarely 50 to 100 μm thick, with larger rims commonly 

associated with amphibole xenocrysts (Thornber and others, this 

volume, chap. 32). Variations in thickness of amphibole disequilibrium-reaction rims decreased over the course of the eruption, 

with only rare thick rims in March 8, 2005, tephra. The presence 

of thick-rimmed amphibole grains is characteristic of 1980–86 

magmas (Rutherford and Hill, 1993) but differs from the uni-
634 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample Analysis SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 SO2 Cl Total

E2A03_A1 2 73.30 0.47 14.72 1.19 0.05 0.22 1.40 4.37 2.75 0.09 0.01 0.14 98.71

E2A03_A1 7 74.00 0.27 15.91 0.89 0.00 0.06 2.09 4.28 2.83 0.07 0.00 0.04 100.44

A20_10_11 11 72.32 0.26 15.02 1.21 0.04 0.09 1.66 4.77 2.28 0.08 0.01 0.12 97.86

A20_10_16 2 76.81 0.24 13.79 1.42 0.02 0.16 0.52 2.92 4.52 0.04 0.00 0.09 100.53

A20_10_16 4 77.63 0.34 12.64 0.97 0.00 0.05 0.57 3.15 4.21 0.03 0.00 0.02 99.61

A21_10_20 14 74.75 0.34 14.24 0.53 0.00 0.02 1.24 3.97 3.41 0.07 0.01 0.01 98.59

A04_11_2 16 71.52 0.22 18.38 0.97 0.00 0.06 2.70 4.26 3.28 0.06 0.00 0.00 101.45

JP_1_14A 1 75.49 0.22 14.00 0.47 0.00 0.10 1.00 4.47 4.12 0.05 0.01 0.05 99.98

JP_1_14A 2 76.06 0.23 12.60 1.32 0.00 0.00 0.33 5.49 3.19 0.18 0.01 0.05 99.46

JP_1_14A 3 73.83 0.18 15.34 0.63 0.02 0.04 1.70 4.45 4.49 0.03 0.01 0.01 100.73

JP_1_14A 6 76.01 0.22 13.78 0.53 0.00 0.09 1.22 2.77 5.33 0.23 0.01 0.03 100.22

JP_1_14A 9 78.38 0.16 12.69 0.64 0.01 0.03 1.29 1.62 5.01 0.04 0.00 0.03 99.90

JV_1_19 4 74.14 0.26 16.05 1.44 0.05 0.07 1.50 5.54 3.03 0.09 0.00 0.02 102.19

JV_1_19 8 71.17 0.09 16.05 0.42 0.01 0.03 1.68 4.55 2.28 0.05 0.02 0.03 96.38

SH304-2 1 A9b hb1 75.2 0.19 13.5 0.87 0.00 0.01 0.88 2.32 5.09 0.01 nd 0.103 98.27

SH305-1 1 kc g-22 76.6 0.34 12.4 1.35 0.01 0.03 0.27 3.01 5.72 0.08 nd 0.110 99.92

1

Representative analyses of dome samples SH304 and SH305 from Pallister and others (this volume, chap. 30).

Feldspar

0

5

10

15

20

25

30

35

50 60 70 80 90 100

SiO2, IN WEIGHT PERCENT

Al2O3, IN WEIGHT PERCENT

Table 2. Electron microprobe analyses of clean, juvenile glass.

[Juvenile glass as described in text is based on MgO and K2O fields defined by glass and inclusion analyses of samples SH304 and SH305. Specific analysis 

from among a suite is listed in “Analysis” column. Total Fe reported as FeO. nd, not determined. Sample prefixes MSH04 and MSH05 have been removed; see 

table 1 for complete sample numbers.]

Figure 4. Silica-variation 

diagram showing Al2O3 versus 

SiO2

 contents in groundmass 

glass of ash fragments. 

Compositional range between 

feldspar and quartz end members 

is characteristic of highly 

crystalline groundmass. Red data 

points, clean glass (see text for 

explanation). Inset X-ray map 

(Al Kα) shows crystallization of 

feldspar (orange) and quartz 

(blue) in juvenile groundmass.
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 635

68 70 72 74 76 78 80

F

4

3

2

1

0

5

6

K2O, IN WEIGHT PERCENT

D

4

3

2

1

5

6

K2O, IN WEIGHT PERCENT

Pre-2004 glass and inclusions Sample SH304 and 305 glass and inclusions

B

4

3

2

1

5

6

7

K2O, IN WEIGHT PERCENT

1.2

0.9

0.6

0.3

A

MgO, IN WEIGHT PERCENT

0.9

0.6

0.3

0

66 68 70 72 74 76 78

Nov. 4 dome collapse (MSH04MR_11_4)

Jan. 13 dome collapse (MSH05JP_1_14A)

Jan. 16 vent explosion (MSH05JV_1_19)

Mar. 8 vent explosion (MSH05DRS_3_9_4)

E

SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT

0.9

0.6

0.3

Oct. 5-11 (MSH04A20_10_11)

Oct. 4-12 (MSH04A09_10_12)

Oct. 11-16 (MSH04A20_10_16)

Oct. 15-20 (MSH04A21_10_20)

Oct. 16-Nov. 2 (MSH04A04_11_2)

C

MgO, IN WEIGHT PERCENT

Oct. 4 vent explosion (MSH04E2A03_A1)

Oct. 1 vent explosion (MSH04E1DZ_1)

Oct. 5 vent explosion (MSH04E3RANDLE_2)

EXPLANATION

Figure 5. Silica-variation diagrams of MgO and K2O versus SiO2 contents in clean glass (see text for explanation). Blue field, 

pre-2004 tephra and dome glass compositions (Sarna-Wojcicki and others, 1981a; Melson, 1983); yellow field, glass and meltinclusion compositions for dome samples SH304 and SH305 from current explosions (see Pallister and others, this volume, 

chap. 30). Plots are paired to show tephra from early vent explosions (A, B), tephra from October 5 to November 2, 2005 (C, D),

and tephra from rockfalls and vent explosions from November 4, 2004, to March 8, 2005 (E, F).
636 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

N=13 E1_DZ_1

N=11 E2_A03_1

N=6 RANDLE_2

N=10 A20_10_11

N=14 A20_10_16

N=13 A09_10_12

N=11 A21_10_20

N=16 A04_11_2

N=14 MR_11_4

N=31 JP_1_14

N=20 JV_1_19

N=32 1980-1986

En Fs

Wo 20% 20%

N=9 DRS_3_9_4

En Fs

20%

20% 40% 60% 80%

20%

N=19 SH304-2A/2G

6

7

8

9

10

11

12

13

14

15

9 10 11 12 13 14 15 16 17 18 19

FeO*, IN WEIGHT PERCENT

Al2O3, IN WEIGHT PERCENT

Source of amphibole

EXPLANATION

2004–5 ash 2004–6 dome lava

1980–86 eruptive products

Figure 6. Compositions of hypersthene phenocrysts 

analyzed from 1980–86 Mount St. Helens, dome sample 

SH304-2A, and 2004–5 ash samples. Shaded field, range 

of 1980–86 hypersthene compositions from top panel.

Figure 7. Al2

O3 versus total Fe (FeO*) in amphiboles in ash 

samples. Blue field, amphibole in 1980–86 eruptive material; 

yellow field, amphibole in 2004–6 dome material.

formly thin (~5 μm thick) decompression rim around 2004–5 

dacitic amphiboles, an observation suggesting that the early 

ash samples have older Mount St. Helens material mixed with 

juvenile ash particles.

In addition to the presence of thick reaction rims on 

amphibole phenocrysts in the ash, the identification of 

clinopyroxene and rare olivine phenocrysts in the early ash 

supports the interpretation that these samples were a heterogeneous mixture containing some proportion of older Mount St. 

Helens eruptive material. Discrete clinopyroxene phenocrysts 

are present in the early ash samples but rare in the new dome 

material, where they typically are associated with xenolith 

fragments. Clinopyroxene phenocrysts were not observed in 

the ash samples collected after October 20, 2004.

Owing to the compositional variation of hypersthene and 

amphibole phenocrysts in the ash samples and the complete 

compositional overlap between these phases in the new dome 

and 1980–86 dome, mafic-mineral phases are of limited value 

for identification of juvenile material and are not discussed 

further here. (See Rutherford and Devine, this volume, chap. 

31, and Thornber and others this volume, chap. 32, for discussion of the textural and compositional variations of Mount St. 

Helens 2004–5 amphiboles.)

Feldspars

Feldspar phenocrysts in ash, gouge, and crater debris vary 

widely in composition. Feldspar phenocryst compositions in 

the ash range from An87 to An28, overlapping samples from 

both the new dome (An53–33) and 1980–85 dome (An51–34). As 

with mafic-mineral phases, major-element contents in feldspar 
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 637

in the ash and new dome samples are indistinguishable from 

older Mount St. Helens eruptive material (Rowe and others, 

2005; Streck and others, this volume, chap. 34). A total of 266 

LA-ICP-MS analyses were completed on 14 of the samples 

included in this study. Significant variations were noted in the 

contents of Ba (14.5–224 μg/g), Sr (532–1,603μg/g), La (1.0–

9.7μg/g), Pb (0.3–7.0μg/g), and Li (6.9–48.5μg/g) (Rowe and 

others, 2008). La has a well-defined, and Pb a weakly defined, 

positive correlation with Ba, whereas Sr is negatively correlated with Ba (fig. 8). Correlations between trace elements 

(excluding Li) and anorthite content are believed to be due 

both to variations in feldspar/melt partition coefficients with 

anorthite content and to changes in melt composition as a 

result of fractionation of plagioclase, hornblende, hypersthene, 

and oxides during melt evolution (Kent and others, this volume, chap. 35).

Li in feldspar, in contrast to other trace elements, does not 

correlate with other major or trace elements. In addition, on 

the basis of feldspar/melt partitioning, the highest Li contents 

would require >>200μg/g Li in the melt (Bindeman and others, 1998), significantly greater than that observed in the bulk 

dacite (21–28μg/g; Kent and others, 2007; Thornber and others, 2008b). In ash samples collected after November 4, 2004, Li, IN MICROGRAMS PER GRAM

10

20

30

40

50

Pb, IN MICROGRAMS PER GRAM

Ba, IN MICROGRAMS PER GRAM Ba, IN MICROGRAMS PER GRAM

0

1

2

3

4

5

6

7

0 50 100 150 200

250

500

750

1,000

1,250

1,500

1,750

2,000

0

La, IN MICROGRAMS PER GRAM Sr, IN MICROGRAMS PER GRAM

0 50 100 150 200 250

2

4

6

8

60

MSH04E1DZ_1

MSH04E2A03_A1

MSH04E3RANDLE_2

MSH04A20_10_11

MSH04A21_10_20

MSH04A04_11_2

MSH04MR_11_4

MSH05JP_1_14A

MSH05JV_1_19

MSH04A20_10_16

MSHDRS_3_9_4

SH300-1

Li contents are more homogenous and consistently lower relative to earlier ash samples, with a maximum Li content of 30

μg/g (fig. 9). This decrease in Li content was also observed in 

dome samples collected after Mount St. Helens dome sample 

SH304, which was estimated to have been erupted on or about 

October 18, 2004 (Kent and others, 2007; Pallister and others, 

this volume, chap. 30; Kent and others, this volume, chap. 35).

Discussion

October 1–5, 2004, Ash Explosions

Identification of a Juvenile Component

Identification and quantification of the juvenile component in the products of early Mount St. Helens eruptive events 

can be used to evaluate the likely course of a reawakening 

volcano. By comparing glass compositions and feldspar traceelement contents in the ash with those of 1980–86 and 2004–5 

dome materials, the proportion of juvenile material (if present), as previously defined, in the early erupted ash may be 

Figure 8. Selected trace-element contents in feldspar in ash samples.
638 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1

2

3

4

5

1

2

3

4

1

2

3

4

1

2

3

4

5

6

7

1

2

3

4

5

6

7

1

2

3

4

5

6

7

1

2

3

4

5

6

7

<10

10-12

12-14

14-16

16-18

18-20

20-22

22-24

24-26

26-28

28-30

30-32

32-34

34-36

36-38

38-40

40-42

42-44

44-46

46-48

48-50

>50

Li, IN MICROGRAMS PER GRAM (μg/g)

NUMBER OF SPECIMENS

MSH04E1DZ_1

MSH04E2A03_A1

MSH04E3RANDLE_2

MSH04A20_10_11

MSH04A20_10_16

MSH04A21_10_20

SH300-1

(10/1/04)

(10/4/04)

(10/5/04)

(10/5/04-10/11/04)

(10/11/04-10/16/04)

(10/15/04-10/20/04)

(10/15/04)

2

4

6

8

10

1

2

3

4

5

6

7

8

1

2

3

4

1

2

3

4

1

2

3

4

1

2

3

4

5

6

7

1

2

3

4

<10

10-12

12-14

14-16

16-18

18-20

20-22

22-24

24-26

26-28

28-30

30-32

32-34

34-36

36-38

38-40

40-42

42-44

44-46

46-48

48-50

>50

MSH04A04_11_2

(10/16/04-11/2/04)

SH303-1

(10/18/04)

MSH04MR_11_4

(11/4/04)

MSH05JP_1_14A

(01/13/05)

MSH05JV_1_19

(01/16/05)

SH307-1

(02/12/05)

MSH05DRS_3_9_4

(03/8/05)

Figure 9. Histograms of Li contents in feldspar from each of 14 analyzed samples. Feldspar Li contents >30 μg/g (dashed line) in 

ash samples collected before November 4, 2004, coincide with Li contents in juvenile feldspar from dome sample SH304. Eruption 

dates of samples are given in parentheses.
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 639

calculable. Most major-element contents in glass overlap with 

those in eruptive products of 1980–82 (Sarna-Wojcicki and 

others, 1981a; Melson, 1983) and pre-1980 (Sarna-Wojcicki 

and others, 1981a). However, matrix glass and melt inclusions 

from 2004–5 Mount St. Helens samples SH304 and SH305 

have distinctly lower MgO contents (<0.5 weight percent) and 

higher K2O contents (max 5.5 weight percent) at comparable 

SiO2 contents (fig. 5; table 2; Pallister and others, this volume, chap. 30). On the basis of variations in MgO and K2

O 

contents, glasses can be divided into what appear to be older 

glassy ash fragments and juvenile material. Judging from these 

criteria, none of the clean-glass analyses from the October 

1 and 5, 2004, explosions are considered juvenile, although 

glasses with juvenile characteristics were observed in ash from 

the October 4, 2004, explosion (fig. 5). However, the high 

crystallinity of the groundmass, resulting in few clean-glass 

analyses, precludes definitive estimates of the proportion of 

juvenile glass observed in early eruptive material. In addition, 

all of the analyzed pumice fragments fall within the compositional fields defining older volcanic deposits (fig. 5).

Backscattered electron images provide a textural comparison between juvenile crystalline groundmass (fig. 10B) and 

a pumiceous glass fragment compositionally similar to older 

volcanic deposits (fig. 10D). Both textural and compositional 

comparisons between the various groundmass types are critical 

to avoiding a common misconception that glassy ash particles 

(for example, fig. 10D) are always representative of the juvenile 

component associated with an eruption, as previously attributed 

(for example, Sarna-Wojcicki and others, 1981b; Swanson and 

others, 1995; Watanabe and others, 1999).

Previous studies of 1980s tephra and dome samples indicate 

that the Li content in feldspar from 1980–85 lavas and tephras 

is as high as ~25μg/g (fig. 11; Berlo and others, 2004; Kent 

and others, this volume, chap. 35). In contrast, the Li content in 

plagioclase from early 2004 dome sample SH304-2A (erupted 

approx. Oct. 18, 2004) ranges from 28 to 43μg/g Li, distinctly 

higher than that in older Mount St. Helens feldspar (Kent and 

others, this volume, chap. 35). This difference provides a means 

to trace the contribution to ash samples collected between 

October 1 to October 5, 2004, from juvenile magma (figs. 9, 

11). As previously discussed, after dome sample SH304-2A 

was collected, Li contents dropped to levels similar those in 

1980–85 materials (<30 μg/g), thus limiting the usefulness of this 

approach to material produced only at the start of an eruption.

Measured Li contents in plagioclase in ash erupted from 

October 1, 2004, to March 9, 2005, form two populations. 

The Li contents of most plagioclase phenocrysts range from 

10 to 25 μg/g and appear to be broadly normally distributed 

around a mean of 19 μg/g (fig. 9), whereas a smaller subset 

(~25 percent) of grains have distinctly higher Li contents (max 

46 μg/g). On a cumulative-probability plot (fig. 11B), most 

plagioclase phenocrysts lie on a straight-line segment, with a 

mean of ~18 μg/g, consistent with a near-normal distribution; 

however, a distinct break in slope occurs at ~30 μg/g, suggesting that the higher Li contents may derive from a different 

population. If we divide the samples into two groups, using a 

Li content of 30 μg/g as a boundary, application of Student’s 

t test (appropriate for comparing the means of small samples 

that may show some departure from a normal distribution; for 

example, Borradaile, 2003) suggests that these two sample 

Figure 10. Backscattered 

electron images of dome rock, 

ash fragments, and pumice 

fragments. A, Groundmass 

from dome sample SH304. B,

Ash fragment from sample 

MSH04MR_11_4 (feldspar 

Li content of 33.7 μg/g). 

C, Glassy fragment of ash 

with extensive plagioclase 

microlite crystallization from 

sample MSH04E2A03_A1. D,

Pumice fragment from sample 

MSH04E2A03_A1.

100 μm

C

500 μm

A

100 μm

B

50 μm

D
640 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

50

40

30

20

10

0

.01 .1 1 5 20 50 80 95 99 10 30 70 90 99.9

PROBABILITY, IN PERCENT

Feldspar in 2004–5 Ash

0

5

10

15

20

25

30

<10

10-1212-14

14-1616-1818-2020-2222-2424-2626-2828-30

30-3232-34

34-36

36-38

38-4040-4242-44

1980–1985 (N=535)

SH304-2A/2C (N=53)

A

B

RELATIVE FREQUENCY, IN PERCENT

Li, IN MICROGRAMS PER GRAM

Li, IN MICROGRAMS PER GRAM

Figure 11. Li contents in feldspar in dome rocks. A, Histogram 

of Li contents in feldspar in 1980–85 dome rocks (black bars, 

N=535) and in sample SH304 (gray bars; N=53). Data for 1980–85 

samples from Berlo and others (2004) and Kent and others (this 

volume, chap. 35); data for sample SH304 from Kent and others 

(this volume, chap. 35). B, Li content versus probability for 2004–5 

feldspar in ash. In a single normally distributed population, this 

relation would plot as a straight line. Distinct change in slope at 30 

μg/g Li content suggests that multiple populations of plagioclase 

grains are present in ash samples.

groups are significantly different at the >99.9-percent confidence level. Furthermore, the >30-μg/g-Li sample group is 

indistinguishable from plagioclase from samples SH304-2A 

and SH304-2C, whereas the <30-μg/g-Li sample group is 

indistinguishable from plagioclase from the 1981–85 dome 

samples (table 3). From this we suggest that the plagioclase 

containing >30 μg/g Li in ash produced by explosions from 

October 1 to 5, 2004, was derived from Li-enriched juvenile 

magma similar to that sampled by dome samples SH304-2A 

and SH304-2C.

Li enrichment in feldspar, which is interpreted to result 

from transfer of a magmatic vapor phase, is expected to occur 

only within juvenile material, not in older conduit material 

or debris covering the conduit (Kent and others, 2007). This 

interpretation is supported by (1) the absence of anomalous 

Li contents in plagioclase in gabbroic inclusions; and (2) the 

high Li contents (max ~200 μg/g) in amphibole-hosted melt 

inclusions, consistent with concentrations required in the melt, 

as estimated from plagioclase/melt partitioning (Bindeman 

and others, 1998; Kent and others, 2007). On the basis of this 

evidence, 8 out of 39 (~20 percent) of the feldspars from the 

October 1 to 5, 2004, ash deposits are juvenile. Though not 

definitive, the proportion of juvenile material also increases 

from October 1 to October 4, 2004, suggesting a possible 

increased involvement of juvenile material in explosions. One 

explanation for this increase is upward movement of magma 

within the shallow conduit, consistent with the appearance 

of hot dacite on the surface of the crater floor on October 11, 

2004.

A discrepancy is evident between estimates of the 

proportion of juvenile material in ash based on Li contents 

in feldspars versus the compositions of glass fragments. A 

possible cause for this discrepancy is that only a few cleanglass analyses were obtained as a result of quartz and feldspar 

microlite crystallization (figs. 4, 10). This same microcrystalline texture is also observed in dome samples, suggesting that 

the shallow magmatic crystallization observed in dome rocks 

occurred before fragmentation of the dacite. Because most 

of the juvenile groundmass had crystallized before eruption, 

clean-glass analyses ultimately underestimate the proportion 

of juvenile material in early ash samples.

Evidence for a Volatile Eruption Trigger

The high Li contents measured in early-erupted feldspar 

from the ash (max 48μg/g) and dome samples (max 41μg/g) 

may indicate an increase in volatile components at the initiation 

of the 2004 eruption. As previously discussed, the high Li content measured in feldspar would require a Li content in melt of 

~200μg/g, on the basis of plagioclase/melt partition coefficients 

(Bindeman and others, 1998), despite the lower Li contents (21–

28μg/g) measured in bulk dacite. As suggested by Berlo and 

others (2004), the Li-rich feldspar could be explained by enrichment of shallow-stored magma through upward movement of an 

alkali-enriched vapor from deeper within the magmatic system. 

Because the highest Li contents were measured only before 

collection of sample SH305-1 (estimated eruption date, Nov. 20, 

2004), this difference suggests greater vapor enrichment during 

the early stages of the eruption. Kent and others (2007) estimate that vapor enrichment must have occurred within the year 

before the eruption (potentially within just a few days) before 

the eruption, because earlier enrichment would have resulted 

in diffusional equilibration of feldspar in gabbroic inclusions—

which have lower Li contents than groundmass plagioclase. The 

decrease in feldspar Li contents over time observed in the dome 

dacite samples may indicate that the volatile-enriched magma 
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 641

Feldspar in dome rocks Feldspar in 2004–5 ash

2004 1981–85 1980 Li-poor Li-rich

Mean 33.4 17.1 13.1 17.6 37.5

Standard deviation (1σ) 3.8 4.3 4.6 5.8 4.5

Number of samples (N) 53 28 507 177 18

Table 3. Statistics of Li contents in feldspar in 2004–5 ash, new dome lava, and 1980–85 dome lava.

[2004 sample is SH304-2A, erupted about Oct. 18 and collected Nov. 4, 2004, showing data from Kent and others 

(this volume, chap. 35). Li contents in feldspar for 1980 materials from Berlo and others (2004). Data for 1981–85 

from samples SH100, SH141, and SH187, analyzed at Oregon State University (Kent and others, this volume, chap. 

35). Li contents in feldspars from ash produced during the 2004–5 eruption are divided into“Li-poor” (<30 μg/g) and 

“Li-rich” (>30 μg/g) groups on the basis of statistical analysis (fig. 11; see text for discussion).]

was only a shallow cap on a larger magma body (Kent and others, this volume, chap. 35).

A CO2 output of 2,415 t/d, measured on October 7, 2004, 

followed by a steady decline in ensuing months also suggests 

a greater volume of gas early in the eruption, although the 

higher gas output could also result from higher extrusion rates 

early in the eruption (Gerlach and others, this volume, chap. 

26). Together, both the high Li contents in feldspar and the 

highest CO2 output measured on October 7, 2004, may indicate the presence of an excess volatile phase, partly resulting 

from volatile components transferred from a deeper magma, 

before the onset of the eruption. An increase in vapor pressure 

may have helped drive the explosions of October 2004. (For a 

more detailed discussion on the formation of the vapor phase, 

see Berlo and others, 2004, and Kent and others, this volume, 

chap. 35.)

Ash Generation

Airborne ash at Mount St. Helens is produced by both 

vent explosions and rockfalls or rock avalanches. Two general 

mechanisms are typically described for the generation of ash 

during vent explosions. In magmatic explosions, exsolution 

and expansion of gases in magma during its ascent leads to 

vesiculation and fragmentation of the magma (Heiken, 1972; 

Cashman and others, 2000). In contrast, phreatomagmatic 

eruptions result primarily from the rapid expansion, at Mount 

St. Helens, of meteoric water as it interacts with magma and 

(or) hot rock at depth (Morrissey and others, 2000).

Ash from Mount St. Helens’ 2004–5 vent explosions is 

dispersed more widely than ash from rockfalls, with fallout 

reported as far as Ellensburg, Wash., after the March 8, 2005, 

vent explosion (Mastin and others, 2005; Scott and others, 

this volume, chap. 1). Rockfalls and rock avalanches in this 

eruption resulted from dome growth and oversteepening (Vallance and others, this volume, chap. 9). Rockfalls and rock 

avalanches, with elutriation of fine ash particles, occur much 

more frequently than vent explosions and are typically smaller, 

sending ash as high as 3,000 m into the air before it falls proximal to the edifice (Scott and others, this volume, chap. 1). In 

addition to vent explosions and rockfalls, small volumes of 

airborne ash may be distributed by elutriation during steaming 

of the dome or resuspended off crater walls by strong winds—

relatively low energy mechanisms for ash transportation.

Mechanisms of ash generation and distribution of the 

2004–5 Mount St. Helens eruption are not easily discernible 

by using criteria of grain shape, volume of glass, or textural 

and compositional heterogeneity of ash particles (see above 

descriptions). In this eruption, very little glassy material was 

present in the ash, and <5 percent of the clean-glass samples 

are considered juvenile. Juvenile groundmass observed in the 

dome and in the ash is mostly crystalline (quartz and feldspar), 

with localized vesiculation and devitrification (figs. 10A, 10B). 

Water-by-difference measurements of melt inclusions in phenocrysts from dome samples SH304 and SH305 support gas 

emission data which, for this eruption, indicate that juvenile 

magma had degassed extensively before eruption (Gerlach and 

others, this volume, chap. 26; Pallister and others, this volume, 

chap. 30). The degassed and crystallized nature of the magma 

may have reduced the role of expansion of magmatic volatile components in the generation of ash, during either vent 

explosions or rockfalls and rock avalanches. In addition, the 

abundance of lithic fragments and paucity of glass shards or 

pumice fragments suggest that vent explosions (Oct. 1, 4, and 

5, 2004; Jan. 16, 2005; Mar. 8, 2005) are dominantly phreatic in origin and that crystallization of the rising magma had 

occurred before the explosive events.

By comparing geochemical analyses of feldspar and glass 

in conjunction with the backscattered electron imaging of ash 

derived from vent explosions and rockfalls or rock avalanches, 

we attempted to distinguish between possible causes and 

mechanisms for the dispersal of ash. Major-element compositions of matrix glass and glass inclusions from the January 13, 

2005, rockfall ash correlate well with juvenile compositions 

determined from glass analyses of dome samples SH304 and 

SH305 (fig. 5). In contrast, less than half of the glass analy-
642 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

ses from the January 16, 2005, vent explosion correlate with 

juvenile fields. This discrepancy demonstrates that the vent 

eruptions entrain more of the exotic fragments derived from 

crater debris in the vent area.

A distinct contrast is observed when comparing glass 

compositions from the rockfall events of November 4, 2004, 

and January 13, 2005. Whereas the ash from the November 

4, 2004, event is dominated by older glass, the ash from the 

January 13, 2005, event is dominated by juvenile glass (fig. 

5), suggesting that the part of the early dome (or uplift area) 

responsible for ash generation contained a significant amount 

of older material but by January 2005 was composed dominantly of juvenile dacitic material. This conclusion is also 

supported by field observations and pre-November 2004 dome 

samples (Pallister and others, this volume, chap. 30; Reagan 

and others, this volume, chap. 37).

Juvenile ash fragments from large rockfalls (Nov. 4, 2004, 

and Jan. 13, 2005) and vent explosions (Jan. 16, 2005, and Mar. 

8, 2005) have nearly identical groundmass textures, as described 

above (fig. 12). Both processes produce strikingly similar grain 

shapes. Ash particles dominated by large phenocrysts commonly reflect the general shape of those phenocrysts and are 

more angular, whereas ash particles dominated by groundmass 

typically range in shape from subangular to round. In principle, particle shape in magmatic eruptions is controlled by the 

shape and abundance of vesicles, whereas glassy fragments 

in phreatomagmatic eruptions are commonly pyramidal or 

blocky (Heiken, 1972). In both magmatic and phreatomagmatic 

eruptions, however, relatively rapid cooling should result in the 

generation of glassy material. In contrast, in an experimental 

study of dacite fragmentation using Mount St. Helens 1980 

gray dacite, similar to the cryptodome with low H2O contents 

and composed of ~30 volume percent phenocrysts, the material 

produced brittle, angular fragments (Alidibirov and Dingwell, 

1996). This material more closely resembles some of the particles observed in recent rockfalls and explosions, suggesting 

Figure 12. Backscattered 

electron images of 

characteristic ash particles 

from dome rockfalls (A, B), 

vent explosions (C, D) and 

dome gouge (E, F). A, Rounded 

ash fragment from sample 

MSH05JP_1_14A, from rockfall 

on January 13, 2005. B, More 

highly disaggregated ash 

particle from same sample as 

in figure 12A. C, Subrounded 

ash fragment with grain shape 

controlled by larger feldspar 

and hypersthene phenocrysts 

from sample MSH05JV_1_19, 

from explosion of January 16, 

2005. D, Largely devitrified ash 

particle (see X-ray map in fig. 

3) from explosion of March 

8, 2005. E, Low-magnification 

image of disaggregated dome 

gouge from sample SH307-1. 

F, Enlargement of rounded 

clast in figure 12E, showing 

characteristic ash morphology.

100 μm

100 μm 200 μm

200 μm

1,000 μm 200 μm

A B

C D

E F
29. Identification and Evolution of the Juvenile Component in 2004–2005 Mount St. Helens Ash 643

that fragmentation may have occurred by brittle failure without 

frothing or vesiculation of the magma. Similarities in grain 

shape and texture of the juvenile matrix from different types of 

events suggest that the ash is generated by related processes.

The disaggregation of the outer 1 to 3 m of the dome 

(dome fault gouge; see Cashman and others, this volume, 

chap. 19, for a detailed description) provides an additional 

potential source for ash particles, independent of magmatic 

fragmentation. Gouge particles are dominantly subangular to 

rounded, similar to the particles observed in a large proportion of the ash (figs. 12E, 12F). Groundmass textures of dome 

gouge (sample SH303-1, table 1) extruded on October 18, 

2004, are heterogeneous, suggesting that at that time the gouge 

contained a high percentage of exotic wallrock fragments.

The heterogeneity of the dome gouge may also explain 

the dominance of glass compositions representative of older 

eruptive events in the November 4, 2004, dome-collapse 

ash. In addition, Li contents in feldspar from the November 

4, 2004, ash closely resemble those of the gouge, which are 

dominantly low (<30 μg/g) and do not correlate well with 

those of either dome samples SH304 or SH305 (figs. 9, 11; 

Kent and others, this volume, chap. 35). Recall that feldspars 

with Li contents below 30 μg/g are a statistically distinct group 

from the Li-rich population and correlate with older Mount St. 

Helens eruptive products (fig. 11). In contrast, the January 13, 

2005, dome-collapse ash is significantly more homogenous, 

with glass compositions matching that of the growing dome 

(fig. 5). This homogeneity was also observed in the dome fault 

gouge (sample SH307-1, table 1) collected on February 22, 

2005, with groundmass textures resembling those of the growing dome, and groundmass textures and grain shapes similar to 

those observed in the ash.

Textural and geochemical evidence presented here 

suggests that airborne ash in the 2004–6 eruption is largely 

generated from the dome fault gouge. Large dome-collapse 

events and explosions may also produce a significant component of material resulting from brittle failure of the dacite 

during expansion of volatile components, despite low volatile 

contents. These particles would have groundmass textures 

similar to dome gouge but would likely have angular to subangular grain shapes instead of the more rounded particles found 

within the gouge (Alidibirov and Dingwell, 1996; Cashman 

and others, this volume, chap. 19).

The particle size distributions of ash samples generated 

from both dome collapse and vent explosions overlap significantly, largely owing to the distance to the vent at which 

samples were collected; therefore, particle size is inconclusive 

for distinguishing source mechanism (fig. 13). Mean particle 

size varies but generally decreases with increasing distance 

from the vent. In addition, the mean particle size of ash from 

vent explosions is significantly greater than that from domecollapse events, despite collection farther from the vent, 

illustrating the farther transport of coarse ash during more 

explosive events. The particle size distribution of dome gouge 

differs from that of ash deposited from either dome-collapse 

events or vent explosions (fig. 13). However, because variables 

resulting in mechanical sorting of the ash, including explosivity of the eruption, windspeed, height of the ash cloud, and 

distance to the vent, are not taken into account in this comparison, particle size distributions cannot be used to corroborate 

textural evidence which suggests that a large proportion of the 

ash is derived from the dome fault gouge.

Distal ash believed to remobilize during steaming or 

from strong winds is dissimilar to ash from the rockfalls and 

vent explosions. The ash collected between October 6 and 

November 2, 2004, contains abundant exotic particles and 

pumice fragments and substantially less of the highly crystalline groundmass characteristic of the juvenile dome material 

commonly erupted during discrete ash-producing events. 

Only a small proportion of clean juvenile glass is present in 

all of the ash samples collected during this interval, and no 

systematic variation is observed in the proportion of juvenile 

glass over time, despite the appearance of hot dacite on October 11, 2004. Similar results are observed for Li contents 

in feldspar during this period, with high Li contents in only 

~4 percent (3 of 79) of analyzed feldspar, despite identifying ~22 percent (5 of 23) of juvenile feldspar in the fine 

material in the sample collected from the October 20, 2004, 

crater debris flow (sample SH300-1, table 1). Remobilization 

of ash particles from the crater walls by strong winds, and 

elutriation of fine ash particles from rising steam, appear to 

have (1) preferentially transported low-density material, such 

as pumice fragments; and (2) resulted in apparent dilution 

of the estimated proportion of juvenile material as a result 

of greater contamination from older volcanic rocks. Despite 

the continuous ash collection, therefore, sampling of discrete 

events appears to provide a more representative depiction of 

current eruptive conditions.

10

0

20

30

40

50

60

70

80

90

100

0.1 1 10 100 1,000 10,000

PARTICLE SIZE, IN MICROMETERS

(0.2)

(1.6) (14.3)

(3.5)

CUMULATIVE WEIGHT PERCENT

Figure 13. Particle size distribution for ash derived from rockfalls 

and rock avalanches (black solid lines, N=13), vent explosions 

(dashed lines, N=7), and dome gouge (red solid lines, N=2). Wide 

variation is partly attributable to varying distance from vent 

(numbers in parentheses are distance in kilometers to vent) and 

explosivity of eruption. Note that even at greater distance from vent, 

explosive deposits generally have a larger average particle size.
644 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Conclusions

This textural and geochemical examination of 2004–5 

Mount St. Helens ash provides several important observations and conclusions with regard to early eruptive events and 

processes of ash generation and transportation.

1. The Li content in feldspar from dome sample SH304 

is >30 μg/g, significantly higher than in older Mount St. 

Helens eruptive material, allowing for the identification of 

juvenile plagioclase in the ash. Li contents may indicate 

an increase in the proportion of juvenile material from 

October 1 through 4, 2004. High Li contents in feldspar 

are useful as a tracer for juvenile material only during the 

beginning stages of the eruption—within 6 weeks after 

eruptive onset—because after about November 20, 2004 

(estimated eruption date of sample SH305), Li contents in 

feldspar dropped to levels similar to that observed during 

the period 1980–86. The decrease in Li content also suggests that the volatile-enriched magma was only a shallow 

cap on a larger magma body.

2. The clean glass in ash samples is mostly older volcanic material. Juvenile groundmass is typically crystalline 

(quartz and feldspar), with localized devitrification and 

patchy, very fine vesiculation, contrary to the traditional 

assumption that clean glass is juvenile volcanic material. 

Estimates of the proportion of juvenile material based on 

glass analyses are significantly lower than those based on 

Li content in plagioclase.

3. Groundmass textures, grain shapes, and compositional similarities among ash from vent explosions, 

rockfalls, and dome fault gouge particles suggest that the 

gouge is a significant source of ash material. Despite differences in particle size, the grain shapes and textures of 

particles from rockfalls and vent explosions are relatively 

similar, suggesting that the ash is derived by a similar 

mechanism in both events. The heterogeneity observed 

in early ash samples correlates with observations that the 

early extruded dome material and fault gouge contained 

a significant proportion of older volcanic material. The 

increase in the homogeneity of ash over time is consistent 

with a similar trend in dome and fault gouge samples.

4. The presence of an excess volatile phase, as indicated by 

the high Li contents in feldspar, may indicate the presence of 

a greater proportion of vapor at the top of the conduit where 

shallow-stored magmas had stalled. This excess vapor may 

have been the catalyst for initiation of the 2004–6 eruption.

5. Continuous monitoring of ash samples during periods of relatively low eruptive activity may not provide a 

representative depiction of the current eruptive processes, 

owing to contamination from other debris. Therefore, collection and monitoring of discrete dome-collapse events 

and vent explosions provides the most accurate record of 

the current eruption.

Acknowledgments

We thank all of the Cascades Volcano Observatory staff 

and volunteers who assisted in the collection of Mount St. 

Helens ash, especially S. McConnell, T. Herriott, and A. 

Eckberg, who assisted with drying, sorting, and weighing of 

ash samples, and D. Gooding and D. Ramsey for assistance 

in the field and with the collection site mapping. Electron 

microprobe time at Oregon State University was provided 

by F. Tepley. LA-ICP-MS analyses were performed at the 

W.M. Keck Collaboratory for Plasma Spectrometry, Oregon 

State University. Detailed reviews by J. Larsen and R. Hoblitt 

greatly improved the manuscript. The Mount St. Helens 

petrology working group provided valuable discussion and 

feedback for this study. This research was supported in part by 

National Science Foundation grant EAR 0440382.

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 30

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 Department of Geological Sciences, 1272 University of Oregon, Eugene, 

OR 97403

3

 U.S. Geological Survey, 345 Middlefield Rd, Menlo Park, CA 94025

4

 U.S. Geological Survey, Box 25046, Denver Federal Center, Denver, CO 

80225

5

 American Museum of Natural History, Central Park West at 79th Street, 

New York, NY 10024

Petrology of the 2004–2006 Mount St. Helens Lava 

Dome—Implications for Magmatic Plumbing and Eruption 

Triggering

By John S. Pallister1, Carl R. Thornber1, Katharine V. Cashman2, Michael A. Clynne3, Heather A. Lowers4, 

Charles W. Mandeville5

, Isabelle K. Brownfield4, and Gregory P. Meeker4

Abstract

Eighteen years after dome-forming eruptions ended in 

1986, and with little warning, Mount St. Helens began to 

erupt again in October 2004. During the ensuing two years, 

the volcano extruded more than 80×106 m3 of gas-poor, 

crystal-rich dacite lava. The 2004–6 dacite is remarkably 

uniform in bulk-rock composition and, at 65 percent SiO2, 

among the richest in silica and most depleted in incompatible 

elements of the magmas erupted at Mount St. Helens during the past 500 years. Since shortly after the first spine of 

lava appeared, samples have been collected using a steel box 

dredge (“Jaws”) suspended 20–35 m below a helicopter and, 

occasionally, by hand sampling. As of the spring of 2006, 25 

age-controlled samples have been collected from the seven 

spines of the new lava dome. Samples were obtained from 

both the interiors of spines and from their carapaces, which 

are composed of fault gouge and cataclasite 1–2 m thick. The 

dacite lava is crystal rich, with 40–50 percent phenocrysts. 

The groundmass is extensively crystallized to a cotectic 

assemblage of quartz, tridymite, and Na- and K-rich feldspar 

microlites, raising the total crystal content to more than 80 

percent on a vesicle-free basis in all but the earliest erupted 

samples. Early samples and those collected from near the 

spine margin are more glassy and vesicular that those collected later and from the interior of the spines. Oxide thermobarometer determinations for the earliest erupted samples 

we collected cluster at temperatures of approximately 850oC 

and at an oxygen fugacity one log unit above the nickel-nickel 

oxide (NNO) buffer curve. In contrast, samples from relatively glass-poor samples erupted in late 2004 and early 2005 

have zoned oxides with apparent temperatures that range to 

greater than 950oC. The higher temperatures in these microlite-rich rocks are attributed to latent heat evolved during 

extensive and rapid groundmass crystallization. Low volatile 

contents of matrix glasses and presence of tridymite and 

quartz in the high-silica rhyolite matrix glass indicate extensive shallow (<1 km) crystallization of the matrix, driven by 

degassing of water and solidifying the magma below the level 

of the vent. The mode of eruption of the dacite as a series of 

fault-gouge-mantled spines is explained by this process of 

extensive subvent degassing and solidification.

Although the dacite from this eruption is more silica 

rich than 1980–86 dome rocks, most major and trace element 

concentrations of the 1980–86 and 2004–6 magma batches are 

similar, and magmatic gas emissions have been low and have 

had similar ratios to those of the 1980s, raising the possibility 

that the magma might be residual from the 1980–86 reservoir. 

However, titanium and chromium are enriched slightly relative 

to the most recent 1980–86 and Goat Rocks (A.D. 1800–1857) 

eruptive cycles, and heavy rare-earth-element abundances are 

slightly depleted relative to those erupted during the past 500 

years at Mount St. Helens. These data suggest either addition 

of new gas-poor dacite magma or tapping of a region of the 

preexisting reservoir that was not erupted previously.

A relatively low pressure of last phenocryst growth 

suggests that the magma was derived from near the apex of 

the Mount St. Helens magma reservoir at a depth of about 
648 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

5 km. Viewed in the context of seismic, deformation, and 

gas-emission data, the petrologic and geochemical data can 

be explained by ascent of a geochemically distinct batch 

of magma into the apex of the reservoir during the period 

1987–97, followed by upward movement of magma into a new 

conduit beginning in late September 2004.

The question of new versus residual magma has implications for the long-term eruptive behavior of Mount St. Helens, 

because arrival of a new batch of dacitic magma from the deep 

crust could herald the beginning of a new long-term cycle of 

eruptive activity. It is also important to our understanding of 

what triggered the eruption and its future course. Two hypotheses for triggering are considered: (1) top-down fracturing 

related to the shallow groundwater system and (2) an increase 

in reservoir pressure brought about by recent magmatic replenishment. With respect to the future course of the eruption, 

similarities between textures and character of eruption of the 

2004–6 dome and the long-duration (greater than 100 years) 

pre-1980 summit dome, along with the low eruptive rate of the 

current eruption, suggest that the eruption could continue sluggishly or intermittently for years to come.

Introduction

End members of natural phenomena offer important 

constraints on processes, in part because they provide a means 

to understand boundary conditions. The ongoing eruption of 

Mount St. Helens is such an end member, as it represents a 

sustained low-rate dacite eruption at a potentially explosive 

volcano. A fortunate combination of conditions has allowed us 

to investigate fundamental questions about volcanic plumbing 

systems, triggers for eruptions, and controls on explosivity.

The dacite has largely solidified to crystalline rock 

beneath the surface, such that the first two years of the eruption have been characterized by nearly continuous extrusion 

of lava spines. Furthermore, because the spines have been 

extruded at an angle, the vent has been partially exposed 

through most of the eruption, allowing up-close access by 

helicopter to install instruments, document the character of the 

spines, and collect a series of lava and gouge samples through 

time. The sample collection, along with field observations and 

seismic, deformation, thermal, and gas monitoring, form the 

basis for this and other petrologic papers in this volume.

We have employed an unusual method of sampling borrowed from the oceanographic community. A dredge bucket 

suspended beneath a helicopter was used repeatedly to collect 

samples from the actively extruding spines over the course of 

the eruption. Samples have been distributed widely to petrologists and geochemists at a variety of institutions in the U.S. 

and overseas. Accordingly, petrologic studies of the 2004–6 

eruption were undertaken through a team approach, which 

brought together more than 25 petrologists from a dozen 

institutions to contribute to the interpretation of the eruption 

(Pallister and others, 2005). We thank these scientists and 

acknowledge their contributions and those of our USGS colleagues, as through this community approach and synthesis of 

different types of data we have advanced our understanding.

This report deals primarily with the general petrology of 

the dacite lava and what it tells us about magmatic processes 

when viewed in the broader context of monitoring data and 

Mount St. Helens’ history. We report results from field observations and petrographic studies using optical and scanning 

electron microscopy, microprobe analyses of glasses and Fe-Ti 

oxide minerals, Fourier-transform infrared (FTIR) spectroscopy, and a comparison of the whole-rock major- and traceelement compositions of 2004–6 samples with those of previous eruptive episodes at Mount St. Helens. Other petrologic 

reports in this volume focus on silicate mineral chemistry and 

zoning, isotopic data, phase equilibria, and the petrology and 

chemistry of ash samples, as well as detailed studies of the 

chemistry of the dacite and textures of the fault gouge that 

mantles the lava spines.

Field Geology and Sampling

A chronology of the 2004–6 eruption is given by Scott 

and others (this volume, chap. 1), and details of lava-dome 

growth are described and illustrated by Vallance and others 

(this volume, chap. 9) and summarized in digital elevation 

models and photogeologic maps (Schilling and others, this 

volume, chap. 8; Herriott and others, this volume, chap. 10). 

Here we provide only a brief outline of field relations used to 

establish eruption ages of samples and field observations that 

are important to interpretation of the petrology. The 2004–6 

eruption sequence began with intense seismic unrest, uplift of 

crater floor, and deformation of glacial ice during late September 2004, which were followed by a series of phreatic explosions during the first week of October. Beginning October 11, 

2004, spines of juvenile lava extruded in the deformed area, 

and these have continued to extrude virtually continuously 

ever since. On January 16, 2005, and March 8, 2005, the only 

sizable explosions since early October 2004 blanketed the 

crater floor with ash and ballistic fragments. For safety, a steel 

dredge (fig. 1), sling-loaded to a helicopter, and a few shortduration landings were used to collect rock and gouge samples 

from the spines. A total of 25 locations were sampled on 16 

days during the period between October 2004 and February 

2006, as shown in table 1 and in the photogeologic maps of 

Herriott and others (this volume, chap. 10).

There have been no pumiceous eruptions in this eruptive 

sequence, and, except for small flows formed during collapse 

of the initial spine in October–November 2004 (fig. 2), there 

have been no dome-collapse pyroclastic flows. Descriptions of 

ash samples collected from deposits of the few small explosive 

eruptions and many small rockfalls and rock avalanches are 

given in Rowe and others (this volume, chap. 29), and descriptions of the fault gouge, which mantles the dacite spines, are 

given in Cashman and others (this volume, chap. 19).
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 649

A

B

February 15, 2006

February 15, 2006

Site of image B

Spine 7

Poorly consolidated

gouge on surface

of spine 7

Figure 1. Photographs illustrating use and design of rock dredge. Dredge allowed collection 

from normally inaccessible locations, as shown in the main photos. A, Upper lip of spine 7, where 

fragments of consolidated gouge and lava from close to the gouge carapace were collected on 

February 15, 2006. USGS photo by J.S. Pallister B, View looking down at dredge from helicopter while 

sampling blocks of dacite from near gouge contact. As shown in the inset photos, the dredge is a 

plate-steel frame, 30 by 60 cm and 30 cm deep, beneath which hangs a removable heavy-gauge wirescreen basket for collecting rocks (lower inset). Upper edge of frame is serrated to better bite into 

rock or gouge, giving dredge its informal name, “Jaws.” Basket can be replaced with steel box for 

sampling fine-grained material such as gouge (upper inset). Dredge is lifted from a steel ring attached 

to 1-m-long steel lift arm bolted to sides of frame to allow adjustment of bite angle. A steel leader 

cable 6 mm in diameter and 5 m long is attached to steel ring on lift arm. The ring allows side loads to 

be imparted on dredge to help dislodge it if it becomes wedged in rock face. Steel cable is attached to 

standard electrically released sling-load cable from helicopter. USGS photo by J.S. Pallister.
650 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample No. Spine Collection date Eruption date Limiting dates

SH300 1 10/20/2004 10/15/2004 10/11/04–10/20/04

SH3001 10/20/2004 <19861

SH3012 10/27/2004 <19862

SH302 1 or 2 10/27/2004 10/14/2004 10/12/04–10/27/04

SH304 3 11/4/2004 10/18/2004 10/12/04–11/04/04

SH305 3 1/3/2005 11/20/2004 10/27/04–12/01/04

SH306 4 1/14/2005 12/15/2004 12/07/04–12/21/04

SH307 4 2/22/2005 2/12/2005 02/01/05–02/15/05

SH308 4 2/22/2005 1/21/2005 01/15/05–02/01/05

SH309 4 2/22/2005 1/13/2005 01/07/05–01/25/05

SH310 4 2/22/2005 1/16/2005 01/10/05–01/28/05

SH311 4 1/19/2005 1/16/2005 01/16/05–01/19/05

SH312 4 4/10/2005 3/8/2005 03/08/05–03/08/05

SH313 4 4/19/2005 4/1/2005 03/15/05–04/05/05

SH314 5 4/19/2005 4/17/2005 04/16/04–04/19/05

SH315 4 4/19/2005 4/1/2005 03/15/05–04/05/05

SH316 5 5/24/2005 4/15/2005 04/01/05–05/10/05

SH317 5 6/15/2005 5/1/2005 04/10/05–05/30/05

SH318 5 7/13/2005 7/7/2005 07/01/05–07/09/05

SH319 5 7/13/2005 5/15/2005 05/01/05–05/25/05

SH320 5 7/13/2005 7/1/2005 06/26/05–07/06/05

SH321 6 8/19/2005 8/10/2005 08/05/05–08/19/05

SH322 6 8/19/2005 8/15/2005 08/10/05–08/19/05

SH323 6 10/18/2005 9/10/2005 08/10/05–10/19/05

SH324 7 12/15/2005 12/5/2005 11/23/05–12/07/05

SH325 7 02/07/2006 12/20/2005 12/05/05–01/01/06

SH326 7 02/15/2006 1/10/2006 01/01/06–01/20/06

1 Dredge sample SH300 contained 2004 dacite and rock fragments uplifted from crater floor (1980–86 dome or shallow 

conduit rock).

2

 Hand sample SH301 contained several dacite blocks pushed up from crater floor (1980–86 dome or shallow conduit rock).

Table 1. Rock samples collected in the crater of Mount St. Helens from October 2004 to February 2006.

[All samples of new dome are multiple rock fragments collected by helicopter dredging except SH301 and SH325, 

collected by hand. Sample numbering sequence continues that adopted for Mount St. Helens reference collection 

from 1980–86. Sample numbers with hyphenated suffixes (for example, SH304-2G1) elsewhere in the report refer to 

individual fragments in these composite samples (Thornber and others, 2008b). Eruption dates are estimated by using 

collection locations and known lineal rate of extrusion to track samples back to vent, as described in text. Limiting dates 

specify time interval for emplacement, from known field relations.]

On October 11, 2004, the first spine (spine 1) of new lava 

was extruded through uplifted glacier near the south flank of 

the 1980–86 lava dome, and we began collecting lava samples 

soon after. During the remainder of October, spine 2 emerged 

and extended to the south, while a new area of the glacier was 

uplifted east and then southeast of spine 1. Our initial samples 

were collected from spine 1 by helicopter dredging using a 

makeshift bucket dredge, followed by landings and hand sampling on October 27 and November 4. By October 27, protuberances of hot lava had breached the area of uplifted glacier 

in the southeastern part of the crater, and by November 4 an 

elongate wedge of new lava had emerged from the uplift, risen 

to a height of about 100 m (fig. 3), and been thrust laterally 

toward the south crater wall, plowing the glacier aside. Because 

of its elongate, recumbent form, this spine 3 became known as a 

“whaleback,” a term used to describe similar recumbent spines 

during an early stage of the eruption at Soufrière Hills volcano, 

Montserrat (Watts and others, 2002). Initially, the whaleback 

had a smooth east face mantled by soft white fault gouge; this 

face dipped 40o–60o degrees east and was marked by down-
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 651

Spine 1

October 27, 2004

Figure 2. Oblique aerial photograph (view to southeast) of small 

pyroclastic flow (with vapor clouds rising from it) produced by 

rockfall from spine 2 (shrouded in cloud) on October 27, 2004. 

Snowmelt from flow produced watery lahar in small channel on 

snow-covered glacier, forming small debris fan in basin at base of 

spine 1. USGS photo by J.S. Pallister.

plunge striations and darker colored “bathtub rings,” the latter 

recording uplifted positions where the spine had previously 

been in contact with a debris apron at its base. During November and December 2004, global positioning system (GPS) data 

and fixed-camera observations recorded the whaleback moving 

southward in conveyorlike fashion, transporting new lava from 

a vent at its north end and shedding it as hot talus blocks to the 

south and west (Schilling and others, this volume, chap. 8; Vallance and others, this volume, chap. 9).

By mid-December 2004, oblique longitudinal fractures 

developed in spine 3. By early January 2005, spine 3 had 

broken into multiple fracture-bounded segments and spine 4 

began to emerge from the vent. Spine 4 was also extruded to 

the southeast as a whaleback. It overrode rubble from spine 3 

until it too reached the south crater wall and fragmented in mid 

to late April as spine 5 emerged from the vent. Spine 5 overrode the two earliest spines. It began to fragment into rubble 

in late July 2005. Smaller, rubbly spine 6 then emerged, but it 

was extruded to the southwest from the vent and extended farther west than previous spines. During the fall of 2005, a sag 

opened between spines 5 and 6 as spine 6 continued to extrude 

and thrust upward and to the southwest. Finally, during 

October 2005 and continuing into 2006, spine 7 emerged and 

was thrust steeply to the southwest; it overrode spine 6, and 

part of it collapsed to form a talus apron that extended onto 

the narrowing west arm of Crater Glacier, which was trapped 

between the westward-thrusting dome and the crater wall. The 

eruption of spine 7 has continued to the time of this writing 

(September 2006) and has been characterized by near-vertical 

spine growth and collapse, producing a large mass of dacite 

rubble in the western part of the crater.

The relatively low angle of extrusion through most of the 

eruption produced inclined spines that were extruded to the 

south (spines 2–5), such that the northern half of the vent margin was exposed at the surface. This margin consisted of an 

arcuate rim of hot gouge, convex on its north side. The southern margin of the vent has not been exposed. Extrapolation of 

the exposed rim as a circle yields a vent diameter of 100–200 

m; however, variably exposed extensions of the gouge rim 

along the east and west margins of the spines (on occasions 

for several hundred meters) suggest an oval vent outline. This 

oval outline may be the result of the inclined intersection of a 

north-plunging cylindrical conduit with the crater floor, or it 

may indicate a true elongation of the vent.

Each of the spines displayed a carapace of white to tan fault 

gouge or cataclasite 1–2 m thick as it emerged from the vent 

(Cashman and others, this volume, chap. 19). The gouge zone is 

striated, with multiple layers of subparallel slickensides oriented 

parallel to the transport direction at the point of emergence 

from the vent. The orientation of the slickensides, combined 

with photographically or GPS-determined rates of linear spine 

transport, provide a means to “backtrack” sample localities to the 

vent and thereby establish eruption ages. The vent is assumed to 

be at the altitude of the crater floor below Crater Glacier. Using 

this method, approximate eruption dates have been estimated 

for each of the dome samples, along with age ranges, the latter 

reflecting uncertainty in rate of transport from the vent to the 

collection locality (table 1). One sample (SH312) is assigned 

a unique eruption age, coincident with the explosion of March 

8, 2005, because it was melted into a nylon rope on a near-vent 

monitoring station (McChesney and others, this volume, chap. 7; 

LaHusen and others, this volume, chap. 16).

Laboratory Methods

Petrography and Microbeam Analysis

Polished thin sections of representative samples were 

prepared and examined using petrographic and scanning electron microscopes (SEM), electron microprobe, and imaging 
652 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

B

October 27, 2004

November 4, 2004

Spine 

window

Annulus of roof-rock debris

SH304 

dacite

Annulus of roof-rock debris

Weathered gouge

Roof-rock pendant

Fresh gouge

SH303

SH304 

dacite

Annulus of roof-rock debris

Weathered gouge

Roof-rock pendant

Fresh gouge

SH303

SH301

Location of 

inset photo

Figure 3. Sampling sites for new dome rocks early in eruption. A, View to west of lava exposed at 

crest of uplifted glacier in southeast crater on October 27, 2004; USGS photo by J.S. Pallister. Steam 

emanates from surface of spine 3, surrounded by dark (wet) debris from the crater floor and root of 

the 1980–86 dome that was brought to surface atop spine. Inset shows initial window through craterfloor debris, exposing gouge on spine 3 surface. Our sample SH301 was from roof rock collected near 

helicopter (circled). B, View to west of same area on November 4, 2004, showing fault-gouge mantled 

and striated east face of spine 3, now about 100 m above uplifted surface of the deformed glacier. 

USGS photo by J.S. Pallister. First dense dacite sample of the spine interior (SH304 ) was dredged from 

southeast end of spine 3 and was still hot when returned to a staging area 10 km distant. First gouge 

sample was collected during a short-duration landing at SH303. Note roof-rock pendant of crater-floor 

debris, which was transported upward as spine emerged from glacier and rose upward and obliquely to 

the southeast. See Cashman and others (this volume, chap. 19) for additional details of gouge carapace.
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 653

methods. Because of extensive, very fine grained groundmass 

crystallization, modal analyses were done using a combination 

of optical point counting, low- and high-magnification SEM 

image analysis of backscattered electron images, and multielement microprobe-stage raster maps. Thin-section raster map 

data were processed using image analysis software to give area 

percentages for a series of color bins, which were adjusted to 

differentiate phases in the sample and then combined through 

a series of additions, subtractions, and normalizations to yield 

approximate area percentages of the phases. Stage raster maps 

were made for large (1 cm2) and small (2.56 mm2) areas, the 

latter to yield a high-resolution image (2-μm pixel spacing), 

which was required to resolve microlites. For example, the 

color spectrum was adjusted so that red areas in the large-area 

aluminum (Al) map equal the area percentage of plagioclase 

phenocrysts (fig. 4). When added to percent of small green 

lath-shaped microlites in the high-resolution aluminum map, 

this gives the total feldspar abundance. Similarly, red to yellow 

areas of the two iron (Fe) maps (fig. 4) gives the percentage of 

oxides, which is subtracted from the percentage of nonblack 

areas in the high-resolution silicon (Si) map to determine 

percentage of void space. Yellow areas in the high-resolution 

Si map give the percentage of SiO2 phases, which are present 

only in the groundmass. Glass abundance is obtained from 

the potassium (K) maps by subtracting the area percentage of 

K-rich phases (yellow, alkali feldspar) in the high-resolution 

map from the percentage of glass (blue) in the large-area, 

lower resolution map. Using these types of mapping and 

image analysis techniques, it has been possible to quantify the 

abundance and textures of crystals, glass, and vesicles at scales 

ranging from centimeters to microns (table 2).6

Matrix glass, glass inclusions, oxide minerals, and 

selected silicate minerals were analyzed using the electron 

microprobe, and the results are summarized in tables 3 and 

4 and in appendix 1 (included in the digital version of this 

paper). Most microprobe analyses were conducted at 15-kV 

accelerating voltage and 20-nA beam current, using the JEOL 

8900 microprobes at U.S. Geological Survey (USGS) laboratories in Denver, Colo., and Menlo Park, Calif. SEM analyses 

were conducted using a JEOL 5800LV SEM equipped with an 

energy-dispersive X-ray system (EDS) at the USGS microbeam laboratory in Denver. Natural and synthetic silicate 

standards were used and off-peak background corrections were 

applied to standards and unknowns; secondary silicate-mineral 

standards were used to verify and monitor the analytical 

routines. Glass standard NIST SRM 610 (Ihinger and others, 

1994; Pearce and others, 1997) was used as a secondary standard for sulfur (S) and chlorine (Cl).

Beam spot sizes were varied according to the material 

being analyzed and the nature of the problem being examined. 

Point beams were used for oxides and to examine zoning 

profiles, larger spot sizes (to 40 μm) were used where possible 

to analyze glasses, and iterative counting techniques, similar 

to those pioneered by Nielson and Sigurdsson (1981), were 

used during all glass analyses to monitor and, where indicated, 

to correct for alkali and silica migration. To avoid loss related 

to electron beam damage, the elements sodium (Na), K, and Si 

were analyzed before other elements in glasses, and count rates 

were collected for these elements starting immediately when 

the beam was unblanked. Curves were individually fit to the 

count-rate data for each analysis, and the intercept at time zero 

was used in the data reduction routines. Long count times (>300 

seconds on and off peak) were used for S and Cl analyses, 

yielding detection limits of 70–90 ppm for S and about 50 ppm 

for Cl. On the basis of long-term reproducibility of standard 

analyses, analytical reproducibility for microprobe analyses 

is estimated as 1–2 percent of the reported amounts for major 

elements, except for H2O, which is calculated by the difference 

of complete major- and minor-element probe analyses from 100 

percent and is subject to uncertainties of about 1 weight percent.

In addition, two sets of approximately 700–1,000 microprobe point analyses each were done in polygonal grids and 

averaged to characterize the bulk composition of two small 

quenched inclusions in two samples. A third set of gridded point analyses were done over a similar-size grid in the 

host dacite of one of the samples to test the method. To save 

time, mean atomic number background corrections and short 

(approximately 20 s) dwell times were used for these analyses, 

cutting analytical time to approximately 36 hours for each 

area. The resulting data for the dacite host were also used to 

calculate an analytical mode by binning analyses into mineral 

types, glass, and void space according to a series of compositional and stoichiometric rules (see tables 2 and 5).

IR Spectroscopic Analysis

Dissolved water and carbon dioxide in two unusual 

glassy samples, SH304-2G1 and SH304-2G2, were determined by Fourier-transform infrared (FTIR) spectroscopy 

in the Department of Earth and Planetary Sciences at the 

American Museum of Natural History, using a Nicolet 20SXB 

FTIR spectrometer attached to a Spectra Tech IR Plan microscope, according to methods described by Mandeville and 

others (2002). A total of 10 to 16 spots were measured in the 

most microlite-free (Fe-Ti oxides and orthopyroxene) regions 

of matrix glass in each sample, and IR results are reported in 

table 6. Because of the small thicknesses of films of glass in 

the matrix of typical microcrystalline dacite, it has only been 

possible to determine water contents by FTIR on the matrix 

glass in the one glass-rich sample (SH304-2G).

Bulk-Rock Analysis

Representative samples of the dacite lava were coarsely 

crushed in a hydraulic press and hand-picked for bulk-rock 

analyses. Splits of the crushed dacite were cleaned of xenolithic 

material, and the cleaned samples were submitted for analysis at 

USGS labs in Denver, Colo. Splits of many of the samples were 

also analyzed at Washington State University’s Geoanalytical 

6

 Tables 2 through 9 are grouped at the back of this chapter, after 

“References Cited.”
654 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Fe MAP, 1 x 1 CENTIMETERS Fe MAP, 1.6 x 1.6 MILLLIMETERS

Al MAP, 1 x 1 CENTIMETERS Al MAP, 1.6 x 1.6 MILLIMETERS

20.7

1.1

7.7

14.2

57.1

3.2

21.2

39.1

Fe CONCENTRATION, IN PERCENT

Al CONCENTRATION, IN PERCENTAl CONCENTRATION, IN PERCENT

Fe CONCENTRATION, IN PERCENT

99.4

5.8

37.0

63.2

59.0

3.5

22.0

40.3

Figure 4. Electron microprobe-stage raster maps of sample SH305-1A illustrating the procedure for deriving abundance 

of phases from image analysis. Each map was produced by driving microprobe stage over areas approximately 1 cm2

(left images) and then 2.6 mm2

 (right images; to differentiate matrix phases) while simultaneously recording X-ray peak 

intensities from five wavelength-dispersive spectrometers and from backscattered electron detector. The X-ray intensities 

are compared to standard intensities to derive approximate concentrations of major elements as oxides and are plotted 

according to stage position to produce images shown here. Note that each map has a unique color scale.

Laboratory. Major-element analyses of whole-rock samples 

were obtained by X-ray fluorescence (XRF) methods (Taggart and others, 1987). Trace-element abundances for selected 

samples were determined by Induction Coupled Plasma–Mass 

Spectrometric (ICP–MS) and Instrumental Neutron Activation 

Analysis (INAA) methods (Baedecker and McKown, 1987). On 

the basis of replicate analyses of standards, analytical reproducibility of major-element abundances is estimated to be better 

than 0.4 percent of the reported values for SiO2 and Al2O3. 

For other elements, reproducibility is better than ±2 percent of 

values in the range 1–10 wt. percent and better than ±6 percent 

for abundances less than 1 wt. percent. Coefficients of variation 

(100 × (σ / x), in percent) for INAA analyses, based on counting statistics, are given in table 8.

Modal Compositions and Textures of 

the Dacite Lava

Microcrystalline Dacite

The spines erupted from Mount St. Helens in 2004–6 are 

composed of remarkably crystal-rich hornblende-hypersthene 

dacite. For consistency with past usage at Mount St. Helens, 

the term hornblende is used informally herein for rock names. 

In fact, the amphiboles in the 2004–6 dacite range in composition from magnesiohornblende and tschermakite to magnesiohastingsite (Rutherford and Devine, this volume, chap. 31; 

Thornber and others, this volume, chap. 32), and therefore 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 655

Si MAP, 1 x 1 CENTIMETERS Si MAP, 1.6 x 1.6 MILLIMETERS

8.6

52.0

95.4

18.3

38.0

57.6

77.2

Si CONCENTRATION, IN PERCENTSi CONCENTRATION, IN PERCENT

K MAP, 1 x 1 CENTIMETERS K MAP, 1.6 x 1.6 MILLIMETERS

2.2

14.0

25.6

37.3

K CONCENTRATION, IN PERCENT

1.4

4.4

7.3

10.2

K CONCENTRATION, IN PERCENT

Figure 4.—Continued.

we use the more general term amphibole when referring to 

individual crystals. Microbeam mapping of three microcrystalline dacite samples (SH304-2A, SH305-1A, SH315-4) yields 

total crystal contents that range from 69 percent to 87 percent 

(phenocrysts+microlites) on a void-free basis. Phenocryst 

content of the microcrystalline dacite averages 46.0 percent 

(±4.3 percent), also on a void-free basis (table 2). The high 

total crystal content is a result of extensive groundmass crystallization, which produced fine-grained mosaics of microlites, 

residual glass, and vesicles. For consistency with previous work 

(Cashman, 1992) a cutoff of approximately 30 μm is used to 

distinguish microlites from phenocrysts in modes. The microlite-rich groundmass texture of the 2004–6 dacite is referred 

to as “microcrystalline” to distinguish it from the more typical 

hyalopilitic textures of the 1980s lavas.

Samples collected from spine 3 during the early stages 

of the eruption (table 1), when the eruption rate was high 

(that is, greater than 5 m3/s; Schilling and others, this volume, 

chap. 8; LaHusen and others, this volume, chap. 16), and from 

the exterior margin of the spines are relatively glassy. For 

example, interior sample SH304-2A has, by microbeam map 

analysis, approximately 45 percent phenocrysts of plagioclase, 

amphibole, hypersthene, and Fe-Ti oxides set in a microcrystal-rich matrix that contains approximately 33 percent microlites and 22 percent glass on a vesicle-free basis (table 2). 

Exterior sample SH305-1A is also relatively vesicular (37 per-
656 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

cent) and glassy (30 percent on a vesicle-free basis). However, 

samples erupted subsequently have less glass. For example, 

sample SH315-4 from spine 4 (erupted approximately April 1, 

2005) has only about 13 percent glass (table 2), and samples 

collected later in 2005 and in 2006 have only trace amounts 

of glass. The presence of two populations of vesicles in the 

microcrystalline dacite, one with larger (100–500 μm long) 

irregular vesicles, and a second with small (1 μm) spherical 

vesicles implies continued degassing during microlite crystallization (figs. 5A, 5B).

A detailed look at SH305-1A shows several features 

that are characteristic of the 2004–6 dacite. First, plagioclase 

phenocrysts commonly have overgrowth rims that are partly 

irregular in form and appear to have grown rapidly (fig. 5A). 

Amphibole crystals are complexly zoned and have thin, 

fine-grained breakdown rims (fig. 5C) that average 5 μm in 

thickness (Thornber and others, this volume, chap. 32). The 

groundmass contains abundant plagioclase microphenocrysts 

and microlites in a glassy matrix (figs. 5B, 5D). The matrix 

of this sample shows the onset of very late stage cotectic 

precipitation of Na- and K-rich feldspar and quartz, which 

commonly nucleate at the margins of larger plagioclase crystals. Additionally, there is a minor component of tridymite 

(acicular dark-gray crystals in fig. 5D), as also reported by 

Blundy and Cashman (2001) for slowly cooled samples of 

the March 1983 Mount St. Helens dacite. 

Later erupted samples preserve noticeably less glass 

in the matrix. For example, sample SH315-4 from spine 4 

(erupted approximately April 1, 2005) has only about 13 

percent glass (table 2; fig. 6A). The groundmass has some 

still-distinguishable plagioclase microlites, but these have been 

largely subsumed into a near-holocrystalline matrix (fig. 6B). 

Also common in this sample are fairly large, discrete crystals 

of quartz (fig. 6B). Sample SH316 from spine 5 shows similar 

features to samples from spines 3 and 4, with evidence of 

rapid rim growth on many plagioclase phenocrysts (figs. 6C) 

and small patches of preserved glass in a near-holocrystalline 

matrix (fig. 6D). Sample SH315 also contains broadly folded, 

vesicular, pink-colored veins a few millimeters to a few centimeters in cross section and extending for tens of centimeters 

across large blocks (fig. 7A). Scanning electron micrograph 

images from within one of the veins display an irregular 

network of matrix phases between the phenocrysts, formed by 

finely intergrown tridymite, quartz, Na- and K-rich feldspar, 

and void space (fig. 7B). Where larger cavities are transected 

by the plane of the thin section, the silica-feldspar intergrowth 

is seen to extend into the vesicle space (fig. 7C), exposing 

delicate networks of tridymite and feldspar (figs. 7D, 7E). 

Magnetite grains in the vein zones have undergone oxidation 

exsolution (fig. 7F). These features document a continuum of 

matrix crystallization, extending from a fluid-saturated magma 

into a vapor-dominated phase of crystallization.

Samples collected later in 2005 and in 2006 have only 

trace amounts of glass and show even more extensive groundmass and vapor-phase crystallization. An examination of 

SH321 and SH322 from spine 6 (samples erupted on approximately August 10 and August 15, 2005, respectively) reveals 

groundmasses composed entirely of microlites (figs. 8A–D) 

and consisting of graphic intergrowths of a “granite minimum” 

assemblage of sodic plagioclase (An20–30), quartz, tridymite, and a more potassic feldspar (anorthoclase to sanidine). 

Quartz and the feldspars occur as subhedral crystals within the 

intergrowths. Tridymite occurs within the graphic intergrowths 

and protrudes into vesicle space (fig. 8A), attesting to crystallization from both melt and gas phases. Microprobe analyses 

of representative groundmass phases are given in table 3, and 

many are keyed to figure 8.

Cristobalite is a common groundmass phase in samples from each of the spines. It displays a characteristic 

microbotryoidal texture and commonly fills vesicles and 

interstices in the dacite, as well as within rare diktytaxitic 

quenched inclusions. The occurrence is indicative of latestage vapor-phase crystallization, similar to that described 

for cristobalite in the 1980 Mount St. Helens cryptodome 

(Hoblitt and Harmon, 1993).

Glassy Dacite

The dredge haul from spine 3 on November 4, 2004 

(SH304 sample series, Thornber and others, 2008b), contained 

principally blocks of still-hot, pink to gray, microcrystalline dacite. However, the haul also contained several small 

(2–5-cm diameter) dark glassy dacite fragments. Most of 

these have bulk compositions, phenocryst assemblages, Fe-Ti 

oxide temperatures, and textures that indicate they are derived 

from uplifted parts of the 1980–86 lava dome. However, one 

flow-banded glassy fragment (SH304-2G) has characteristics 

that suggest it may be juvenile (for example, overlapping 

Fe-Ti oxide temperature and similar bulk composition compared to the microcrystalline dacite). Sample SH304-2G is a 

hornblende-hypersthene dacite with 72 percent glass (fig. 9, 

table 2). The ratio of K2O in the glass compared to that of the 

whole rock is indicative of a glass content of only 54 percent 

and suggests that the sample is enriched in glass and depleted 

in crystal components (table 2), such as would occur by flow 

segregation or filter-pressing, an effect that is consistent with 

the flow-banded character of the fragment (fig. 9A). Although 

there is some uncertainty in establishing that this sample is 

juvenile, it is investigated in more detail here as a possible 

analogue for the melt phase of the 2004 magma before the 

extensive groundmass crystallization. Unlike the amphibole 

phenocrysts in the microcrystalline dacite, which have reaction-rim thicknesses that cluster at about 5 μm (Rutherford and 

Devine, this volume, chap. 31; Thornber and others, this volume, chap. 32), the amphibole phenocrysts in SH304-2G lack 

reaction rims (fig. 9B), indicating more rapid ascent, as might 

be expected during the earliest phase of the eruption. The glass 

in this sample contains swarms of very small (less than 1 μm 

diameter) hypersthene microlites (figs. 9C, 9D), similar to 

those found in the outer zones of plagioclase phenocrysts in 

the 2004–6 dacite (Streck and others, this volume, chap. 34).
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 657

A

C

B

D

100 μm 100 μm

50 μm 25 μm

Figure 5. SEM backscattered electron images (A, B, D) and one optical image (C) of microcrystalline dacite sample 

SH305-1. This early 2004 sample (eruption date, November 20, 2004, table 1) has relatively glass-rich groundmass and 

shows early stages of shallow cotectic groundmass crystallization. Magnification varies; scale bars labeled on images. 

A, Plagioclase phenocryst. Note irregular, cuspate margins, indicative of rapid growth. B, Groundmass. Visible in A and B 

are two size populations of vesicles: large (irregular) and smaller (oval to spherical). C, Zoned amphibole with thin (5 μm) 

reaction rim. D, Small plates of tridymite (dark gray).

Inclusions

Inclusions in volcanic rocks provide glimpses of country 

rock surrounding magma reservoirs and conduits, plutonic 

equivalents or cumulates from host magmas, quenched 

blebs of other magmas, and potentially even partially melted 

source rocks for the magmas (Bacon, 1986; Costa and others, 

2002; Jackson, 1968; Smith, 2000). Mount St. Helens dacites 

contain a rich suite of inclusions, consisting primarily of 

dioritic to gabbroic plutonic rocks, and the 2004–6 lava dome 

is no exception.

Plutonic and Metamorphic Inclusions

The 2004–6 dacite contains relatively abundant 

medium-grained gabbroic to dioritic inclusions as well as 

minor banded and granoblastic amphibolite inclusions. 

They are present in both dome lava and gouge samples; we 

estimate their abundance at 1–5 percent. Although not yet 

studied in detail, in hand sample the inclusions appear to 

be similar to the gabbroic and metamorphic inclusions that 

make up 3–5 percent of the 1980–86 lava dome (Heliker, 

1995; Pallister and others, 1991). We previously thought that 
658 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

A

C

B

D

250 μm

100 μm 50 μm

50 μm

Figure 6. SEM backscattered electron images of microcrystalline dacite samples emplaced in April 2005. These samples 

have experienced rapid and extensive groundmass crystallization, producing a nearly holocrystalline groundmass. 

Magnification varies; scale bars labeled on images. A and B, sample SH315-4, erupted April 1, 2005. In B note dark-gray 

quartz grain (center) and micrographic intergrowth of feldspar and quartz forming the groundmass. C and D, Sample SH316, 

erupted April 15, 2005.

some of the gabbroic inclusions represented cumulates from 

the Mount St. Helens magma reservoir, but U-Pb ion-probe 

zircon ages of about 25 Ma for several samples (our unpublished ion microprobe data) suggest that many of these may 

instead be derived from a Miocene gabbroic to quartz diorite 

intrusive complex inferred from gravity data to underlie the 

volcano (Williams and others, 1987).

Dacite Inclusions

Some samples collected from the later spines of the 

2004–6 eruption contain small (1–5 cm) inclusions of dacite 

with slightly different textures or degrees of alteration compared to the host dacite. Petrographic examination and electron microprobe analyses of one of these inclusions in sample 

SH321-1C show little difference in mineralogy or composition across the inclusion boundary. Another (SH316-1A) has 

a bulk composition distinct from the host dacite, with 63.4 percent SiO2, and it likely represents a xenolith entrained in the 

conduit magma and derived from the root of the 1980–86 lava 

dome. We suspect that many such “dacite-in-dacite” inclusions 

are fragments of shallow conduit lava and wall rocks, incorporated as the conduit shifted slightly to feed subsequent spines. 

Consistent with this idea, an examination of inclusions in 

the gouge carapace and variations in bulk composition of the 

gouge (Cashman and others, this volume, chap. 19) indicate 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 659

B

C

A

B 500 μm

9 mm

Figure 7. Photograph and SEM 

backscattered electron (BSE) 

images of microcrystalline dacite 

sample SH315-5. Sequence 

shows progressive increase in 

magnification, and locations are 

mapped on preceding images. 

A, Polished thin section, 4.5 cm 

in long dimension, showing pinkcolored vesicular vein. B and C, 

BSE images showing network of 

fine-grained microlite-rich matrix 

between phenocrysts. D and E, 

BSE images showing phases 

identified by morphology and 

energy-dispersive spectrometric 

(EDS) analysis: Fsp = Na+K-rich 

feldspar; Mt = titanomagnetite; Opx = 

orthopyroxene (hypersthene). F, BSE 

image of titanomagnetite grains in 

the vein zone showing development 

of oxidation-exsolution lamellae.

that the conduit margins incorporated a range of crater floor 

rocks as the gouge formed.

Quenched Magmatic Inclusions

Two small (<1-cm diameter) inclusions, composed of 

diktytaxitic networks of plagioclase and amphibole crystals 

and containing variable amounts of glass and void space, were 

found in samples SH315-4 (figs. 10A, 10B) and SH321-1C. As 

magma mixing has been involved in the evolution of magmas erupted at Mount St. Helens during the past 4,000 years 

(Pallister and others, 1992; Smith and Leeman, 1993) and has 

been implicated as a trigger for explosive eruptions (Pallister 

and others, 1996), we searched the 2004–6 rock collection 

for additional examples, and we examined these two inclusions in detail. Our search indicates that although dacite-indacite inclusions are relatively common, quenched inclusions 

are exceedingly rare in the suite of 2004–6 samples. In the 

approximately 500 kg of samples collected to date (Thornber and others, 2008b), we have found only the two small 

quenched inclusions, with a total mass less than 5 g.
660 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

C

D

E

D

SiO2 (Tridymite)

SiO2 (Tridymite)

Fsp

Fsp

Fsp

Opx

Fsp

Fsp

Mt

Mt

20 μm

200 μm

E

F

SiO2 (Tridymite)

100 μm

10 μm

Figure 7.—Continued. 

A more-detailed examination of an inclusion in sample 

SH315-4 reveals elongate and hopper-shaped crystals of amphibole and plagioclase (figs. 10C, 10D), as would be produced by 

rapid growth during undercooling of a hotter and more mafic 

magma in the dacite host magma. The amphiboles are high in 

Al2O3 (to 14 percent) and they lack reaction rims, except at the 

margins of the inclusion, where they reacted with the host dacite 

during decompression to produce thin breakdown rims (Thornber and others, this volume, chap. 32). Plagioclase phenocrysts 

have normally zoned An60–40 cores bounded by cuspate dissolution zones, which are rimmed by much more sodic and potassic 

plagioclase (to An20, Or9). These rims extend into the rhyolitic 

groundmass glass (75–76 percent SiO2) as angular and hopperform crystals (figs. 10D, 10E). Cristobalite, with a characteristic 

microbotryoidal texture, is relatively abundant as a late vaporphase product.

To determine the bulk compositions of the inclusions, we 

used a microprobe to analyze 736 and 960 points distributed 

over the polished surface of each of the two inclusions and 

1,021 points over a similar area of the host dacite (fig. 10A; 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 661

SH321-1C-7b SH321-1C-7d

SH322-1-b SH322-1-d

A

C

B

D

10

17 19

20

18

9

3

2

4

4

5

2

2 1 5

4

3

5

50 μm

10 μm

10 μm

10 μm

Figure 8. SEM backscattered electron images of microcrystalline dacite samples, documenting continued matrix 

crystallization. Microprobe analysis points keyed by number to table 3 (final suffix numbers in sample-number list of table 

3 refer to locations shown here). A and B, Sample SH322-1. C and D, Sample SH321-1C. Minerals probed include tridymite 

with approximately 1 percent Al2O3 (A, 9–10; B, 2), quartz with less than 0.7 percent Al2O3 (B, 3; C, 4; D, 4–5), An19–39 /Or3–9

plagioclase microlites (B, 4; C, 2, D, 1–3) and An33–36 

/Or1–4 plagioclase microphenocryst (A, 17–20), anorthoclase (B, 5), and 

cristobalite with more than 3 percent Al2O3 (C, 5).
662 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

B

A

am

pl hy

hy

am

1 mm

100 μm

300 μm

100 μm

C

D

Figure 9. Photomicrographs (A, C, D ) and SEM backscattered image (B) of glassy fragment SH304-2G1, collected on November 

4, 2004. A, Phenocrysts of plagioclase (pl), hypersthene (hy), and amphibole (am) lie within a flow-foliated groundmass marked by 

dark bands bearing micron-size oxides. B, Microlites of hypersthene and oxides appear as bright specks in backscattered image, 

and subhedral to euhedral amphibole microphenocryst lacks reaction rims. C and enlargement D, Glass-rich matrix is peppered with 

submicron-size curvilinear hypersthene microlites and oxide grains.

table 5). Averages of these analyses reveal that the inclusion 

in sample SH315-4 is andesitic (60.0 percent SiO2), consistent 

with the relatively high An content of plagioclase cores. It is 

regarded as a bleb of a more mafic magma that mingled with 

the dacite at some point in the past (see under “Discussion”). 

The second inclusion, in SH321-1C, has a bulk composition that overlaps with the host dacite, at 64.8 percent SiO2

(table 5). However, this sample has a similar quench texture 

and phenocryst assemblage to the inclusion in SH315-4. It 

probably represents an entrained bleb of hotter dacite that 

quenched in the cooler host dacite. However, we cannot rule 

out the possibility that formerly it was a more mafic magma 

that quenched in the dacite and then accumulated late-stage 

rhyolite melt in its matrix through filter pressing (to account 

for its higher silica content). Quenched inclusions are often 

heterogeneous in composition because of the effects of postentrainment crystallization (Bacon, 1986). During crystallization 

of a more mafic inclusion in a cooler and more silicic melt, the 

melt reaches vapor saturation; it vesiculates, and pressure in 

the inclusion increases. This process may force residual melt 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 663

A

4 mm

Figure 10. Backscattered electron SEM stage raster-map images of sample SH315-4. 

Abbreviations: am, amphibole; cr, cristobalite; pl, plagioclase; gl, glass; vo, void space. A, Composite 

of 24 smaller images, shows a mafic (60 percent SiO2) quenched inclusion at upper right, with 

acicular amphiboles (light gray) and glassy vesicular groundmass. Irregular box encloses an 

area that was analyzed using a grid pattern to determine inclusion’s bulk composition (table 5). 

B, Enlargement of part of inclusion, showing diktytaxitic texture (visible also in A) and void-filling 

cristobalite (cr) with microbotryoidal texture. C, D, and E, differing views and magnification showing 

cristobalite (in image C ), elongate and hopper crystals of amphibole and plagioclase (C,D,E ), brown 

high-silica rhyolite glass (D, E ) and irregular dissolution and overgrowth textures in plagioclase (E ). 

Numbers in images D and E refer to An content of plagioclase and percent SiO2 in glass (the latter 

determined by broad-beam microprobe analysis).

toward the margins of the inclusion as the interior gains void 

space (vesicles). Thus, inclusions may autofractionate, some 

parts gaining high-silica residual melt and some parts losing 

residual melt.

Compositions of Matrix Phases and 

Oxides

Silicate Minerals

Matrix silicate minerals include hypersthene, plagioclase, 

anorthoclase, quartz, tridymite, and cristobalite. The compositions of phenocrysts in the dacite are the subject of other reports 

in this volume (Rutherford and Devine, chap. 31; Thornber and 

others, chap. 32; Streck and others, chap. 34; Kent and others, 

chap. 35). Here we focus only on the compositions of mineral 

grains in the microcrystalline groundmass of samples—grains 

that have textures indicative of late and shallow crystallization. 

These include microlites (grains <30 μm in diameter), matrix 

silica phases, and edge compositions of plagioclase phenocrysts. 

Compositions of these minerals are listed in table 3, and textures 

of the matrix phases are illustrated in figures 5 through 8. 

Cristobalite is a common minor phase in all but the most glassy 

dome rocks. The cristobalite grains have a distinctive botryoidal texture and multiple contraction fractures. They commonly 

occur at the margins of irregular void space, indicating deposition from a late-stage, silica-rich vapor phase. This occurrence 

is similar to that seen in the 1980 cryptodome (Hoblitt and 

Harmon, 1993). Bladed crystals of tridymite are present in low 
664 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

B

C

200 μm

50 μm

am

vo

pl

cr

vo

vo

pl

pl

pl

gl

gl

gl

cr

am

D

E

pl

pl

pl

gl

vo

gl

pl

75.9

vo

pl

54

pl

gl

75.4–76.3

23

39

59 43 60

52

40 μm

50 μm

Figure10.—Continued.

abundance in most samples, although they are common in one 

of the samples, SH309-1. Quartz microlites occur as anhedral 

intergrowths with anorthoclase and sodic plagioclase, in some 

cases forming a patchy texture extending into small glass pools 

at the margins of plagioclase grains, which have experienced 

rapid edge growth. The compositions of plagioclase microlites 

and the final few microns of some phenocryst rims extend to 

An20, well below the average rim composition of An35 measured 

in hundreds of zoning profiles (Rutherford and Devine, this 

volume, chap. 31; Streck and others, this volume, chap. 34) and 

thought to represent the late stages of plagioclase equilibration 

at depth. A second feldspar, which ranges from anorthoclase to 

sanidine in composition, with as much as 43 percent Or component (table 3), is also present in the matrix of some samples, 

where it is intergrown with sodic plagioclase and quartz, forming a “granite minimum” assemblage (Tuttle and Bowen, 1958).

Glasses

Microprobe analyses of matrix glass and glass inclusions 

reveal mostly dry, high-silica rhyolite compositions (table 4). 

Most of our analyses of matrix glasses come from early phases 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 665

0

2

3

4

5

68 70 72 74 76 78 80

SiO2, IN WEIGHT PERCENT

H2O, IN WEIGHT PERCENT BY DIFFERENCE

SH305 MI (PL, AM)

EXPLANATION

SH305 matrix glass

SH304-2G1 matrix glass

SH304 MI (AM)

20 MPa

100 MPa

SIMS data

range for

SH305, 315

H2O IN MELT INCLUSIONS (MI) AND MATRIX GLASS

(DEFOCUSSED BEAM AND CORRECTED POINT BEAM ANALYSES)

SH304-2G1 and SH304-2G2

matrix glass by FTIR

Figure 11. Abundance of H2O in matrix glass and glass inclusions from spine 3, determined by differencing the totals 

of microprobe analyses from 100 percent and plotted against SiO2. Abbreviations: PL, plagioclase; AM, amphibole; MI, 

melt inclusion. Circled fields represent average and standard deviation for Fourier-transform infrared (FTIR) analyses 

of matrix glass in SH304-2G1 and SH304-2G2, which overlap with microprobe analyses of same sample. Concentrations 

of water at saturation in silicate melts at 20 and 30 MPa (Moore and others, 1995) are shown for comparison, as is the 

range of secondary ion mass spectrometer (SIMS) data for matrix glass in samples erupted later (J. Blundy and K. 

Cashman, written commun., 2005).

of spine 3. In subsequent samples, matrix glass occurs only as 

thin films that are difficult or impossible to analyze with the 

electron microprobe without exciting nearby mineral grains. 

Available matrix-glass compositions are all rhyolitic, ranging 

from 73 to 78 percent SiO2 and containing <0.5 to 2.2 wt. percent H2O (fig. 11). Most glass inclusions also have low waterby-difference values. A single inclusion in amphibole from 

SH304 has about 3.6 percent H2O (by difference); all others 

have <3 percent and many have <0.5 percent H2O. Secondary 

ion mass spectrometric (SIMS) analyses of glass inclusions 

in samples SH305 and SH315 range from 0.03 to 0.20 wt. 

percent (Jon Blundy, written commun., 2005), confirming the 

dry character of the inclusions.

Two sections of the glassy fragment SH304-2G were analyzed using FTIR spectroscopy. These measurements of total 

dissolved water in the matrix glass of sample SH304-2G2, 

based on the intensity of the absorbance at 3,570 cm−1, yielded 

1.9±0.18 wt. percent H2O (16 spots, table 6), with a range 

from 1.4 to 2.1 wt. percent, overlapping the microprobe waterby-difference value reported here for the same sample (fig. 

11). Total water in matrix glass of SH304-2G1 is 1.7±0.14 

wt. percent on the basis of summation of water dissolved as 

hydroxyl groups determined from the absorbance peak at 

4,500 cm−1 and molecular water determined from the intensity 

of the absorbance peak at 1,630 cm−1, in agreement with the 

value determined from the 3,570 cm−1 peak. Total dissolved 
666 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

water concentration in the matrix glass of sample SH304-2G2, 

estimated from the measured intensity of the absorbance peak 

at 3,570 cm−1, is 2.3±0.15 wt. percent (10 spots, table 6) and 

ranging from 2.0 to 2.5 wt. percent. Total dissolved water 

in SH304-2G2 matrix glass, determined from summation of 

water dissolved as hydroxyl and molecular components, is 

2.1±0.17 wt. percent, in agreement with the value from the 

3,570 cm−1 peak (table 6).

In matrix glass from samples SH304-2G1 and 

SH304-2G2, the concentration of dissolved water present as 

molecular H2O slightly exceeds that of water dissolved as 

hydroxyl groups (table 6). The observed water speciation in 

SH304-2G matrix glasses is in contrast to that observed in 

rapidly quenched (cooling by hundreds of degrees per minute) 

experimental glasses or in silicate melts at high temperature 

and pressure with similar total dissolved water concentrations (Dixon and others, 1995; Mandeville and others, 2002; 

Nowak and Behrens, 1995; Shen and Keppler, 1995; Silver 

SH305,

matrix glass

SH304,

matrix glass

70 71 72 73 74 75 76 77 78 79 80

0

200

400

600

800

1,000

1,200

1,400

1,600

S or Cl, IN PARTS PER MILLION

SiO2, IN WEIGHT PERCENT

SH304,

inclusions 

EXPLANATION

S (m)

S (m)

S (m)

Cl

Cl (m)

MELT INCLUSIONS MATRIX GLASS S

S (i)

S (i)

Cl (i)

Cl (i)

Cl (i)

S (i)

Cl (m)

Cl (m)

SH304

SH304-2G1

SH305

SH304-2G1

SH305

MSH05WS_1_16

SH304

SH304-2G1

SH305

SH304-2G1

SH305

MSH05WS_1_16

and Stolper, 1989; Sowerby and Keppler, 1999). This most 

likely reflects species reequilibration during slower cooling 

and magma ascent. In none of the acquired spectra from SH3

04-2G1 and SH304-2G2 matrix glass was a CO2 absorbance 

peak above background observed. Dissolved CO2 concentration is below our detection limit (<20 ppm).

Most matrix glasses and glass inclusions have sulfur 

abundances less than detection limits (70–86 ppm, table 

4). However, four glass inclusions from dacite erupted in 

October 2004 have between about 80 and 130 ppm S, and 

one inclusion contains about 250 ppm S (table 4, fig. 12). 

Chlorine abundances in glass inclusions reach a high of 

1,450 ppm; they average 829±309 ppm in microcrystalline 

sample SH305-1 and 1,155±142 ppm in microcrystalline 

sample SH304-2A9b. Chlorine abundance in matrix glasses 

averages 278±53 ppm in the high-silica rhyolite glass of the 

microcrystalline dacite; in contrast, it averages 852±47 ppm 

in low-silica rhyolite glasses from the flow-banded glassy 

Figure 12. Abundance of sulfur (S) and chlorine (Cl) in matrix glass and glass inclusions from spine 3 plotted against SiO2. 

Note low concentrations for S in matrix glass and most glass inclusions. Only glass inclusions in sample SH304 (eruption 

date, October 18, 2004), and one inclusion from SH305 have more than 200 ppm SO2 (or >100 ppm S). Chlorine abundances in 

glasses show a wider range, with early matrix glass of the glassy fragment SH304-2G1 having more than 800 ppm Cl, whereas 

matrix glass in SH305, erupted later, is depleted, with less than 400 ppm Cl, consistent with more extensive degassing of 

microcrystalline SH305. The matrix glass with the lowest SiO2 is in a fragment from tephra sample MSH05WS_1_16, collected 

from the January 16, 2005, explosion deposit.
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 667

sample SH304-1G and in glassy fragments from the January 

16, 2005, tephra. As the high-silica rhyolite is a differentiation product of decompression-driven matrix crystallization, 

the lower Cl in these glasses is consistent with very shallow 

degassing of Cl during the latest stages of solidification, 

accompanied by development of fracture permeability in the 

conduit (Edmonds and others, this volume, chap. 27).

Kalama

andesite

We-jgWn-jp

Wn-jg Wb -jp

Ps-jg

Ye-jg

Yn-jp

Yn-jg

Cy-jp

Jg-jp

Pu-jg

Mc-jp

T -jg

80-bm

80-mr

80-jp

HM MNO NNO to grain-pair

contact

700 1,000 1,050 1,100

−8

−9

−10

−11

−12

−13

−14

−15

750 800 850 900 950

TEMPERATURE, IN DEGREES CELSIUS

OXYGEN FUGACITY, AS LOG fO2

A

11/04/04 (10/18/04)

850±5°C and −12.29±0.07 fO2 (SH304-2A9)

844±2°C and −12.43±0.07 fO2 (SH304-2G1)

EXPLANATION

Eruption date, with collection date in parentheses, and

analytical details (for example, 11, 94 = 11 pairs, 94 analyses)

01/16/05 (Ash) 11, 94

11/20/04 (1/03/05, SH305-1) 5, 45

11/20/04 (1/03/05, SH305-1) 14, 98

10/18/04 (11/04/04, SH304-2A, -2G) 9, 118

1985-86 (10/20/04) 3, 20

Oxides

Equilibrium temperatures and oxygen fugacities were 

calculated using a spreadsheet routine (Lepage, 2003), which 

yields results from multiple solution models for the Fe-Ti 

exchange thermometer and oxygen barometer (Andersen and 

Lindsay, 1988). Results calculated using the solution model of 

Figure 13. Results of Fe-Ti oxide thermobarometry for 2004–6 lava dome (A, B), plotted with respect to oxygen fugacity and 

temperature. Buffer curves for nickel-nickel oxide (NNO), manganese-manganese oxide (MNO) and hematite-magnetite (HM), 

and results for samples from earlier eruptive episodes at Mount St. Helens plotted for comparison. Samples color coded by 

eruption date. Number of grain pairs of the two solid-solution series ilmenite (FeTiO3)–hematite (Fe2O3) and magnetite (Fe3O4)–

ulvöspinel (Fe2TiO4) for each sample are indicated in the explanation. In cases in which temperatures and oxygen fugacities 

cluster, error bars represent ranges derived from the thermobarometer by using maximum and minimum values of standard 

deviations of Fe and Ti. In cases in which zoning was apparent from analytical profiles, backscattered-electron imagery, or 

stage mapping, temperatures and oxygen fugacities are plotted for each analyzed point in the zoned phase. Arrows show 

trends as contact between the two oxide grains is approached. Most zoned grains show increasing apparent temperature and 

oxygen fugacity toward grain contact, indicative of recent heating, although latest erupted samples analyzed (April 2005) show 

apparent cooling and reduction in oxygen fugacity. A, Data from samples erupted between October 18, 2004, and January 16, 

2004, as well as data from 1980 and previous eruptions at Mount St. Helens. Two-letter symbols such as “Cy” refer to tephra 

units of Mullineaux (1996), suffixes refer to source of data as follows: -jg = Jim Gardner (Gardner and others, 1995a, 1995c),

-mr = Malcolm Rutherford (Rutherford and others, 1985), -bm = Bill Melson (Melson and Hopson, 1981), -jp = John Pallister 

(unpublished). B, Data from samples erupted between January 21 and April 17, 2005. C, Average atomic Mg/Mn ratios for 

ilmenite-magnetite pairs used in generating thermobarometric results presented in A and B. Also shown are equilibrium line and 

error limits (2σ) for oxide pairs from volcanic rocks (Bacon and Hirschmann, 1988).
668 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

HM MNO NNO to grain-pair

contact

TEMPERATURE, IN DEGREES CELSIUS

B

700 1,000 1,050 1,100

−8

−9

−10

−11

−12

−13

−14

−15

750 800 850 900 950

OXYGEN FUGACITY, AS LOG fO2

EXPLANATION

04/17/05 (4/19/05, SH314-2) 7, 61

04/01/05 (4/19/05, SH315-4) 7, 57

03/08/05 (4/10/05, SH312) 6, 93

01/21/05 (2/22/05, SH308-3A) 6, 65 

01/16/05 (Ash) 11, 94

11/20/04 (1/03/05, SH305-1) 5, 45

11/20/04 (1/03/05, SH305-1) 14, 98

10/18/04 (11/04/04, SH304-2A, -2G) 9, 118

1985-86 (10/20/04) 3, 20

Eruption date, with collection date in

 parentheses, and analytical details

 (for example, 11, 94 = 11 pairs, 94 analyses)

0.4

0.5

0.6

0.7

0.8

0.5 0.6 0.7 0.8 0.9 1.0

ILMENITE ATOMIC RATIO OF Mg/Mn, IN LOG UNITS

MAGNETITE ATOMIC RATIO

OF Mg/Mn, IN LOG UNITS

C

Figure 13.—Continued.

Figure 13.—Continued.

Stormer (1983) are most consistent with experimental calibrations and are given in table 9 and displayed in figure 13. Only 

analyses of touching Fe- and Ti-rich oxide pairs were used for 

thermobarometry, and individual analyses of oxide minerals in 

the dacite are given in appendix 1 (in the digital version of this 

chapter). Typically 5 to 25 analyses were made of each grain 

pair. For unzoned grains or for grain pairs that were too small 

to demonstrate zoning, the average compositions of each phase 

are plotted, along with error bars for temperatures and oxygen 

fugacities (fO2

) (figs. 13A, 13B). The error bars were calculated 

from the maximum and minimum of the ±1 standard deviation 

ranges for Ti and Fe in each grain pair. As shown in figure 13C, 

all of the grain pairs used for thermobarometry in this study 

have Mg/Mn ratios within error limits of the equilibrium line of 

Bacon and Hirschmann (1988).

We have observed apparent zoning in both the titanomagnetite and ilmenite with respect to distance from the boundary between these two phases (fig. 14), although the abrupt 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 669

931 928 938 921 907 903 898

A

57

36

12

Ti CONCENTRATION, IN PERCENT

B 58

36

12

Ti CONCENTRATION, IN PERCENT

Figure 14. Electron microprobe stage maps showing 

abundance of TiO2 in oxide grain pairs from samples 

SH305-1 (A) erupted approximately November 20, 2004, and 

MSH05JV_1_19 (B) from a dacite fragment in the ash deposit 

from the January 16, 2005, explosion. Maps are 100 X 100 μm 

in dimension. The SH305-1 grain pair (A) shows a relatively 

uniform or slightly decreasing TiO2 trend across ilmenite grain 

(orange), with a steep increase in abundance within 10 μm of 

the contact with the adjacent titanomagnetite grain (blue) and 

a gradual increase in TiO2 across the titanomagnetite grain 

toward the ilmenite grain. Apparent temperatures, calculated 

for individual analytical points in the titanomagnetite (as 

indicated by the asterisks and numbers in the figure) 

compared to the average composition of the ilmenite show a 

decrease from 928°C at the grain contact to 898°C at the rim. 

The apparent temperature for the two analytical points closest 

to the grain boundary yielded 931°C. The MSH05JV_1_19 grain 

pair (B) displays the same abrupt decrease in Ti in the ilmenite 

with distance from the titanomagnetite, but the opposite trend 

in the titanomagnetite, which shows a gradual decrease in 

TiO2 abundance as the grain boundary is approached. The 

superimposed line graphs show the relative abundance of 

TiO2 in the band outlined by the two horizontal lines in each 

image. The MSH05JV_1_19 grain pair (number 7 in table 9) 

shows a decrease from an average of 830°C for points within 

10 microns of the grain boundary to 938°C for the average 

compositions of the two grain interiors.

increase in Ti in the ilmenite near the boundary has been 

attributed to submicron-scale recrystallization of the ilmenite (Devine and others, 2003; Rutherford and Devine, this 

volume, chap. 31). For grain pairs in which zoning is apparent 

in one phase but not the other, we calculated multiple apparent temperatures and oxygen fugacities using each individual 

analysis of the zoned phase and the average composition of the 

unzoned phase. Temperatures for the first seven months of the 

eruption are plotted against eruption dates and permissive date 

ranges (from table 1) in figure 15.

The early sample of microcrystalline dacite (SH304-2A9; 

eruption date approximately October 18, 2004; table 1) yields 

a temperature of 850±5°C at an oxygen fugacity of 10−12.29 ± 0.07

(table 9), which we take to be the conditions of last equilibration 

of the magma before ascent in 2004. Zoning in oxides erupted 

subsequently yield Fe-Ti oxide thermometer results that range 

to apparent temperatures of 950°C or more before plunging to 

less than 800°C in April 2005. Our samples that were erupted 

since April 2005, after the eruption rate slowed substantially, 

all contain titanomagnetite with oxidation lamellae of titanohematite, such that we have been unable to obtain meaningful 

thermometry results for the later part of the eruption.

Bulk-Rock Geochemistry

No significant variation in average major- and trace-element whole-rock compositions of the 2004–6 dome samples 

is apparent in the major-element data; standard deviations for 

more than 20 samples are close to the range of uncertainty 

expected for replicate analysis of standard rocks (table 7). The 

2004–6 dacite has 64.93±0.09 percent SiO2 and a relatively 

high Na2O/K2O ratio of 3.2, similar to other Mount St. Helens 

dacite. The lack of variation in SiO2 with eruption date is 

evident in figure 16, which also shows SiO2 abundances for 

gouge and dacite collected from the roof of spine 3 (SH300, 
670 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

7

In this report we use an age range of A.D. 1479–1750 for the Kalama 

period, on the basis of a tree-kill date of A.D. 1722 for a summit-dome lahar 

in the Muddy River drainage and a bracketing youngest age of A.D. 1750 

determined for trees that subsequently rooted on a late summit dome-derived 

hot lahar in upper Pine Creek (Yamaguchi and Hoblitt, 1995). The age of final 

summit dome eruptions in the Kalama eruptive period remains uncertain.

SH301), when it first breached glacial ice in the southeastern 

crater on October 25, 2004. These roof-rock samples have 

distinctly lower SiO2 and alkalis and higher MgO, CaO, and 

FeTO3 (total iron expressed as Fe2O3), and their compositions 

overlap with those of lava erupted in 1985–86 (table 7, fig. 

17), suggesting that they are crater-floor rocks uplifted by 

the 2004 dacite as it punched through the crater floor. Analyses of several small glassy fragments that were collected in 

dredge samples from early spines (spines 1 and 3) are also 

listed in table 7. These are higher in silica than samples from 

the 1980–86 dome and may represent fragments of dacite that 

were intruded early in the 2004 eruption.

Compositions of the 2004–6 samples lie along the highSiO2 projections of most major- and minor-element variation 

trends defined by the Goat Rocks and 1980–86 dome lavas (Pallister and others, 1992). However, they lie off the trend for TiO2, 

which is present in higher abundance than would be expected 

from the 1980–86 trend (fig. 17). High-precision analyses of Cr 

by instrumental neutron activation analysis (INAA) are diagnostic of the presence of basaltic mixing components within Mount 

St. Helens andesites and dacites (Pallister and others, 1992). 

Chromium abundances in 2004–6 dome rocks are compared 

to those of Kalama (A.D. 1479–17507

), Goat Rocks (A.D. 

1800–1857), and the 1980–86 eruptive periods in figure 18 (full 

INAA analyses and uncertainties are reported in Thornber and 

others, 2008b). As with TiO2, the 2004–6 dacites have higher 

Cr abundances than would be expected from a continuation of 

the Goat Rocks and 1980–86 trends, as well as lower Cr at 65 

percent SiO2 than the dacites from the early Kalama eruptive 

period. Analyses were conducted at the same laboratory using 

the same procedures and standards to ensure reproducibility, and 

duplicate samples from 1985 and 1986 dome rocks were rerun 

to verify the differences.

A characteristic feature of Mount St. Helens dacites, 

including 2004–6 lava, is the lack of europium anomalies and 

a decrease in total rare-earth-element (REE) abundance with 

increasing SiO2, thereby ruling out progressive crystal fractionation in their origins (table 8, fig. 19). The 2004–6 dacite 

is distinct in having the lowest heavy-REE abundances of any 

of the dacites we have analyzed from eruptive products of the 

past 500 years at Mount St. Helens.

Discussion

Petrology of the 2004–6 dacite at Mount St. Helens 

provides constraints on both shallow and deep magmatic processes and insights into the roles that degassing and crystallization play in controlling the explosivity of volcanic eruptions 

and how the past history of the volcano helps in forecasting 

the future course of the eruption. Accordingly, the discussion is presented in four parts: (1) shallow conduit processes, 

(2) deep reservoir processes, (3) implications for magmatic 

plumbing, explosivity, and eruption triggering, and (4) lessons 

from the past and implications for the future.

Shallow Conduit Processes

Field relations, petrography, volatile contents of glasses, 

and oxide thermobarometry help us understand the processes 

that were operative in the shallow conduit beneath the 2004–6 

vent. Together with geophysical and gas geochemical monitoring data, they show that magma within the conduit solidified 

at shallow levels beneath the vent, producing a nearly solid 

seismogenic plug (Iverson and others, 2006) that has fed the 

series of spines that constitute the new lava dome.

Constraints from Field Geology

The 2004–6 eruption of Mount St. Helens represents a 

nonexplosive end member in the range of eruptive behavior 

of dacitic magmas. By the spring of 2006 the volume of the 

new dome exceeded 80×106

 m3 (Schilling and others, this 

volume, chap. 8), a volume similar to the 1980–86 lava dome 

(74–77×106 m3, Swanson and Holcomb, 1990). Eruption rates 

during 2004–6 have ranged from as high as 9 m3/s in the initial 

phase to less than 1 m3/s (Schilling and others, this volume, 

chap. 8). The style of eruption, with continued extrusion 

of solid dacite spines, contrasts with most of the 1980–86 

600

650

700

750

800

850

900

950

1,000

1,050

1,100

NOV JAN MAR APR

2004 2005

ERUPTION DATE

OCT DEC FEB

10/18/04

11/20/04

1/16/05

1/21/05

4/1/05

4/17/05

3/8/05

TEMPERATURE, IN DEGREES CELSIUS

Figure 15. Fe-Ti oxide temperatures for 2004–5 Mount St. Helens 

dacite lava samples plotted against eruption dates; diamonds 

indicate best estimate of eruption dates. Bars show full range 

of permissible eruption dates (table 1). Wide range in apparent 

temperatures results from zoned oxide pairs (fig. 14).
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 671

Figure 16. Content of SiO2

in Mount St. Helens samples 

plotted against likely eruption 

date (table 1), except for SH300 

and SH301, which are samples 

of crater-floor rocks and plotted 

against collection date. Note 

uniform abundance of SiO2 (65 

percent) in most dome and 

gouge samples. Exceptions 

include low abundances (<63 

percent) in inclusion samples 

and in two gouge samples, 

and wide range in crater floor 

samples collected on October 

20, 2004.

Figure 17. Variation diagrams of MgO 

and TiO2 against silica, showing trends in 

juvenile eruptive products from 1980–86 

eruptions at Mount St. Helens (Pallister and 

others, 1992) and 2004–6 eruptions. Note that 

2004–6 samples lie at high-SiO2 projection 

of the linear arrays, except for TiO2, which 

is distinctly higher for given SiO2 content in 

the 2004–6 lavas than in previous eruptive 

episodes.

Lava

EXPLANATION

Gouge

Uplifted crater-floor rock

SIO2, IN WEIGHT PERCENT

58

OCT DEC JAN

2004 2005 2005 2006

MAR MAY JUL SEP DEC JAN

59

60

61

62

63

64

65

66

67

68

ERUPTION DATE

SH321-1C dacite inclusion

(probe) 

SH316-1A dacite inclusion

(XRF) 

SH315-4 andesite inclusion

(probe) 
672 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Q

Q Q

Q O

OO

R

R

P

S

RRR

R

R

R R

T

T

T

T

T SS

S

SSS

S

S S

S P

U U U

VV W

7

7

7

W

XX YY

MMM

M

MM

M

M M

M

M

H

H H

H

I I I J J J

F E

DC A7

2

2 2

2

b

b

c

c c a

1 1

1

1

11

1111

1

1111

5 9 9

W 0

V

50%

40%

30%

Mixing trend to MSH OIB-type basalt

(49% SiO2, 175 ppm Cr)

20%

10%

N

N

N

A.D. 1500 mix

Goat Rocks—A.D. 1800–1857 (lowercase letters)

Kalama cycle—A.D. 1479–1750 (symbols are

 uppercase letters, with alphabetical arrangement

 matching time sequence)

1980–86 cycle (numerals)

~A.D.1750

57

90

80

70

60

50

40

30

20

10

0

100

58 59 60 61 62 63 64 65 66 67

2004–6 dome

SiO2, IN WEIGHT PERCENT, VOLATILE FREE

Cr, IN PARTS PER MILLION, BY INAA

Mount St. Helens magma-mixing 

trends of the past 500 years

EXPLANATION

A.D. 1650

A.D. 1479

a–c

A–Y

1–9

K A L A M A

C Y C L E

1980–86 and GOAT ROCKS

Figure 18. Variation diagram of Cr against SiO2 for juvenile eruptive products from Mount St. Helens erupted 

during past 500 years (Pallister and others, 1992). Note that 2004–6 samples have higher Cr abundance than 

the projection of the Goat Rocks eruptive period and 1980–86 trend but lower abundance than the early phase 

of the Kalama eruptive period (A.D. 1479–1750). Analytical uncertainty for Cr is ±3 ppm at the 10-ppm level (by 

instrumental neutron activation analysis, INAA) and for SiO2 (XRF data) is ±0.2 percent on the basis of counting 

statistics and replicate analyses. Magma mixing with basaltic magmas at Mount St. Helens is called on to 

explain the cyclic trends (Pallister and others, 1992). Letters and numbers refer to stratigraphic sequence, and 

percentages refer to proportions of a basaltic mixing end member as explained in Pallister and others (1992).

dome-forming eruptions at Mount St. Helens, as well as with 

most other dome-forming eruptions worldwide. The development of gouge-mantled spines is unusual but not unique. 

Photographs of the Showa-Shinzan lava dome (extruded in 

a Japanese wheat field on the flank of Mount Usu in 1944–

45) show spines with striations that look remarkably like 

those on the Mount St. Helens dome (Mimatsu, 1995). The 

200-m-diameter cylindrical and striated spine of Mont Pelée, 

Martinique, is another example. The original Peléean spine 

rose at rates of 3–6 m per day and reached a height of 115 m 

above the crater floor by the time of the explosive eruptions 

on May 6–8, 1902 (Chrétien and Brousse, 1989; Smith and 

Roobol, 1990). A larger spine grew in 1903, rose to 300 m, 

and bore vertical striations and slickensides (Williams, 1932, 

quoting Lacroix, 1904). Similarly, the andesitic Soufrière Hills 

volcano, Montserrat, has produced multiple, gouge-covered 

and striated spines during some phases of dome construction 

(Watts and others, 2002). What is unusual about the 2004–6 

Mount St. Helens eruption is continuous spine extrusion over 

such a long period, nearly continuous extrusion of relatively 

thick gouge and cataclasite at the spine margins, and the low 

level of explosivity that has accompanied spine growth and 

disintegration. Of known analogs, the gouge-covered spines 

of Showa-Shinzan are considered most similar to those at 

Mount St. Helens (Vallance and others, this volume, chap. 9).

The 2004–6 spine-forming mode of eruption also contrasts with eruptions that produced the 1980–86 lava dome 

at Mont St. Helens. The 17 eruptive episodes in the 1980s 

produced 20 short (200–400 m), thick (20 m) flows (lobes) and 

included periods of endogenous growth and small explosions 

(Swanson and others, 1987; Swanson and Holcomb, 1990). 

Lava flows produced through mid-February 1984 had scoriaceous carapaces, as opposed to the gouge carapaces on spines 

of the current eruption. Two small spines were extruded during 

these episodes (in late February–March 1983 and during May 

1985), but these also lacked gouge carapaces. Also in contrast 

with the hotter dome lavas of the 1980s, the 2004–6 lava generally lacks augite phenocrysts and is much more crystal rich 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 673 CHONDRITE-NORMALIZED ABUNDANCE 

~2,000-yr-old OIB-type

MSH basalt

Andesite from

middle Kalama

eruptive period

(n=3, n=7) Dacite from

early and late

Kalama eruptive period

(n=24, n=13)

100

2004–6 

dome dacite

(n=11)

1986 dome

dacite (n=3)

1980–85 

dome dacite

(n=7)

1

La Ce Nd Sm Eu Gd

RARE-EARTH ELEMENTS

Tb Ho Tm Yb Lu

10

Figure 19. Chondrite-normalized rare-earth-element (REE) abundance patterns for juvenile eruptive 

products from Mount St. Helens erupted during the past 500 years. Note that the 2004–6 lavas have 

the lowest heavy-REE abundances. Error bars (±1σ) are given for analyses of 1980 and subsequent 

samples. Error bars (±1σ) for older, higher-REE abundance samples are smaller than the symbols and 

are not plotted.

and glass poor compared to the more glass-rich augite-bearing 

hornblende-hypersthene dacite of the 1980s dome (Cashman, 

1992; Geschwind and Rutherford, 1995; Melson, 1983). We 

suggest that a primary control on the low level of explosivity of 

the 2004–6 lava is the combined effect of lower gas abundance 

in the magma and a greater degree of shallow matrix crystallization, which produced virtually dry and rheologically solid dacite 

lava hundreds of meters beneath the vent.

The “megaspines” of the early (1995–98) eruptive phase 

of the Soufrière Hills volcano, Montserrat, also had striated 

surfaces and cataclastic textures, judging from observations by 

one of us (Cashman). Because their morphologies are correlated with eruption rate and character of seismicity (Watts 

and others, 2002), it is instructive to compare these with the 

Mount St. Helens examples. At Montserrat, vertical spines and 

whaleback structures dominated at rates of <1 m3/s, and their 

emplacement was accompanied by periodic hybrid earthquake 

swarms. Megaspines were emplaced aseismically at 1–2 

m3

/s. Broad shear-lobe spines were emplaced aseismically 

at 2–5 m3/s, but their collapse was accompanied by intense 

hybrid swarms. Blocky shear-lobes were emplaced at 2–5 

m3

/s, accompanied by repetitive hybrid swarms and tremor. 

Pancake lobes were emplaced at 7–9 m3/s, also with repetitive 

hybrid swarms and tremor. Explosions at Montserrat occurred 

at eruption rates greater than 9 m3/s and commonly followed 

large dome-collapse events.

The Mount St. Helens spines are similar to the vertical 

spines, whaleback structures, and megaspines of Montserrat. However, instead of being restricted to eruption rates 

of less than 2 m3/s, these morphologies at Mount St. Helens have occurred over the full range of eruption rates, and 

collapse events have been small—rockfalls and sparse rock 

avalanches that have not triggered substantial pyroclastic 

flows. These differences are likely related to compositional 

differences (andesite at Soufrière Hills, dacite at Mount St. 

Helens), and to a much larger vent diameter of 100–200 m 

at Mount St. Helens, compared to about 30 m at Soufrière 

Hills. The larger vent diameter allows slower ascent rates and 

more extensive shallow crystallization for similar volumetric 

rates of extrusion.

Compared with activity at Soufrière Hills, Montserrat, 

collapse events at Mount St. Helens have been minor. Lim-
674 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Mi

Mm

Gm

Gi

Ash

Gouge

Ab

Qz

Or

100 MPa

0.1 MPa

50 MPa

200 MPa

300 MPa

EXPLANATION

Figure 20. Projection of glass compositions from SH304 and 

SH305 Mount St. Helens samples onto modified quartz-albiteorthoclase haplogranite ternary diagram of Blundy and Cashman 

(2001). Cotectic lines and water-saturated minima and eutectics 

are plotted as function of pressure for comparison. Dashed line is 

approximate position of the 1-atmosphere cotectic. Solid symbols 

refer to melt inclusions, open symbols to matrix melt. Other labels 

as follows: Mi, melt inclusion in dacite lava; Mm, matrix melt in 

dacite lava; Gi, melt inclusion in dacite fragment in gouge; Gm, 

matrix glass in dacite fragment in gouge; Ash, melt inclusion 

within tephra fragment in ash; Gouge, matrix melt within dacite 

fragment in gouge.

ited disintegration of collapsing blocks during short runouts 

onto the low-relief crater floor at Mount St. Helens is likely 

due to a drier composition and the more solid (microlite-rich 

and glass-poor) character of the Mount St. Helens dacite lava 

compared to the more glass- and water-rich matrix glasses of 

the Montserrat andesite lava (Couch and others, 2003; Harford 

and others, 2003). Furthermore, greater topographic relief at 

Soufrière Hills has contributed to greater runout distances.

Constraints from Petrography and Major-Element 

Compositions of Matrix Glasses and Microlites

As shown in figures 4 through 8, the matrix of the 2004–6 

dacite is remarkably rich in microlites, even compared to the 

dome lava of the 1980s (with the exception of spine samples 

from 1983 and 1985, which are also highly crystalline). 

We interpret these matrix textures to result from extensive 

depressurization-driven crystallization during ascent of the 

magma through the conduit, similar to the process envisioned 

for the 1980–86 dome (Blundy and Cashman, 2001; Cashman, 

1988, 1992). In contrast to the episodes of the 1980s, we see 

no evidence for stagnation of magma at multiple levels of the 

conduit. Instead, an analysis of eruptive volumes and likely 

conduit dimensions (discussed below) and of consistency of 

amphibole reaction-rim thicknesses (Rutherford and Devine, 

this volume, chap. 31; Thornber and others, this volume, 

chap. 32) indicates that dacite lava of the current eruption is 

a product of continuous ascent of magma through a conduit 

that taps the top of the deeper reservoir at about 5-km depth. 

We see a continuum in the degree of matrix crystallization, 

from early 2004–early 2005 samples with variable fractions 

of matrix glass to late 2005–6 samples, which have only trace 

amounts of glass remaining. Such a pattern is consistent with 

more rapid ascent and eruption rates during the early phase of 

the eruption.

The crystallization of a granite-minimum microlite 

assemblage of tridymite or quartz, An20–30 plagioclase, and 

anorthoclase (table 4) is consistent with extensive groundmass crystallization. The high-silica rhyolite composition of 

residual glass is indicative of very shallow final crystallization. 

Following the arguments of Blundy and Cashman (2001), the 

presence of quartz and the projection of glass compositions 

onto the modified Qz-Ab-Or ternary diagram of Blundy and 

Cashman (2001) yield crystallization pressures for the most 

evolved matrix glasses between 50 MPa and 0.1 MPa (fig. 20). 

This result indicates crystallization at a depth considerably less 

than 2.2 km, on the basis of a density model (Williams and 

others, 1987, section B–B’) in which the crustal column consists of 1.5 km with density of 2.15 g/cm3

 overlying an extensive diorite to gabbroic pluton with average crustal density of 

2.7 g/cm3

. The presence of tridymite in some samples further 

restricts the pressure for matrix crystallization and solidification to the range 11–25 MPa (depth 0.5–1.0 km) at temperature 885–915°C (Blundy and Cashman, 2001), consistent with 

the shallow location of earthquakes that have accompanied the 

eruption (Thelen and others, this volume, chap. 4).

Constraints from SO2, Cl, and H2O Abundances in 

Glasses

The abundance of H2O in glasses can also be used to 

constrain depths, owing to the variation in solubility of water 

in silicate melts with pressure (Moore and others, 1995, 1998). 

Matrix and inclusion glasses show a decline in water content 

from approximately 2.3 wt. percent H2O at 73.5 percent SiO2

to less than 0.1 percent at 77 percent SiO2 (fig. 11). The upper 

end of this H2O range is indicative of quenching at a pressure 

of about 30 MPa (depth of 1.4 km) followed by decompression-driven crystallization and quenching of residual melt, 

which continued to pressures of less than 10 MPa (depth about 

0.5 km). Most of the glass inclusions analyzed so far have 

water and SO2 contents that overlap with the matrix glasses, 

indicating that most of the inclusions have leaked (fig. 11).

Total dissolved water concentrations determined in the 

matrix glass of samples SH304-2G1 and SH304-2G2 were 

used to constrain the pressure and depth at which residual melt 

in this glassy sample solidified. Water solubility in the melt 

at 850°C was computed from electron microprobe determinations of the matrix-glass bulk composition and the 0–300-MPa 

water solubility model of Moore and others (1998). Given 

saturation, the total water concentration in matrix glass of 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 675

sample SH304-2G2, which has the most microlite-free matrix 

glass and the greatest water content (2.3±0.15 wt. percent), 

indicates quench to glass from approximately 33 to 41 MPa, 

or approximately 1.5 to 2 km depth, on the basis of the density 

distribution in crustal section B–B’ of Williams and others (1987). Assuming this unusual glassy sample represents 

a quenched equivalent of the microcrystalline 2004 dacite, 

these results would indicate incipient matrix crystallization 

at a depth of 1.5–2 km, although, as noted above, most of the 

groundmass crystallization and solidification is constrained by 

the microlite phase assemblage in the more typical microcrystalline dacite to depths of less than 1 km.

The abundance of sulfur in glass inclusions (table 4, 

fig. 12) is lower than that in glass inclusions from the 1980 

Mount St. Helens dacite (Blundy and Cashman, 2005; Devine 

and others, 1984; Gardner and others, 1995a; Melson, 1983). 

This indicates that the magma was depleted in sulfur before 

eruption and is consistent with the hypothesis that the 2004–6 

magma was residual from the reservoir of the 1980s. The 

highest sulfur contents (as well as highest Cl and H2O) are 

found in the earliest erupted samples, coincident with the 

highest SO2 and CO2 emissions (Gerlach and others, this volume, chap. 26). This suggests more rapid ascent, less matrix 

crystallization, and less gas loss as the initial magma made its 

way relatively rapidly to the surface.

Abundance of Cl in matrix glass is lower than that in 

glass inclusions (table 4, fig. 12) and decreases with increasing SiO2 and K2O (indicators of increasing matrix crystallization), consistent with degassing accompanying decompression crystallization at shallow levels, as also seen at 

Soufrière Hills volcano (Edmonds and others, 2001; Harford 

and others, 2003). The abundance of Cl in glass inclusions 

and matrix melts can be used to calculate emission rates of 

Cl for the eruption. We use our average Cl abundance of 

1,155±142 ppm in glass inclusions as representative of the 

dacite at magma reservoir depths (4–5 km) (Rutherford and 

Devine, this volume, chap. 31), 852±47 ppm for the magma 

at a depth of 1.4 km (as inferred from the water content of 

the glass in SH304-2G1), and 278±53 ppm for the residual 

Cl in the microcrystalline dacite lava (table 4). The average 

Cl value for glass inclusions (1,155±142 ppm) is based on 

inclusions in amphibole crystals (a Cl-bearing phase) from 

sample SH305-2A9b. Although exchange of Cl with the host 

crystal is a possibility, the average abundance is similar to 

the higher Cl abundances in plagioclase and hypersthene 

inclusions, which range widely from 400 to 1,450 ppm; the 

lower abundances likely result from leakage. The differences 

between the abundance in the glass inclusions and that in 

the matrix glasses, multiplied by the dacite density (approximately 2,400 kg/m3

) yields emissions of 400±249 metric tons 

per million cubic meters of dacite for the ascent interval from 

5 to 1.4 km depth and 758±132 metric tons for the interval 

from 1.4 km to the surface. The total emission is 1,158±381 

metric tons per million cubic meters of dacite extruded, 

or about 100×103 metric tons for the 85×106 m3 extruded 

through August 2006.

On the basis of this analysis, we would predict Cl emissions of 100±33 and 700±231 metric tons per day for extrusion rates of 1 m3

/s and 7 m3/s, respectively, exclusive of any 

contribution from a separate fluid phase. This value is greater 

than the rate of 12–25 metric tons per day measured on August 

31, 2005, by open-path FTIR (Edmonds and others, this 

volume, chap. 27). The discrepancy may be due to (1) precipitation of chlorides in dome rocks (yellow efflorescence has 

been common near fumaroles and attributed to deposition of 

iron chlorides), (2) shallow scrubbing of Cl into groundwater 

(however, only 5.2 tons per day of dissolved Cl was measured 

in water from Loowit and Step springs at the crater mouth, and 

no spike in Cl abundance has yet been detected (Bergfeld and 

others, this volume, chap. 25)), or (3) a lower Cl abundance in 

the August 2005 dacite compared to that from late 2004, when 

our glass inclusion-bearing samples were erupted. We consider 

the last possibility most likely.

Constraints from Oxide Thermobarometry

Zoned oxides similar to those in the 2004–6 dacite are 

known to occur in volcanic rocks that have undergone recent 

heating due to magma mingling, and the time scales for 

preservation of zoning in oxide minerals have been determined experimentally to be on the order of a few days to a few 

months, depending on the size of the oxide grains and degree 

of heating (Devine and others, 2003; Gardner and others, 

1995a; Nakamura, 1995; Pallister and others, 1996; Venezky 

and Rutherford, 1999). Although there is clear petrographic 

and geochemical evidence for magma mixing in older rocks 

from Mount St. Helens (Pallister and others, 1992; Smith and 

Leeman, 1993), there is little such evidence in the 2004–6 dacite, which is uniform in bulk-rock composition and has only 

extremely rare and small (less than 1-cm diameter) quenched 

mafic inclusions.

In several cases, the source of the heating to produce 

zoned oxides is evident, such as in the case of the June 7–12, 

1991, andesites at Mount Pinatubo, which show clear petrographic and geochemical evidence of derivation by mingling 

of basalt with dacite in proportions of about 1:2 (Pallister 

and others, 1996). In other cases, the physical evidence of 

heating by a hotter and more mafic magma consists only of a 

small proportion of quenched mafic inclusions, such as in the 

andesite of Soufrière Hills volcano (Murphy and others, 2000). 

In the Soufrière Hills case, the zoned oxides indicate apparent 

heating of about 30°C, which led Devine and others (2003) to 

propose a model in which underplated basalt heated a boundary zone in an andesitic reservoir. The boundary layer was 

then mobilized and erupted, with only minor entrainment of 

the underlying basalt, as seen in ubiquitous blade-shaped laths 

of pargasite in the groundmass of all samples.

In their study of microlite formation in the Soufrière Hills 

andesite, Couch and others (2003) calculated that a temperature rise of as much as 45°C would be expected from latent 

heat evolved during the extensive (32 percent), shallow (less 
676 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

than 2 km), and relatively rapid (1–2 day) decompressiondriven groundmass crystallization, given the measured extrusion rates and estimated conduit diameter (Melnik and Sparks, 

2002). However, neither Couch and others (2003) nor Devine 

and others (2003) concluded that latent heating at shallow 

levels was responsible for the zoned oxides, because it was 

not possible to experimentally produce the zoning profiles of 

the Soufrière Hills titanomagnetites in short-duration (2 day) 

heating experiments without extensive amphibole breakdown, 

which was not seen in natural samples. However, longer-duration (2 week) heating experiments at magma reservoir pressure 

of 130 MPa (5–6 km depth) did reproduce the natural zoning 

profiles without amphibole breakdown.

The zoned oxides in the Mount St. Helens samples 

share some features with those from Soufrière Hills. First, 

the earliest samples from October 2004, when eruption rates 

were high (greater than 5 m3/s) have oxide pairs that lack 

zonation. Second, the Mount St. Helens oxide pairs erupted 

during the winter of 2004–5 are zoned in a sense that indicates heating (fig. 14). As was also seen in a sample from the 

explosive eruption of September 29, 1997, at Soufrière Hills 

(Devine and others, 2003), the zoning in the winter 2004–5 

oxides is best developed at the grain boundary between the 

titanomagnetite and ilmenite phases, with little zonation 

developed at melt-crystal boundaries. As is evident from the 

range of temperature-fO2

 trends for grain pairs in figure 13, 

the Mount St. Helens thermobarometric results are parallel 

to the common buffer curves, a characteristic of heating with 

melt present (Devine and others, 2003). The above features 

could be interpreted as indicating a similar process of heating 

by hot magma underplating. However, there is considerable 

variation in degree of zoning from grain pair to grain pair 

in the Mount St. Helens samples, even within the same thin 

section. These features indicate that heating was of relatively 

short duration and was variable at the scale of millimeters. 

That the zoning is best developed at the two-crystal grain 

boundaries indicates redistribution of Ti and Fe within the 

crystal structures of titanomagnetite and ilmenite, but with 

little, if any, redistribution between the oxide minerals and 

the adjacent groundmass glass.

We believe that these small-scale relations are best 

explained by latent heating during groundmass crystallization, 

as was also suggested for 1980–86 pumice and lava samples 

(Blundy and others, 2006). Variation of heating at the millimeter scale is most readily explained by inhomogeneity in rate and 

extent of groundmass crystallization and localized variability 

in transport of heat by vapor expansion, conduction, and access 

to the hydrothermal system at shallow conduit levels. The 

fact that zoning is developed with respect to distance from the 

titanomagnetite-magnetite phase boundary, and not at grainmelt boundaries, is probably a result of the rapidly changing 

composition and crystallinity of the adjacent groundmass, such 

that consistent zoning of Fe-Ti oxides by diffusive exchange 

with melt was not possible. In contrast to the situation at Soufrière Hills, where transport from 2 km to the surface is thought 

to have been too rapid for latent heating, by the beginning of 

2005 the eruption rate at Mount St. Helens had slowed to less 

than 2.5 m3

/s (Schilling and others, this volume, chap. 8) and 

the conduit at shallow levels was of much larger diameter—100 

to 200 m, as opposed to about 30 m for Soufrière Hills (Melnik 

and Sparks, 2002). Consequently, the ascent time through the 

shallow interval of groundmass crystallization was longer at 

Mount St. Helens. Using our estimate that most groundmass 

crystallization took place mainly in the uppermost 500 m, with 

a near-surface conduit diameter of 200 m and a volumetric 

ascent rate of 2.5 m3/s, we derive an ascent rate of 7 m/d, which 

is consistent with the measured rates of linear extrusion of 

spines (Dzurisin and others, this volume, chap. 14; LaHusen 

and others, this volume, chap. 16), and an ascent time of about 

2.5 months. Allowing the conduit to decrease in diameter with 

depth, such that it averages only 100 m in diameter over the 

uppermost 0.5–1 km, would yield an average rate of 28–56 m/d 

and an ascent time of 2.6–5.2 weeks. These ascent times are in 

the appropriate range for development and preservation of zoning in oxides (Devine and others, 2003).

We can estimate the maximum amount of latent heating 

that is theoretically possible from the amount of groundmass 

crystallization in the 2004–6 dacite using thermodynamic 

data, much as was done by Couch and others (2003) for the 

Soufrière Hills andesite. Latent heat is equal to the enthalpy 

of melting divided by the heat capacity at the temperature of 

interest. As the microlite population is dominated by plagioclase, we can use the range of enthalpies of melting of 

plagioclase, from 59,280 J/mol for albite to 81,000 J/mol for 

anorthite, to approximate the amount of latent heat available 

and the maximum temperature rise. At an initial temperature 

of about 850°C, the heat capacity of albite would be about 318 

J/mol per K; and of anorthite, about 328 J/mol per K. Therefore, isenthalpic crystallization of 30–40 percent of the dacite 

would produce a temperature rise of about 30–45°C. A more 

complete analysis, by Mark Ghiorso using the thermodynamic 

modeling program pMELTS (Ghiorso and others, 2002) for 

decompressing a water-saturated 2004–6 dacite composition 

at 855°C and fO2

 = 10−12 from 20 MPa to 0.1 MPa, estimates 

a temperature rise of 30°C and results in a magma with 0.5 g 

ilmenite, 2.5 g magnetite, 7.6 g orthopyroxene, 13 g quartz, 

84 g of An31 plagioclase, and 12.3 g of 76.6 percent SiO2 melt 

with 80 ppm H2O (Mark Ghiorso, written commun., 2005)—a 

reasonable match to the natural assemblage. Loss of some of 

the heat would result in appearance of alkali feldspar in the 

assemblage, an increase in the abundance of quartz, and a 

decrease of the melt fraction remaining. Further discussion of 

the role of latent heating in decompression crystallization of 

volcanic rocks is available in Blundy and others (2006).

The maximum quantity of latent heat available from 

groundmass crystallization is less than the degree of apparent 

heating recorded by the zoned oxides in the dome rocks. We 

attribute this discrepancy to the fact that the zoned oxides can 

only give apparent temperatures. They indicate heating, but the 

extent of heating must be less than the maximum indicated by 

the zoning. If the diffusion rates were fast enough, the distribution of Ti and Fe in the grains would rehomogenize to a value 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 677

intermediate between the minimum and maximum ranges seen 

in the zoned crystals and would also reequilibrate with the 

melt; consequently, the actual temperatures would be less than 

the maximum apparent temperatures in figure 13.

The analyses from samples collected later in 2005 and 

in 2006 offer additional insights into the process responsible 

for the oxide zoning (fig. 13B, table 9). Samples erupted 

during April 2005, when the extrusion rate was less than 2 

m3/s, show a wide range in apparent temperature and oxygen 

fugacity. Seven grain pairs in sample SH314-2 are unzoned 

or weakly zoned; three of these yield average temperatures 

of 794–814oC, two of 840–874oC, and two grain pairs yield a 

wide range of temperatures (726–821oC). The titanomagnetite 

of one grain pair appears to be reversely zoned (cooled), with 

lower Ti near the boundary with ilmenite, yielding an apparent 

temperature for the two core areas of about 872oC but decreasing to 840oC at the boundary.

Sample SH315-4 shows an even wider and cooler range 

in apparent temperatures, with reverse zoning of both titanomagnetite and ilmenite and apparent temperatures ranging 

from a low of 644oC to a high of 839oC. Sample SH315-4 is 

also distinct from most of the other samples with respect to a 

lower oxygen fugacity, extending below the NNO buffer curve 

at temperatures below 800oC. As previously noted, we have 

been unable to obtain meaningful oxide thermobarometry 

results for samples erupted after April 2005 owing to the presence of extensive oxidation lamellae.

The lower apparent oxidation state of April 2005 samples 

may be the result of extensive degassing driving precipitation 

of oxides and extraction of Fe+3 from the melt, thereby lowering fO2

. In addition, as the liquid gains more SiO2, the increased 

concentrations of Na2O and K2O in the melt would complex 

with Fe+3 and lower its activity, having the same effect of 

lowering the fO2. This explanation is consistent with a predicted 

decrease in oxidation accompanying extensive crystallization 

at less than 10 MPa pressure by pMELTS modeling (Mark 

Ghiorso, written commun., 2005).

The progression in time from (1) tightly clustered temperatures and oxygen fugacities for the early erupted samples to (2) 

zoned oxides indicative of heating to (3) wide-ranging apparent 

temperatures, reverse zoning, and low oxygen fugacities to (4) 

development of oxidation lamellae is coincident with a decrease 

in eruption rate from more than 5 m3

/s to less than 2 m3/s and 

with a decrease in the abundance of glass from as much as 

30 percent to less than 10 percent. Consequently, we attribute 

the changes in the oxides to result from latent heating due to 

groundmass crystallization, coupled with, and then dominated 

by, heterogeneous heat loss and vapor transport, and finally by 

oxidation at the shallowest levels of the conduit and vent.

To answer questions about the source of the explosions 

of January 16 and March 8, 2005 (Moran and others, this 

volume, chap. 6; Rowe and others, this volume, chap. 29), we 

analyzed oxide grain pairs in 11 small tephra fragments from 

the January 16, 2005, tephra deposit (table 9). All but two of 

these grain pairs are unzoned, and they each produce tightly 

clustered temperatures and oxygen fugacities ranging from 

780oC at fO2

 of 10−13.8 to 930oC at fO2of 10−11.3. None of these 

overlap with the temperature and oxygen fugacity of the initial 

2004 dacite samples; however, two of the grain-pairs have 

zoned titanomagnetite indicative of heating. With grain size 

about 100 μm, they are larger than the range of 20–40 μm that 

is typical of the 2004–6 dacite. From these data, we conclude 

that most of the dacite tephra fragments in the January 16 

tephra are not juvenile, suggesting that this explosion emanated from the margins of the conduit and favoring a phreatic 

rather than magmatic explosion source. However, the story 

may be different for the March 8, 2005, explosion. The analyses of oxide grain pairs from that ash result in temperatures 

that average 852±22oC, indistinguishable from the initial 2004 

dacite temperature of 850±5oC.

Shallow Conduit Processes—A Petrologic Synthesis

In figure 21 we summarize petrologic features that 

explain how and where the conduit magma solidified to 

become a plug. Viscous and relatively melt-rich dacite magma 

with approximately 55 percent melt ascended from depth, 

decompressed, lost volatiles, and began to undergo rapid 

groundmass crystallization at a depth of about 1 km. The earliest erupted dacite preserved equilibrium Fe-Ti oxide temperatures of about 850oC, which we assume is the temperature at 

the top of the source reservoir. With increasing crystallization 

and decreasing melt fractionation, plagioclase became more 

sodic (reaching An20–25 at final stages), and the crystallizing assemblage was joined by anorthoclase and quartz or 

tridymite (table 3) at depths of less than 0.5 km. Latent heat 

that was evolved during decompression-driven groundmass 

crystallization raised temperatures above 850oC until magma 

approached the surface. During the earliest stages of the eruption, when extrusion rates were high, the residual high-silica 

rhyolite matrix melt quenched through the glass transition, 

especially near the more rapidly cooling conduit margins, 

and locally preserved as much as 30 percent glass. However, 

within the spine interior and during later stages of the eruption 

when eruption rates were slower, groundmass crystallization 

continued virtually to completion, producing microcrystalline groundmasses with a granite-minimum composition, and 

latent heating produced zoned oxides. The petrologic constraints outlined here helped define the pressure and depth of 

solidification used in the quantitative dynamical model of Iverson and others (2006) and Iverson (this volume, chap. 21) for 

seismogenic stick-slip extrusion of the solid Mount St. Helens 

dacite plug.

Deep Reservoir Processes

Constraints from Petrography, Textures, and 

Fe-Ti Oxide Thermobarometry

Isotopic data elsewhere in this volume (Kent and others, chap. 35; Cooper and Donnelly, chap. 36; Reagan and 
678 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

others, chap. 37) indicate a complex history for the 2004–6 

magma. For example, U-Th-series phenocryst model ages vary 

over time scales of decades to thousands of years, and some 

crystals have variable initial Sr isotopic ratios. The dacite 

contains amphiboles with widely varying Al and Fe content, 

which indicate derivation over wide ranges in temperature and 

depth—most readily explained by crystallization during convection of magma over a depth range of about 5 to more than 

12 km within the source reservoir (Rutherford and Devine, 

this volume, chap. 31). In addition, normal and reverse zoning 

of plagioclase phenocrysts and entrainment of hypersthene 

microlites in the outer zones of plagioclase resulted from 

changes in temperature and pressure as magma cycled within 

a convecting reservoir (Streck and others, this volume, chap. 

34). It seems clear that many of the phenocrysts were recycled 

from previous crystallization episodes and from crystal mush 

resident in the reservoir.

Modal data and textures described previously (table 2; 

figs. 4, 5), in combination with isotopic, phase-equilibria, 

and phenocryst relations described above, indicate multiple 

stages in the history of the dacite. The phenocryst assemblage of oxides, amphibole, hypersthene, and plagioclase 

grew slowly in the deep convecting reservoir. Our earliest 

samples of the 2004 dacite, erupted in October, provide our 

best constraints on temperature, oxygen fugacity, PH2

O, and 

depth of magma at the roof of the deep reservoir (850 ±5°C, 

fO2 = 10−12.4 ± 0.1, 130 MPa, and about 5 km depth, respectively; 

see also Rutherford and Devine (this volume, chap. 31)). 

The equilibration depth for 2004 is significantly shallower 

than the equilibration depth for May 18, 1980, dacite, which 

by the same method yields a depth of 8.6±1 km (220±30 

MPa) (Rutherford, 1993). These differences suggest that 

the roof of the reservoir became shallower after the 1980 

eruption, a relation also suggested by shallowing of storage areas inferred from water contents of melt inclusions 

in pumice from eruptions in the summer of 1980 (Cashman 

and McConnell, 2005). A greater abundance of phenocrysts 

and more sodic plagioclase rims in the 2004–6 dacite, along 

with shallower equilibration depths and more evolved bulk 

composition, is consistent with this eruption tapping a shallower and more phenocryst-rich part of the magma reservoir 

than was tapped in 1980–86 (Cashman, 1992; Cashman and 

Taggart, 1983; Geschwind and Rutherford, 1995; Rutherford 

and others, 1985).

Glass-rich dacite sample SH304-2G initially had all 

major phenocryst phases, namely plagioclase, hypersthene, 

V E N T

P E R I M E T E R

V E N T

P E R I M E T E R

Magma degasses

and crystallizes

Solidified

 spine

 extrudes

Vent

<1 km Gouge develops

Gouge develops

Figure 21. Diagram showing 

stages in development of a solid 

spine, drawn approximately to scale, 

using oblique aerial photograph 

of spine 4, February 22, 2005. 

View in photograph is to southsouthwest from above east margin 

of 1980–86 lava dome; USGS photo 

by D. Dzurisin. South crater wall 

is visible in background. Spine 

is inclined and is being extruded 

parallel to slickenside striations, 

which are formed in meter-thick 

gouge carapace. Blocky talus 

flanks spine on east and west and 

was derived from areas where 

dome is disintegrating. Bracket 

on diagram indicates depth range 

of solidification, estimated to be 

less than 1 km. Colors of arrows 

diagrammatically represent intervals 

of decompression crystallization, 

latent heating, and subsequent 

cooling at the surface.
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 679

amphibole, and Fe-Ti oxides, coexisting with melt at 

844±2oC and an oxygen fugacity of 10−12.43±0.07, overlapping 

with the temperature and oxygen fugacity of the microcrystalline dacite SH304-2A9 (table 9). The presence of tiny 

elongate hypersthene and oxides (less than 5 μm in diameter; 

fig. 9D) as the only microlite phases is therefore surprising. 

At pressures less than about 100 MPa, amphibole is unstable, 

and one would expect to see plagioclase on the liquidus and 

crystallizing in the groundmass. In contrast, at high temperatures, the Mount St. Helens dacite composition does 

have hypersthene first on the liquidus (Rutherford, 1993; 

Rutherford and Devine, this volume, chap. 31). Could the 

glassy sample SH304-2G represent magma that was heated 

during convection in the deep reservoir? Heating would 

preferentially melt the groundmass phases, such that subsequent cooling and rapid ascent would lead to nucleation of 

early hypersthene. However, late-stage heating is ruled out 

by tight clustering of Fe-Ti oxide thermobarometric data at 

low temperatures in SH304-2G (844±2oC for 65 analyses 

of six equilibrium grain pairs) and overlap of these oxide 

temperatures with those in the microcrystalline dacite, SH3

04-2A9 (850±5oC). In addition, absence of resorbed margins 

on plagioclase and amphibole phenocrysts argues against an 

earlier stage of extensive heating. Consequently, we are left 

with the interpretation that the absence of plagioclase microlites is a kinetic effect and that SH304-2G is a fragment from 

the selvage of a vanguard dike that was emplaced to shallow 

levels and quenched during the earliest phase of the eruption 

in September 2004.

The presence of extremely rare mafic inclusions in two 

of the samples indicates mingling and quenching of andesite 

magma in the dacite at some time during its history. The lack of 

decompression rims on the high-Al amphiboles of the quenched 

inclusions, except where they were exposed to the dacite host 

magma, is consistent with quenching at depths where amphibole 

was stable (>100 MPa). The difference in bulk composition 

between inclusions in SH315-4 and SH321-1C, and their small 

sizes, suggests that these are fragments of larger inclusions, 

which may display a range in vesiculation and filter-pressing 

textures (Bacon, 1986). How long small fragments of quenched 

inclusions can be preserved in convecting, crystal-rich dacite 

magma is unknown. Their scarcity, coupled with the remarkably uniform composition of the 2004–6 dacite, suggest that 

they date from a past mingling event. As discussed later, geophysical and gas data suggest that no new magma was added 

to the reservoir between the summer of 1980 and late 1987 

(Moran, 1994; Moran and others, this volume, chap. 2). However, it is possible that new magma was added during seismic 

swarms that took place during the period after 1987. Consequently, we suggest that these rare inclusions may have been 

entrained during a period of mingling with hotter and more 

mafic magma at depth preceding the 1980 eruption or during 

the repressurization of the reservoir recorded by seismicity 

between 1987 and 1998 (Moran, 1994).

Constraints from Bulk-Rock Geochemistry

Upper regions of magma reservoirs, by virtue of lower 

temperature and higher viscosity, may be bypassed during 

eruptions (Sparks and others, 1984), such that hot, lower 

viscosity plumes of magma from lower levels of a reservoir 

may intrude through overlying magma to be erupted first, 

only to be followed by higher viscosity magma entrained into 

the newly established conduit. Such a model was proposed 

to explain the cycle of initially decreasing and then increasing SiO2 in 1980–86 dome lava at Mount St. Helens (Carey 

and others, 1990; Pallister and others, 1992). However, the 

difference in minor and trace-element compositions of the 

2004–6 and 1980–86 dacites, including lower heavy rare-earth 

elements (HREE) (fig. 19) and higher Ti and Cr (figs. 17, 

18), indicate that this magma was not derived by continued 

closed-system crystallization of 1980 dacite magma. Instead, 

we suggest that the 2004–6 magma was derived from a distinct 

batch of dacite, as explained below.

The rare-earth-element (REE) abundances in Mount St. 

Helens rocks are especially informative because they decrease 

in abundance with increasing SiO2. Owing to their overall 

incompatibility in major phenocryst phases, the decrease in 

their abundance cannot be explained by major-phase crystal 

fractionation (Pallister and others, 1992; Smith and Leeman, 

1987). The highest REE abundances in recent Mount St. 

Helens magmas are in mafic magmas (basalts and basaltic 

andesites) erupted during the Castle Creek eruptive period, 

about 2,000 years ago (fig. 19). This factor, as well as other 

geochemical and petrologic trends in Mount St. Helens 

magmas, are most readily explained by variable amounts of 

mixing between dacite magma (derived by partial melting of 

lower crustal metabasaltic rocks) and mafic magma enriched 

in high-field-strength elements (HFSE) and REE (Gardner 

and others, 1995b; Pallister and others, 1992; Smith and Leeman, 1987, 1993). The 2004–6 dacites have the lowest HREE 

abundances seen in any Mount St. Helens rocks erupted during 

the past 500 years. Consequently, they appear to be the least 

contaminated by mixing with REE- and HFSE-enriched basaltic magmas. Like other Mount St. Helens dacites, the 2004–6 

samples have several characteristics that have been attributed 

to adakites (Defant, 1993; Drummond and Defant, 1990), such 

as high SiO2, Al2O3, and Na2O, low Y and Yb, and relatively 

high Sr (about 400 ppm), and they lack or have only small 

Eu anomalies. However, they are not as enriched in LREE as 

adakites (with La/Yb ratios of about 10, as opposed to ratios 

greater than 20 in the archetypical adakite). The same residual 

phases responsible for depletion of the HREE (garnet, amphibole, and pyroxene) are also expected at the base of the crust, 

a source we prefer, as it does not require transport of the dacite 

through a hot mantle wedge without modification (Dawes and 

others, 1994). Consequently, we do not consider the Mount 

St. Helens dacites as partial melts of slab eclogites but more 

likely as melts of lower crustal metabasaltic rocks that have 
680 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

undergone variable amounts of crystallization and geochemical modification in the shallow crustal reservoir beneath the 

volcano (Smith and Leeman, 1987).

Regardless of the ultimate source of the 2004–6 dacite, 

its distinct TiO2, Cr, and HREE abundances suggest that either 

it is an entirely new batch of dacite magma from the deep 

crust or it is derived from a geochemically isolated region 

of a crustal reservoir. The question of new versus residual 

magma has implications for the long-term eruptive behavior 

of Mount St. Helens, because arrival of a new batch of dacitic 

magma from the deep crust could herald the beginning of a 

new long-term cycle of eruptive activity (Pallister and others, 

2005). Despite the heterogeneity in phenocryst ages and zoning history, the erupted magma is remarkably homogeneous in 

bulk composition, indicating that this batch of magma is well 

mixed. Although we cannot rule out the possibility that a batch 

of dacite magma from the deep crust is involved in the current 

eruption, the low levels of gas emissions (Gerlach and others, 

this volume, chap. 26), U-series evidence for multiple ages of 

crystals (Cooper and others, this volume, chap. 36), and the 

wide range in crystallization depths of phenocrysts, indicative of convection (Rutherford and Devine, this volume, chap. 

31), lead us to favor the hypothesis that the 2004–6 magma 

is derived from a geochemically distinct batch of magma that 

had accumulated recently at the apex of the crustal magma 

reservoir. This magma had been depleted in SO2 and CO2 during multiple cycles of reservoir convection, as explained in the 

section below on eruption triggering.

An Updated Reservoir and Conduit Model

In figure 22 we update a diagrammatic cross section of 

the Mount St. Helens magmatic system that was drawn on the 

basis of seismic, petrologic, and deformation data (Pallister 

and others, 1992; Scandone and Malone, 1985). The principal 

changes to the diagram that make it different from a 1980s 

interpretation are listed here.

1. The top of the reservoir is raised from a depth of 

about 6 km to about 5 km on the basis of the equilibration 

pressure of 130 MPa for the 2004–6 magma (Rutherford and Devine, this volume, chap. 31). Assuming 

that magma ascent was accompanied by an increase in 

pressure in the reservoir and not accomplished simply by 

collapse of the roof into the magma, this top-of-reservoir 

level is bracketed by seismic and deformation data to the 

time interval 1987–97, as described below.

2. The geometry of the conduit from 5 km to the surface is poorly constrained. It may consist of a complex of 

dikes and irregular intrusions that represent the conduit 

pathways of previous eruptions (suggested by textural and 

petrologic evidence that the 1980–86 dome lavas experienced a multistage decompression history; see Cashman, 

1992). However, because of the distinct and uniform dacite composition, consistency of amphibole reaction-rim 

thicknesses, and continuous eruption, the 2004–6 eruption 

is thought to have established a new conduit pathway, 

at least partly distinct from that of the 1980–86 domeforming eruptions. But how was a new conduit emplaced 

without large-scale inflation prior to the eruption? One 

possibility is that the new volume of dacite magma and 

spine that was emplaced at shallow levels between September 23 and October 11, 2004, was compensated by 

extrusion under the glacier of an equal volume of solidified rock from the 1980–86 conduit. Uplift of 12×106

m3 was recorded during this period before the first spine 

emerged (Schilling and others, this volume, chap. 8). Such 

a model would help explain the presence “crater floor” 

samples that are geochemically similar to 1985 or 1986 

dacite, bits of which were uplifted with the glacier and 

sampled on October 20 and 27, 2004 (table 1).

3. We reduce the model size of the reservoir to a magma 

volume of about 4 to 5 km3, a compromise between the 

smaller volumes inferred from seismic data (Lees, 1992; 

Moran, 1994; Musumeci and others, 2002) and magma 

volumes of as much as 8 km3

, which are suggested by eruptive volumes of the largest historical eruptions at Mount St. 

Helens. The largest plinian eruptions of Mount St. Helens 

during the past 4,000 years had magmatic volumes of 4 km3 

(tephra layer Yn) and 2 km3 (tephra layer Wn) (Carey and 

others, 1995). Reservoir volumes are likely at least twice 

this volume, based on the assumption that individual explosive (but noncaldera-forming) eruptions are unlikely to 

withdraw more than half of their shallow crustal reservoir 

volumes without extensive surface deformation or caldera 

collapse. The presence of a continuous magma reservoir 

within the seismically defined depth range between about 5 

and greater than 12 km during the past 4,000 years is based 

on variation in depths of equilibration of magmas erupted 

during this time span, including changes in equilibration 

depths of as much as 6 km in less than 3 years (Gardner 

and others, 1995c). In addition, conductive cooling models 

suggest cooling times on the order of several thousand 

years for magma reservoirs of these depths and volumes 

(see Hawkesworth and others, 2000). That the magma 

reservoir is spatially continuous over its depth range and 

convecting, rather than discontinuous as in Lees’ (1992) 

model, is consistent with the petrologic evidence of largescale convection cited above and with new seismic tomographic modeling, which indicates a continuous reservoir 

over the interval between 6 and at least 10 km (Waite and 

Moran, 2006).

We follow Gardner and others (1995c) in interpreting the 

rapid change in equilibration depths for successive eruptions 

at Mount St. Helens to reflect tapping of different levels of the 

reservoir at different times. Such would appear to be the case 

for the source regions for the 2004–6 magma (130 MPa, about 

5.2 km) compared to the 1980 magma (220 MPa, about 8.6 

km) (Rutherford, 1993; Rutherford and Devine, this volume, 

chap. 31). An increase in seismicity at depths of 3 to 10 km, 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 681

Depleted dacite from

lower crust,

enriched basalt from

mantle wedge

2004–6

2004–6

dike/conduit 

 plug 

1980–86

dike/conduit

2004

source

depth

1980

source

depth

0

1

2

3

4

5

6

7

8

9

10

11

12

13

14

15

16

17

DEPTH BELOW CRATER FLOOR, IN KILOMETERS

Figure 22. Schematic cross section of Mount St. Helens 

plumbing system, updated from Pallister and others (1992) to 

include constraints from current eruption. Yellow dashed lines are 

inferred convection paths in reservoir (Rutherford and Devine, this 

volume, chap. 31) and are overlain by a stagnant crystal-rich mush 

zone considered as the source for the 2004–6 eruption. Diagram 

used here is modified from an illustration published Nov. 28, 2005, 

by the New York Times News Service; copyright © 2005 The New 

York Times Company.

including an initial cluster at 4–6 km, and accompanying 

repressurization of the reservoir and conduit between 1987 

and 1992 (Moran, 1994), may reflect foundering of the roof of 

the reservoir and migration of a batch of magma to the shallower equilibration depths. As noted above, we suggest that 

the 2004–6 dacite was derived from the apical region of the 

reservoir. It originated from a geochemically distinct batch of 

crystal-rich magma not tapped during the 1980–86 eruptions 

that rose into the apical region after 1986.

The deep levels of the reservoir probably occupy a 

tensional dikelike structure within the St. Helens seismic zone 

(Musumeci and others, 2002), and the 1980–86 conduit may 

well consist of multiple intrusions and dikes (Cashman, 1992). 

However, unlike the situation at other volcanic centers (see, 

for example, Eichelberger and Izbekov, 2000), we see little 

surface geologic evidence for extensive dacitic dikes at shallow levels outside of the immediate crater area or alignment of 

older satellitic domes.

Studies of the plutonic substrate of volcanic arcs offer 

additional insights. Composite arc plutons and stocks with 

volumes of 10 km3

 or more are common at paleodepths of 

5–15 km, and arc batholiths, such as in the Sierra Nevada, are 

increasingly recognized to consist of multiple smaller magma 

bodies. However, many of these plutonic complexes developed over time spans measured in millions of years, not over a 

period of less than 300,000 years (Clynne and others, this volume, chap. 28; Crandell, 1987) in the tensional offset of a fault 

zone (Weaver and others, 1987). A better analogue for the 

Mount St. Helens plumbing system comes from a composite 

stratovolcano model based on ore deposits in the Andes (Sillitoe, 1973). Sillitoe’s composite model is similar to figure 22 

and consists of a complex composite stock of multiple feeder 

dikes and plugs of various ages overlying a shallow crustal 

pluton. We recognize that diagrams such as figure 22 are at 

best oversimplified cartoons compared to the complexity seen 

in eroded volcanic terranes, yet we believe they serve as useful 

constructs to evaluate volcano monitoring and petrologic data 

and to consider how magmas move through the crust.

Reservoir Pressurization and Eruption Triggering

One of the most important questions in volcanology and 

volcano-hazard mitigation is to understand what triggers eruptions. Before addressing this question for the 2004–6 Mount St. 

Helens eruption, however, we need to define what we mean by 

triggering. A pressurized reservoir-conduit system is a necessary 

condition for an eruption, but it is not a trigger, because pressure 

can be released slowly—for example, through passive degassing. A trigger requires a rapid increase in relative pressure of 

sufficient magnitude to fracture a new pathway to the surface 

(for a volcano with a solidified conduit or hydrothermal seal) or 

to raise the level of magma above the vent in an active magmafilled conduit. Accordingly, triggers can result by unloading 

from above or from a rapid increase in pressure or volume at 

depth. Candidates for triggers from above include edifice col-
682 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

lapse (as in 1980 at Mount St. Helens) or more subtle changes 

affecting an already pressurized system (an example: increased 

rainfall, Mastin, 1994). Indeed, Iverson and others (2006) calculate that once the system is pressurized, changes in pressure 

that are equivalent to a magma head of less than one meter are 

sufficient to start and stop the stick-slip cycles that they model 

for the current Mount St. Helens eruption.

Rapid addition of new magma and fluids to a crustal reservoir can increase pressure and trigger an eruption from below. 

This was the case for the eruption of Pinatubo in 1991, in which 

new hydrous mafic magma from depth entered a crystal-rich 

dacite reservoir, vesiculated, mingled, and created a buoyant 

plume, which rose through the viscous and crystal-rich upper 

part of the reservoir, increased pressure in the hydrothermal system, fractured a new pathway to the surface, and triggered the 

eruption (Pallister and others, 1996). In the Pinatubo case, deep 

long-period seismicity recorded and tracked ascent of basalt 

from 35 or 40 km deep to the crustal reservoir at depths of less 

than 14 km before the eruption (White, 1996).

For the 2004 Mount St. Helens eruption, the long-term 

pattern of seismicity since the 1980 eruption indicates that 

the reservoir and conduit system began to repressurize in 

late 1987 after the cessation of 1980–86 dome eruptions 

(Moran, 1994; Musumeci and others, 2002). Pressurization 

led to repeated seismic swarms at depths of less than about 8 

km, and rainfall-induced fracturing weakened hydrothermal 

seals and triggered small gas explosions in 1989–91 (Mastin, 

1994). Only a few deep (>3 km) earthquakes were located in 

the months before the 2004 eruption, and there have been no 

deep earthquakes during the 2004–6 eruptive period; the hundreds of thousands of small hybrid earthquakes that accompanied spine extrusion have been shallow (Moran and others, 

this volume, chap. 2; Thelen and others, this volume, chap. 

4). These relations indicate that the reservoir-conduit system 

was already pressurized before the eruption. The highest rate 

of seismic energy release took place during the September 

30–October 5, 2004, time period of shallow volcano-tectonic 

earthquakes and explosions that preceded extrusion of the 

initial spine (Moran and others, this volume, chap. 2). Given 

the small amount of pressure required for the hypothesized 

stick-slip cycles of spine extrusion (Iverson and others, 2006; 

Iverson, this volume, chap. 21), this initial period of greater 

energy release in early October 2004 indicates that the trigger 

for the eruption involved a somewhat larger pressure differential, which was accompanied by upward migration and 

venting of overpressured hydrothermal fluids.

Pressurization of a magma reservoir may result either 

from addition of mass (replenishment) or by fluid exsolution 

and expansion (driven by cooling and/or decompression). 

Long-term magmatic replenishment and magma mixing is a 

characteristic of the past at Mount St. Helens (Pallister and 

others, 1992; Gardner and others, 1995a; Clynne and others, this volume, chap. 28). However, as previously noted, 

the paucity of mafic inclusions, homogeneity of the dacite 

composition, lack of deep seismicity, and the low levels of gas 

emissions argue against ongoing (2004–6) replenishment, as 

does a lack of volcanic deformation during the period from at 

least as early as late 1997 (when continuous GPS monitoring 

was established) to late September 2004 (Lisowski and others, 

this volume, chap. 15).

Deformation data for the current eruption are best fit by 

an ellipsoidal reservoir with a centroid depth of 7 to 8 km 

(Lisowski and others, this volume, chap. 15) and are consistent with both recent seismic tomography models (Waite and 

Moran, 2006) and with the petrologic constraints outlined 

above. However, the volume of erupted dacite falls short 

of the modeled volume decrease at depth, requiring either 

magmatic replenishment or expansion of the reservoir magma 

(Lisowski and others, this volume, chap. 15). Geochemical 

modeling of gas-emission data suggests about 1.2–1.4 volume 

percent separate fluid phase at 220 MPa (about 8.6 km depth), 

which would increase to about 9–10 volume percent at 130 

MPa (5.2 km), near the apex of the magma reservoir (Gerlach 

and others, this volume, chap. 26). Modeling of magma and 

wall-rock physical properties (Mastin and others, this volume, 

chap. 22) to match the geodetically constrained rate of change 

in eruptive volume indicates only limited magma compressibility (low average volatile content of less than 2 volume 

percent) and a large magma reservoir (a few to as much as 

10 km3 at depths of 5 to >10 km), consistent with the recent 

seismic tomography (Waite and Moran, 2006). The modeling 

by Mastin and others of deformation data from the first 1.5 

year (this volume, chap. 22) also favors ongoing magmatic 

replenishment at a rate of about 1–1.3 m3/s. On balance, and 

given the uncertainties in the models, at this time it remains 

uncertain if the reservoir is being recharged, if the fluid phase 

in the reservoir is expanding, or both. Furthermore, if the 

reservoir is being recharged from below and erupted from 

the top, any significant additions of SO2 and CO2 gases have 

not yet made their way through the >10-km vertical extent of 

magma in the reservoir and conduit system to the surface.

A Petrologic Explanation

Magma in the upwelling limbs of convection cells 

exsolves volatiles because of their reduced solubility in the 

silicate melt on ascent to lower pressures. Therefore, without 

addition of new volatile-rich magma at depth, magma in a 

convecting reservoir will be progressively depleted by multiple 

convection cycles as volatiles in upwelling magma are lost to 

overlying stagnant magma and reservoir roof rocks. Any nonconvecting magmatic mush present at the roof of the reservoir 

would either be enriched or depleted in volatiles, depending on 

the ratio of volatile addition from upwelling magma to volatile 

loss through wall and roof rocks. During periods of frequent 

or continuous magmatic replenishment of the deep reservoir 

by hot and gas-rich magmas, enhanced convection would 

tend to enrich an upper mush zone in volatiles. During such 

periods, undercooling of newly introduced magmas at depth 

could also produce highly gas-rich plumes of buoyant mingled 

magma. Energetic plumes would be capable of penetrating 

and mobilizing previously gas-enriched mush in the reservoir 
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 683

roof to trigger large explosive eruptions. Evidence for such a 

process is seen in the eruptive products of Pinatubo volcano 

(Pallister and others, 1996) and is suggested by the presence 

of mingled pumices found in the basal set W tephra from the 

A.D. 1479 explosive eruption of Mount St. Helens, which initiated the Kalama eruptive period (Pallister and others, 1992). 

In contrast, in the absence of replenishment (or with diminished volumes of replenishment), continued convection would 

slowly deplete the reservoir in the lower-solubility volatiles, 

and the magmatic fluid phase would shift from relatively SO2- 

and CO2-rich to H2O- and Cl-rich compositions.

This latter condition is what we propose for the current eruption of Mount St. Helens. Since 1980, the magma 

reservoir has seen minimal replenishment, such that convection has produced a batch of well-mixed, crystal-rich magma 

with 65 percent SiO2 near the reservoir roof. A slow increase 

in pressure was brought about beginning in late 1987 by 

convection-driven exsolution of a water-rich volatile phase 

(which outpaced volatile losses through roof rocks). Magmatic 

pressure finally exceeded lithostatic load in late September 

2004, and the eruption ensued. As conduit pressures during the 

2004–6 eruption have not greatly exceeded lithostatic pressure 

and have been modulated by continuous crystallization of a 

shallow seismogenic plug (Iverson and others, 2006; Iverson, 

this volume, chap. 21), the eruption has been sustained at a 

low nonexplosive rate.

Lessons from the Past, Implications for the Future

Of dacites erupted at Mount St. Helens during the past 

500 years, the 17th-18th century summit-dome dacite (63–64 

percent SiO2), which erupted at the end of the Kalama eruptive period and formed the pre-1980 summit of the volcano 

(Crandell, 1987; Mullineaux, 1996; Pallister and others, 1992), 

is most akin to the dacite from the current eruption. Pre1980 summit-dome samples are characterized by extensive 

groundmass crystallization, resulting in a vesicle- and glasspoor microcrystalline groundmass. Although geochemically 

distinct from the 2004–6 dacite and texturally finer grained, 

the summit dome dacite has numerous similarities with it. 

Like the 2004–6 eruption, the summit-dome eruptions were 

relatively low in explosivity, producing only one pumiceous 

pyroclastic flow down the Toutle River drainage in early 

summit-dome time (R.P. Hoblitt, oral commun., 2006). The 

dominant deposits from the summit dome originated from 

small lithic pyroclastic flows and hot lahars, which were shed 

off the growing and collapsing lava dome, blanketed the upper 

slopes of the volcano, and fed multiple small lahars that spread 

out onto the flanks of the volcano. The nonpumiceous pink “z” 

ash (Mullineaux, 1996) is the only tephra layer associated with 

summit-dome eruptions. It occurs as a thin (typically less than 

1 cm) ash deposit on all flanks of the volcano and was likely 

derived from small ash clouds produced during dome collapse 

events over an extended period of time. Also, like some of the 

2004–6 dome rocks, lithic fragments in summit-dome deposits 

are characterized by red oxidation fractures and rinds.

The greater extent of summit-dome deposits is a consequence of prolonged emplacement over 100 years or more 

(Clynne and others, 2005; Hoblitt, 1989; Yamaguchi and 

Hoblitt, 1995), large volume, and its location high atop the 

steep-sided cone. We do not know if the summit dome was 

built as a series of spines, for such features are rarely preserved in the geologic record. However, the summit dome was 

clearly built at a low mean eruptive rate. It spalled and disintegrated in place, forming the rubbly pre-1980 summit, and it 

shed multiple lithic pyroclastic flows and lahars. It produced 

only relatively small and diffuse ash clouds, similar to those 

produced during the 2004–6 eruption. If the current eruption 

were to follow a similar path, it could produce continuous or 

intermittent dome-forming eruptions for decades. Although 

low eruption rates do not guarantee a long-duration episode, 

a global compilation of historical dome-forming eruptions 

indicates that only low-rate eruptions last for many years or 

decades (Chris Newhall, written commun., 2005).

Acknowledgments

The authors wish to acknowledge the several dozen 

petrologists and geochemists from a variety of academic and 

government institutions who joined us during the ongoing 

eruption to form an informal Mount St. Helens petrology 

working group (Pallister and others, 2005; Rowe and others, 2008b). The contributions of these scientists are well 

represented in this volume and contribute greatly to our 

understanding of the ongoing eruption. We thank our U.S. 

Geological Survey (USGS) colleagues from the Cascades 

Volcano Observatory (CVO) and other USGS offices and 

from the Pacific Northwest Seismic Network at the University 

of Washington for sharing of data and discussion of ideas and 

interpretations, which have made our work on this eruption 

scientifically stimulating and rewarding. We thank Bobbie Myers of CVO and Jeff Linscott of JL Aviation for their 

emphasis on safety as well as innovation during our helicopter missions at Mount St. Helens. And finally, we thank 

Charlie Bacon and Malcolm Rutherford for technical reviews 

of this manuscript.

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SH304-2A, spine 3

1,000 points counted

SH304-2G1, spine 3

1,000 points counted

SH305-1A, spine 3

1,300 points counted

SH315-4, spine 4

1,300 points counted

Optical mode, percent TS DRE TS DRE TS DRE TS TS DRE TS DRE

Plagioclase 30 31 31 32 22 22 27 27 30 39 39Hypersthene 4 4 4 4 3 3 4 4 5 6 6Hornblende 8 8 2 2 2 2 4 4 4 4 4Oxides 2 2 2 2 1 1 2 2 2 2 2Silica phases n.d. n.d. n.d. n.d. n.d. n.d. 0.2 0.2 0.2 n.d. n.d.Voids 2 -- 2 -- 0.4 -- 9 9 -- 1 --

Matrix 55 55 59 60 72 73 53 53 59 49 49

Sum 100 100 100 100 100 100 100 100 100 100 100Phenocrysts 44 45 39 40 27 27 37 37 41 51 51

Table 2. Modal data for samples from 2004–2006 Mount St. Helens lava dome. Further explanation of methods in “Laboratory Methodssection.[Optical modes, as volume percent, are derived by counting 1,000 or 1,300 points at 100× magnification. Grains and voids smaller than 30 μm in maximum dimension are counted as matrix, consequently most silica phases, which occur mainly as late-stage microlites, were not detected (n.d.) by optical point counting. Some grains were large enough to detect and count, but as late-stage phases, they were not included in the total phenocryst counts. TS, thin section; DRE, dense-rock-equivalent abundances. DRE calculated by subtracting voids and renormalizing the optical modes and microbeam “ALL” modes. Dashes indicate absence of a value (for example, as in no void space in recalculated DRE columns). Hornblende, used informally for a range of amphibole compositions, forms large (1–3 mm) phenocrysts; consequently its abundance in individual sections may not be statistically representative.Microbeam map analysis modes, as volume percent, derived from image analysis of 1-cm2 microprobe-stage raster maps of X-ray intensities for Si, Al, Mg, Fe, K and backscattered electron intensity, combined with image analysis of SEM backscattered electron images of 100-μm2 areas.Columns labeled “All” tabulate complete modal analysis, including microlites (1–30 μm maximum diameter), obtained by microbeam imaging at scales of 1 cmand 100 μm2.Probe DRE for phenocryst abundance calculated from 1,024 analyses of 1-μm spots distributed over a 1-cm2 grid and binned (after excluding 262 analyses with 

totals >102 percent or <85 percent) according to mineral types and glass using the following compositional and stoichiometric rules: Plagioclase if Si+Al = 3.85–4.15 (based on 8 oxygens) and K2O <2 percent;

Anorthoclase if Si+Al = 3.85–4.15 (based on 8 oxygens) and K2O >2 percent;

Hypersthene if MgO >17 percent;Hornblende if MgO = 8–17 percent and Na2O>1 percent;

Quartz if SiO2>78 percent;

Silica phase (tridymite, cristobalite) if SiO2 =72–78 percent;

Fe-Ti oxide if FeO+TiO2 >30 percent and SiO2 <10 percent.

Plagioclase category as part of microprobe analysis may include anorthoclase, except where counted separately in the analysis of SH315-4.Microlite abundance calculated by subtracting total crystals determined by microbeam analysis from phenocryst abundance determined by microbeam (m) or optical (o) methods.Values for glass by K2O balance are weight percent and calculated on a volatile-free basis from the percentage of glass needed to give the K2O determined by 

bulk chemical analysis. To convert to volume percent, multiply by 1.1]
SH319-1, spine 5 1,300 points counted SH321-1A, spine 6 1,300 points counted SH325-1A, spine 7 1,300 points counted Average microcrystalline dacite Optical mode, percent TS DRE TS DRE TS DRE TS DRE SD Plagioclase 36.3 36.9 37.1 37.8 33.8 34.4 33.9 34.8 3.7 Hypersthene 5.2 5.3 6.3 6.4 2.5 2.6 4.7 4.8 1.4 Hornblende 5.5 5.6 3.5 3.6 2.6 2.7 4.6 4.7 2.0 Oxides 1.4 1.4 1.1 1.1 1.4 1.4 1.6 1.6 0.4 Silica phases 1.1 1.1 1.0 1.0 1.1 1.1 0.8 0.9 0.4 Voids 1.6 -- 1.8 -- 1.8 -- 2.8 -- -- Matrix 48.8 49.6 49.2 50.1 56.8 57.8 51.9 53.5 4.4 Sum 100.0 100.0 100.0 100.0 100.0 100.0 100.3 100.3 0.1 Phenocrysts 48.5 49.3 49.0 48.9 41.4 41.1 45.6 46.0 4.3

SH304-2A, spine 3

1,000 points counted

SH304-2G1, spine 3

1,000 points counted

SH305-1A, spine 3

1,300 points counted

SH315-4, spine 4

1,300 points counted

Microbeam map analysis, percent

SH-305-1A Probe (DRE)

TS DRE ALL DRE TS All DRE TS All DRE DRE

Plagioclase 32 35 47 64 22 22 22 32 33 52 48Anorthoclase n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 15Hypersthene 3 4 4 5 2 4 4 5 5 8 5Hornblende 5 6 6 8 2 2 2 3 3 6 3Oxides 1 1 1 1 0.3 0.4 0.4 1 1 2 1

Silica phases n.d. n.d. n.d. n.d. n.d. n.d. n.d. n.d. 1 2 16Voids 8 -- 25 -- 1 1 -- 36 37 -- --Glass n.d. n.d. 17 22 73 71 72 18 19 30 13Matrix 50 55 4

Sum 100 100 100 100 100 100 100 100 100 100 100Phenocrysts 42 45 -- 45 26 -- -- 41 -- -- --

Total crystals -- -- 58 78 -- 28 28 -- 43 69 87

Microlites (m) -- -- -- 33 -- -- 2 -- -- 29 --

Microlites (o) -- -- -- 38 -- -- 1 -- -- 28 37

Glass

K2O

Bulk

K2O

Percent

glass

Glass

K2O

Bulk

 K2O

Percent

glass

K2O weight percent 3.38 1.83

54

5.09 1.41

Glass by K 28 2O balance 

Table 2. Modal data for samples from 2004–2006 Mount St. Helens lava dome. Further explanation of methods in “Laboratory Methods” section.—Continued
Sample No. SiO2 Al2O3 FeO MgO CaO Na2O K2O TiO2 P2O5 MnO BaO Total An 

Anorthoclase and orthoclase microlites (less than 30 μm) SH322-1-d-7 68.3 18.3 0.47 0.00 0.40 5.73 6.70 0.06 0.02 0.01 0.13 100.1 2 

SH322-1-d-5 65.3 21.5 0.37 0.02 2.95 7.54 2.50 0.04 0.03 0.00 0.04 100.2 15 

Plagioclase microlites (less than 30 μm) SH322-1-d-4 60.9 24.0 0.36 0.00 5.05 7.58 1.48 0.03 0.01 0.00 0.07 99.5 25 

SH321-1C-7d-3 (i) 62.3 23.0 0.52 0.00 3.77 8.00 1.42 0.05 0.01 0.00 0.10 99.2 19 SH321-1C-7b-2 (h) 60.8 24.5 0.48 0.00 5.43 7.53 1.11 0.04 0.00 0.01 0.04 99.9 27 SH321-1C-7c-7 (h) 59.9 25.0 0.47 0.00 6.14 7.42 0.93 0.01 0.00 0.00 0.03 99.8 30 SH321-1C-7d-2 (i) 60.3 24.3 0.51 0.02 5.32 7.62 0.82 0.01 0.01 0.01 0.07 99.0 26 SH321-1C-7d-1 (i) 57.0 26.6 0.58 0.00 8.14 6.66 0.45 0.03 0.00 0.01 0.02 99.5 39 

Plagioclase microphenocryst (100 μm) SH322-1-b-17 58.6 26.2 0.43 0.01 7.46 6.94 0.58 0.00 0.01 0.01 0.03 100.3 36 

SH322-1-b-18 59.5 25.8 0.35 0.01 6.76 7.45 0.40 0.02 0.02 0.00 0.05 100.3 33 

SH322-1-b-19 58.5 26.0 0.31 0.00 7.36 7.23 0.23 0.03 0.00 0.00 0.06 99.7 36 

SH322-1-b-20 59.8 25.5 0.49 0.00 6.83 7.35 0.67 0.01 0.00 0.01 0.07 100.7 33 Quartz microlites SH321-1C-7c-6 (h) 100.0 0.14 0.29 0.01 0.01 0.00 0.03 0.04 0.00 0.01 0.02 100.6 SH321-1C-7b-4 (h) 98.3 0.38 0.16 0.01 0.02 0.07 0.05 0.07 0.01 0.00 0.05 99.2 SH321-1C-7d-4 (i) 99.6 0.12 0.17 0.01 0.01 0.02 0.00 0.04 0.00 0.00 0.01 99.9 SH321-1C-7d-5 (i) 99.5 0.07 0.16 0.00 0.03 0.00 0.01 0.06 0.00 0.00 0.02 99.9 SH322-1-d-3 98.7 0.63 0.11 0.02 0.05 0.13 0.05 0.10 0.00 0.02 0.03 99.8 

Tridymite microlites

SH322-1-b-9 99.1 1.06 0.07 0.00 0.01 0.40 0.02 0.12 0.00 0.00 0.02 100.8 SH322-1-b-10 99.3 0.94 0.07 0.00 0.02 0.34 0.02 0.06 0.01 0.01 0.01 100.7 

SH322-1-b-11 98.7 1.02 0.11 0.00 0.03 0.39 0.02 0.10 0.01 0.00 0.00 100.3 SH322-1-d-2 97.0 1.52 0.09 0.00 0.03 0.41 0.11 0.11 0.00 0.02 0.07 99.4 

Cristobalite microlites

SH312-1 93.9 4.27 0.35 0.00 0.15 0.71 0.68 0.15 0.00 0.01 0.04 100.3 SH321-1C-7b-5 (h) 93.3 3.57 0.25 0.03 0.36 1.27 0.39 0.08 0.01 0.02 0.00 99.3 

Table 3. Electron-microprobe analyses of matrix minerals in 2004–2006 dacite.[Oxides in weight percent. Anorthite (An), albite (Ab), and orthoclase (Or) components expressed as molecular percent. Cation totals based on recalculation of atomic proportions using eight oxygen anions. Detection limits are about 0.01 percent; values less than this amount are reported as 0.00 percent.]
Sample and Spot No. Size Phase SiO2 Al2O3 FeO MgO CaO Na2O K2O TiO2 P2O5 MnO BaO Cl SO3 Total H2O 

Glass inclusions in microcrystalline dacite SH305-1kc g-6 (core) 1 ip 77.7 11.2 1.02 0.22 0.34 2.82 5.33 0.33 0.05 0.01 0.08 0.081 n.d. 99.23 0.8 

SH305-1kc g-7 (core) 1 ip 77.3 10.7 1.52 0.26 0.38 2.66 5.32 0.54 0.05 0.00 0.07 0.120 n.d. 98.94 1.1 SH305-1kc g-8 (core) 1 ip 76.8 11.2 1.28 0.18 0.31 2.56 5.36 0.55 0.09 0.02 0.07 0.084 n.d. 98.45 1.6 SH305-1kc g-9 (core) 1 ip 77.2 11.1 1.19 0.18 0.37 2.59 5.40 0.49 0.07 0.01 0.08 0.064 n.d. 98.74 1.3 SH305-1kc g-10 (core) 1 ip 77.9 11.1 1.25 0.18 0.33 2.60 5.24 0.46 0.05 0.03 0.09 0.057 n.d. 99.28 0.7 SH305-1kc g-12 (core) 1 ip 77.4 10.8 0.96 0.04 0.23 2.68 5.16 0.35 0.07 0.00 0.05 0.060 n.d. 97.77 2.2 SH305-1kc g-5 1 ip 77.3 11.0 1.14 0.26 0.37 2.61 5.37 0.27 0.04 0.01 0.07 0.086 0.013 98.49 1.5 SH305-1kc g-21 1 ip 78.0 11.6 1.21 0.11 0.38 2.49 5.01 0.45 0.10 0.01 0.04 0.040 0.007 99.43 0.6 SH305-1kc g-22 1 ih 76.6 12.4 1.35 0.03 0.27 3.01 5.72 0.34 0.08 0.01 0.05 0.110 n.d. 99.97 0.0 SH305-1kc g-27 1 ih 71.5 16.8 1.61 1.42 0.99 4.09 2.37 0.17 0.09 0.05 0.06 0.052 0.062 99.27 0.7 SH305-1kc g-23 1 ih 76.8 12.1 1.12 0.01 0.24 3.02 5.63 0.31 0.05 0.02 0.06 0.096 n.d. 99.45 0.5 SH305-1kc g-25 1 ia 74.3 13.3 0.98 0.08 0.38 3.14 5.50 0.64 0.07 0.00 0.07 0.145 0.007 98.61 1.4 Average 76.6 11.9 1.22 0.25 0.38 2.85 5.12 0.41 0.07 0.02 0.07 0.08 0.022 98.97 1.0 Standard deviation 1.9 1.7 0.20 0.38 0.20 0.44 0.88 0.14 0.02 0.01 0.01 0.03 0.027 0.59 0.6 SH304-2A9b hb1 1 ia 75.0 13.7 0.84 0.01 0.88 2.45 5.23 0.20 0.02 0.01 0.10 0.105 0.012 98.54 1.5 SH304-2A9b hb1 20 ia 75.3 13.6 0.86 0.01 0.85 2.03 4.98 0.23 0.06 0.04 0.11 0.112 0.007 98.16 1.8 SH304-2A9b hb1 20 ia 75.2 13.5 0.87 0.01 0.88 2.32 5.09 0.19 0.01 0.00 0.10 0.103 n.d. 98.37 1.6 SH304-2A9b hb6 5 ia 74.1 12.9 0.81 0.01 0.63 3.44 5.55 0.22 0.03 0.03 0.10 0.125 n.d. 97.93 2.1 SH304-2A9b lghb gl1 1 ia 74.0 14.2 0.91 0.12 0.80 3.69 5.02 0.17 0.05 0.02 0.10 0.139 0.022 99.25 0.7 SH304-2A9b lghb gl2 1 ia 74.0 14.1 0.93 0.11 0.77 3.96 5.17 0.21 0.07 0.03 0.08 0.139 n.d. 99.58 0.4 SH304-2A9b lghb gl3 1 ia 72.1 15.7 1.00 0.24 1.40 4.38 4.20 0.17 0.16 0.02 0.06 0.100 0.012 99.54 0.5 SH304-2A9b lghb gl4 1 ia 73.4 15.6 0.84 0.19 1.35 4.21 3.42 0.19 0.08 0.04 0.08 0.113 0.030 99.55 0.4 SH304-2A9b lghb gl5 1 ia 73.9 15.7 0.83 0.20 1.25 4.12 3.36 0.12 0.10 0.05 0.12 0.112 0.032 99.93 0.1 SH304-2A9b lghb gl6-1 1 ia 73.7 15.7 0.82 0.18 1.33 4.24 3.38 0.16 0.12 0.05 0.08 0.107 0.020 99.92 0.1 Average 74.1 14.5 0.87 0.11 1.01 3.48 4.54 0.19 0.07 0.03 0.09 0.12 0.019 99.08 0.9 Standard deviation 1.0 1.1 0.06 0.09 0.29 0.89 0.87 0.03 0.05 0.02 0.02 0.01 0.010 0.75 0.8 Matrix glass in microcrystalline dacite SH305-1kc g-1 1 m 77.7 13.3 0.63 0.02 0.69 2.20 4.47 0.27 0.06 0.02 0.06 0.026 n.d. 99.38 0.6 SH305-1kc g-30 1 m 75.8 12.1 0.70 0.02 0.41 3.36 5.37 0.34 0.07 0.01 0.03 0.030 0.013 98.24 1.8 SH305-1kc g-33 1 m 76.8 11.8 0.68 0.02 0.19 2.21 5.24 0.33 0.08 0.00 0.04 0.036 0.021 97.41 2.6 SH305-1ag-1 1 m 76.3 14.2 0.64 0.03 0.92 4.00 4.74 0.27 0.07 0.01 0.04 0.024 n.d. 101.28 0.0 SH305-1ag-35 1 m 77.4 11.9 0.83 0.03 0.23 2.97 5.63 0.31 0.05 0.01 0.03 0.023 n.d. 99.42 0.6 Average 76.8 12.7 0.69 0.02 0.49 2.95 5.09 0.30 0.07 0.01 0.04 0.03 0.017 99.15 1.1 Standard deviation 0.8 1.0 0.08 0.01 0.31 0.77 0.48 0.03 0.01 0.01 0.01 0.01 0.006 1.46 1.0 

Table 4. Electron-microprobe analyses of matrix glass and glass inclusions in 2004–2006 Mount St. Helens dacite and two quenched inclusions from [Oxide abundances in weight percent; Cl and S recalculated as ppm. Sample numbers follow standard format (for example, SH305-1) but have additional coding for specific thishow additional feature, “core,” to refer to melt inclusions in reacted core of plagioclase. Size indicates beam diameter, in microns. Phases coded as m, matrix; i, inclusion; p, pamphibole; o, oxide. For inclusions, codes combined to show host mineral. Count rates were monitored for Na, K, and Si during each analysis, and these elements were correct“Laboratory Methods” section. H2O calculated by difference. Detection limit, D.L., for S ranges between 70 and 86 ppm and corresponds to the value at which one standard dev

cent of the reported value. Lower values reported for S whenever a peak is detected above background, along with larger standard deviation. None detected, n.d., indicates no pelevels. Column headed “SD, S” shows standard deviation for S, in percent, based on counting statistics for individual analyses; dashes in this column are for analyses in which nin other columns indicate insufficient data to calculate meaningful averages or standard deviations.]
Sample and Spot No. Size Phase SiO2 Al2O3 FeO MgO CaO Na2O K2O TiO2 P2O5 MnO BaO Cl SO3 Total H2O 

Matrix glass in glassy flow-banded dacite SH304-2G1g1-m1 5 m 73.5 13.9 0.91 0.04 1.31 3.96 3.30 0.22 0.07 0.01 0.09 0.089 n.d. 97.40 2.6 SH304-2G1g1-m2 5 m 73.1 13.9 1.17 0.17 1.40 4.26 3.38 0.16 0.06 0.01 0.04 0.079 n.d. 97.67 2.3 SH304-2G1g1-m3 5 m 73.5 13.9 1.22 0.11 1.37 4.22 3.39 0.21 0.04 0.03 0.07 0.089 n.d. 98.08 1.9 SH304-2G1g1-m4 20 m 73.5 13.8 0.98 0.10 1.06 4.19 3.50 0.19 0.02 0.02 0.07 0.089 n.d. 97.59 2.4 SH304-2G1g1-m5 1 m 73.6 14.0 0.87 0.04 1.30 4.15 3.46 0.21 0.04 0.00 0.07 0.084 0.008 97.87 2.1 SH304-2G1g1-m6 1 m 73.3 13.9 1.02 0.10 1.28 4.05 3.45 0.17 0.05 0.04 0.09 0.089 n.d. 97.49 2.5 SH304-2G1g1-m7 1 m 73.3 13.9 1.09 0.12 1.19 3.92 3.51 0.18 0.07 0.03 0.08 0.087 0.008 97.41 2.6 SH304-2G1g1-m8 1 m 73.2 14.0 0.96 0.08 1.18 4.09 3.55 0.21 0.07 0.03 0.09 0.088 n.d. 97.55 2.5 SH304-2G1g1-m9 1 m 73.7 13.8 0.89 0.04 1.18 4.17 3.59 0.18 0.05 0.02 0.09 0.088 n.d. 97.81 2.2 SH304-2G1g1-m10 1 m 73.3 14.0 0.79 0.04 1.19 4.23 3.62 0.22 0.03 0.01 0.06 0.084 n.d. 97.52 2.5 SH304-2G1g1-m11 1 m 73.6 14.0 0.94 0.06 1.44 4.20 3.16 0.20 0.09 0.04 0.07 0.081 n.d. 97.89 2.1 SH304-2G5g1 40 m 74.3 12.5 1.18 0.10 1.15 4.76 2.61 0.24 0.02 0.00 0.07 0.075 n.d. 96.98 3.0 Average 73.5 13.8 1.00 0.08 1.25 4.18 3.38 0.20 0.05 0.02 0.07 0.09 0.008 97.61 2.4 Standard deviation 0.3 0.4 0.14 0.04 0.11 0.21 0.27 0.02 0.02 0.01 0.01 0.00 0.000 0.29 0.3 Glass inclusions in glassy flow-banded dacite SH304-2G1g1-g1 1 io 73.9 14.0 1.33 0.10 1.43 4.34 3.11 0.48 0.07 0.00 0.09 0.099 n.d. 98.95 1.1 SH304-2G1g3-1 1 ih 73.6 14.0 0.94 0.02 1.48 4.08 3.17 0.21 0.01 0.01 0.08 0.086 n.d. 97.74 2.3 SH304-2Gg4-1 1 ih 76.4 12.2 0.85 0.07 1.00 3.88 3.29 0.03 0.06 0.05 0.08 0.082 0.005 97.96 2.0 SH304-2Gg6-1 1 ip 72.9 14.5 0.54 0.01 2.47 5.59 2.26 0.13 0.06 0.03 0.06 0.081 0.010 98.59 1.4 Average 74.2 13.7 0.92 0.05 1.60 4.47 2.96 0.21 0.05 0.02 0.08 0.09 0.007 98.31 1.7 Standard deviation 1.5 1.0 0.33 0.04 0.62 0.77 0.47 0.19 0.02 0.02 0.01 0.01 0.003 0.56 0.6 Glass inclusion in dacite fragment in gouge11-4 gou gr7 1 ih 63.7 17.8 2.78 0.23 3.32 4.90 2.41 1.24 0.41 0.03 0.06 0.281 0.031 97.17 2.8 2,Matrix glass in fragments from tephra erupted January 16, 2005WS_1_19 g1 20 m 70.6 14.7 2.10 0.52 2.21 4.89 2.17 0.38 0.11 0.07 0.08 0.076 n.d. 97.96 2.0 WS_1_19 g2 20 m 70.7 14.7 2.11 0.54 2.25 4.90 2.19 0.37 0.07 0.05 0.07 0.080 n.d. 98.07 1.9 WS_1_19 2-22 20 m 71.0 14.5 2.13 0.46 2.07 3.88 2.16 0.32 0.07 0.04 0.07 0.088 n.d. 96.73 3.3 WS_1_19 2-23 20 m 71.5 14.8 1.99 0.48 2.24 4.97 2.09 0.33 0.08 0.04 0.04 0.075 n.d. 98.58 1.4 WS_1_19 5-25 20 m 73.5 12.7 1.80 0.28 1.30 4.41 2.56 0.36 0.09 0.03 0.07 0.083 n.d. 97.26 2.7 WS_1_19 5-26 20 m 73.1 12.8 1.84 0.26 1.30 4.43 2.46 0.41 0.08 0.03 0.06 0.087 n.d. 96.84 3.2 Average 71.7 14.0 2.0 0.4 1.9 4.6 2.3 0.4 0.1 0.0 0.1 0.1 n.d. 97.6 2.4 Standard deviation 1.3 1.0 0.1 0.1 0.5 0.4 0.2 0.0 0.0 0.0 0.0 0.0 n.d. 0.7 0.7 Glass inclusions in fragments from tephra erupted January 16, 2005WS_1_19 gi-43 1 io 73.8 14.9 1.01 0.02 0.30 4.71 4.00 0.61 0.05 0.03 0.07 0.389 0.012 99.87 0.1 3,WS_1_19 gi-44 1 io 74.3 15.1 0.94 0.02 0.58 5.84 3.14 0.55 0.12 0.01 0.07 0.369 n.d. 101.05 0.0 3,

Table 4. Electron-microprobe analyses of matrix glass and glass inclusions in 2004–2006 Mount St. Helens dacite and two quenched inclusions from th[Oxide abundances in weight percent; Cl and S recalculated as ppm. Sample numbers follow standard format (for example, SH305-1) but have additional coding for specific thishow additional feature, “core,” to refer to melt inclusions in reacted core of plagioclase. Size indicates beam diameter, in microns. Phases coded as m, matrix; i, inclusion; p, pamphibole; o, oxide. For inclusions, codes combined to show host mineral. Count rates were monitored for Na, K, and Si during each analysis, and these elements were correct“Laboratory Methods” section. H2O calculated by difference. Detection limit, D.L., for S ranges between 70 and 86 ppm and corresponds to the value at which one standard dev

cent of the reported value. Lower values reported for S whenever a peak is detected above background, along with larger standard deviation. None detected, n.d., indicates no pelevels. Column headed “SD, S” shows standard deviation for S, in percent, based on counting statistics for individual analyses; dashes in this column are for analyses in which nin other columns indicate insufficient data to calculate meaningful averages or standard deviations.]
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 695

Sample No. SiO2 Al2O3 FeTO3 MgO CaO Na2O K2O TiO2 MnO Total

SH315-4 host 1

758 point average 62.11 17.54 2.82 1.36 4.83 4.30 1.25 0.34 0.04 94.57

Normalized to 100% 65.67 18.54 2.98 1.44 5.10 4.55 1.32 0.36 0.05 100.00

SH315-4 host by XRF 65.16 17.24 4.38 1.87 4.75 4.49 1.41 0.61 0.07 100.00

Correction factors2 0.99 0.93 1.47 1.30 0.93 0.99 1.07 1.72 1.53 1.00

SH315-4 inclusion 3

724 point average 58.95 16.47 5.95 4.06 6.40 3.54 1.19 0.87 0.06 97.50

Normalized to 100% 60.47 16.89 6.10 4.17 6.56 3.63 1.22 0.89 0.06 100.00

Corrected analysis 60.00 15.71 8.98 5.43 6.11 3.59 1.31 1.54 0.10 100.00

SH321-1C inclusion 4

523 point average 64.31 16.99 4.30 1.85 5.09 4.07 1.23 0.52 0.05 98.42

Normalized to 100% 65.34 17.27 4.37 1.88 5.17 4.14 1.25 0.53 0.05 100.00

Corrected analysis 64.84 16.06 6.43 2.45 4.81 4.09 1.34 0.92 0.07 100.00

1

 Rejected were 263 analyses from 1,021 total points analyzed.

2

 Correction factors are abundance ratios of microprobe to XRF data, used to correct microprobe analyses.

3

 Rejected were 236 analyses from 960 total points analyzed.

4

 Rejected were 213 analyses from 736 total points analyzed.

Table 5. Major-element compositions of host dacite and small quenched inclusions from 2004–2006 Mount St. 

Helens dome lava.

[Oxides reported as weight percent. Compositions determined by automated electron-microprobe, sampling from 736 to 1,021 points 

distributed across gridded areas of polished sections and compared with bulk rock X-ray fluorescence (XRF) analyses of the dacite. 

Rejected were analyses with totals greater than 102 percent or less than 85 percent (resulting from, for example, beam overlap with void 

space in the microprobe mount). Comparison of normalized average composition of host by microprobe to that by XRF shows excellent 

agreement (within 1 percent) for SiO2 and Na2O, good agreement (within 7 percent of reported value) for Al2O3, CaO, and K2O, and 

moderate (microprobe average is 30 percent low relative to XRF analysis MgO) to poor agreement (microprobe average is 47–72 percent 

low for TiO2, FeO, and MnO). Poor agreement attributed to undersampling of small oxide grains in the microprobe analyses, many of 

which were likely rejected due to low totals. See “Laboratory Methods” section for additional information on methodology.]
696 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample No.

Abs

4500 

cm-1

Abs

5200 

cm-1

Abs

3570 

cm-1

Abs

1630 

cm-1

Thickness

cm

Density

g/L

H2Ot

3570 

cm-1

H2Om

1630 

cm-1

H2Om

5200 

cm-1

OH=

4500 

cm-1

H2Ot3

4500 + 

5200 cm-1

H2Ot

4500 + 

1630 cm-1

SH304-2G1-s1 0.0045 0.0069 0.520 0.248 0.00372 2317 1.598 0.942 0.896 0.544 1.439 1.486

SH304-2G1-s2 0.0057 0.0070 0.672 0.307 0.00372 2317 2.066 1.167 0.909 0.689 1.597 1.855

SH304-2G1-s3 0.0041 0.0075 0.469 0.230 0.00372 2317 1.442 0.874 0.974 0.495 1.469 1.369

SH304-2G1-s4 0.0062 0.0070 0.649 0.285 0.00372 2317 1.995 1.083 0.909 0.749 1.658 1.832

SH304-2G1-s5 0.0056 0.0066 0.639 0.289 0.00372 2317 1.964 1.098 0.857 0.677 1.533 1.775

SH304-2G1-s6 0.0070 0.0090 0.837 0.369 0.0053 2317 1.806 0.984 0.820 0.594 1.414 1.578

SH304-2G1-s7 0.0080 0.0100 0.879 0.406 0.0053 2317 1.896 1.083 0.911 0.678 1.590 1.761

SH304-2G1-s8 0.0077 0.0098 0.859 0.382 0.0053 2317 1.853 1.019 0.893 0.653 1.546 1.672

SH304-2G1-s9 0.0088 0.0100 1.020 0.436 0.0062 2317 1.881 0.994 0.779 0.638 1.417 1.632

SH304-2G1-s10 0.0090 0.0110 1.028 0.462 0.0062 2317 1.896 1.053 0.857 0.652 1.509 1.706

SH304-2G1-s11 0.0049 0.0079 0.662 0.292 0.00372 2317 2.035 1.110 1.026 0.592 1.618 1.702

SH304-2G1-s12 0.007 0.0066 0.778 0.347 0.005 2317 1.779 0.981 0.637 0.629 1.267 1.610

SH304-2G1-s13 0.005 0.0070 0.645 0.315 0.00372 2317 1.983 1.197 0.909 0.604 1.513 1.801

SH304-2G1-s14 0.0064 0.0067 0.544 0.256 0.00372 2317 1.672 0.973 0.870 0.773 1.643 1.746

SH304-2G1-s15 0.008 0.0110 0.876 0.396 0.0053 2317 1.890 1.056 1.002 0.678 1.681 1.735

SH304-2G1-s16 0.01 0.0130 1.148 0.516 0.0062 2317 2.117 1.177 1.013 0.725 1.737 1.901

Average 1.867 1.049 0.891 0.648 1.539 1.698

Standard deviation 0.178 0.090 0.096 0.072 0.121 0.139

SH304-2G2-s1 0.0097 0.0137 1.069 0.484 0.0052 2323 2.345 1.312 1.269 0.836 2.105 2.149

SH304-2G2-s2 0.0084 0.0127 1.053 0.488 0.0052 2323 2.309 1.323 1.176 0.724 1.901 2.047

SH304-2G2-s3 0.0145 0.0172 1.558 0.651 0.0078 2323 2.278 1.177 1.062 0.833 1.896 2.010

SH304-2G2-s4 0.0095 0.0155 1.149 0.514 0.0052 2323 2.520 1.394 1.436 0.819 2.255 2.213

SH304-2G2-s5 0.017 0.0160 1.685 0.767 0.0078 2323 2.464 1.387 0.988 0.977 1.965 2.364

SH304-2G2-s6 0.0129 0.0172 1.601 0.723 0.0078 2323 2.341 1.307 1.062 0.741 1.804 2.048

SH304-2G2-s7 0.0092 0.0110 0.988 0.445 0.0052 2323 2.167 1.207 1.019 0.793 1.812 2.000

SH304-2G2-s8 0.0142 0.0163 1.619 0.733 0.0078 2323 2.367 1.325 1.007 0.816 1.823 2.141

SH304-2G2-s9 0.0106 0.0165 1.357 0.612 0.0078 2323 1.984 1.106 1.019 0.609 1.628 1.716

SH304-2G2-s10 0.013 0.0162 1.549 0.716 0.0078 2323 2.265 1.294 1.000 0.747 1.748 2.041

Average 2.304 1.283 1.104 0.790 1.894 2.073

Standard deviation 0.150 0.092 0.147 0.095 0.180 0.168

Table 6. Fourier-transform infrared (FTIR) spectrographic analyses of matrix glass in samples SH304-2G1 and SH304-2G2.

[Abs, absorbance at the indicated wavenumber (for example, 4,500 cm−1). Absorbance at wavenumber 2,350 cm−1 was below detection level of 20 ppm as CO2. 

Water and hydroxyl groups all reported in weight percent; H2Ot, total dissolved water; H2Om, molecular water; OH-, dissolved water as hydroxyl groups. For 

data reduction purposes, standard values from Newman and others (1986) used for molar absorptivity, ε, for water in rhyolitic glass. They are ε3750 cm−1=68; 

ε1630 cm−1=55; ε5200 cm−1=1.61; ε4500 cm−1=1.73.]
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 697

2004–6 dacite Glassy fragments Roof of spine 1 Roof of spine 3 1986

Spine No. 3–7 3 1 1 1 3

Sample -- SH304-

2G1

SH300-

1A3A

SH300-

1C2 SH300 (dense) SH301 SH226

XRF AVG (23) SD (23) AVG(3) SD(3) AVG(3) SD(3)

SiO2 64.93 0.09 66.85 64.64 64.65 63.43 0.35 63.38 0.29 63.78

Al2O3 17.21 0.08 16.77 17.12 17.23 17.45 0.12 17.49 0.02 17.35

FeTO3 4.40 0.05 3.87 4.65 4.43 4.93 0.03 4.97 0.13 4.89

MgO 1.85 0.05 1.24 2.01 1.85 2.17 0.05 2.16 0.08 2.02

CaO 4.73 0.03 3.99 4.86 4.77 5.30 0.09 5.26 0.08 5.04

Na2O 4.61 0.08 4.64 4.50 4.73 4.56 0.05 4.55 0.04 4.60

K2O 1.42 0.02 1.83 1.34 1.44 1.24 0.01 1.27 0.02 1.37

TiO2 0.62 0.01 0.52 0.63 0.62 0.66 0.01 0.67 0.01 0.67

P2O5 0.15 0.02 0.21 0.17 0.21 0.19 0.00 0.17 0.01 0.18

MnO 0.07 0.00 0.07 0.07 0.07 0.08 0.00 0.08 0.00 0.08

Total 100.00 0.00 100.00 100.00 100.00 100.00 0.00 100.00 0.00 100.00

LOI 0.09 0.10 2.21 0.41 0.35 0.06 0.03 0.25 0.15 1.16

ICP-MS AVG (23) SD (23)

Ba 340 18 335 335 373 290 2 312 3 333

Be 1.35 0.11 1.50 1.30 1.70 1.20 0.00 1.30 0.10 1.42

Cd 0.10 0.09 0.07 0.07 0.09 0.04 0.00 0.08 0.06 0.07

Ce 24.61 1.05 25.70 25.40 28.10 21.75 0.49 23.80 0.10 26.00

Co 11.50 0.73 13.40 12.60 7.70 12.90 0.42 13.57 0.45 13.10

Cr 8.91 1.14 24.10 11.40 13.80 12.40 1.56 12.00 2.15 14.00

Cs 1.56 0.10 1.50 1.50 2.10 0.66 0.18 1.37 0.15 1.49

Cu 33.30 3.16 30.80 41.00 25.80 38.20 5.94 37.60 2.19 23.80

Ga 19.50 0.81 20.00 19.40 20.00 17.70 0.14 19.00 0.00 19.90

La 12.69 0.43 13.00 12.50 14.30 11.20 0.14 11.83 0.06 13.00

Li 24.66 1.98 28.30 25.00 11.90 11.45 2.05 18.43 4.15 22.70

Mn 544 23 591 581 529 561 14 590 12 618

Mo 0.75 0.09 1.10 0.69 1.50 0.82 0.04 0.78 0.09 0.95

Nb 5.38 1.26 6.40 7.00 7.60 7.75 2.76 4.50 0.20 7.02

Ni 8.28 0.97 16.90 8.10 7.00 10.20 0.85 12.73 1.94 11.40

P 616 41 720 659 690 563 15 613 6 684

Pb 7.63 0.62 7.80 7.28 8.50 5.24 0.08 6.97 1.34 7.04

Rb 33.65 1.14 32.60 31.40 41.10 23.45 3.46 29.07 0.76 32.40

Sb 0.17 0.12 0.20 0.20 0.26 0.10 0.00 0.20 0.00 <0.02

Sc 9.38 0.64 9.90 10.80 6.00 10.25 0.35 10.50 0.50 10.50

Sr 471 11 462 462 447 457 10 473 9 489

Th 2.75 0.24 2.90 2.55 3.00 2.21 0.01 2.40 0.00 2.56

Tl 0.33 0.34 0.20 0.28 0.40 0.13 0.00 0.11 0.08 0.29

U 1.05 0.11 1.20 1.10 1.20 1.06 0.01 0.97 0.01 1.07

V 75.17 6.12 90.20 83.30 42.00 87.30 3.39 86.43 2.51 86.90

Y 11.79 0.73 12.90 11.70 13.20 12.40 0.00 13.37 0.25 13.30

Zn 60.35 3.40 68.70 66.40 64.00 66.70 6.93 70.10 1.61 65.20

ED-XRF

Zr 119 19 136 119 1 149

Table 7. Major elements and Zr by X-ray fluorescence (XRF) and selected trace-element analyses by inductively coupled plasma 

mass spectrometry (ICP-MS) for bulk samples from 2004–2006 Mount St. Helens lava dome and sample of lava erupted in October 1986.

[Oxide abundances in wt. percent, elemental abundances in parts per million. Numbers in parentheses represent number of analyses in average (AVG) and 

standard deviation (SD). Complete analyses of all samples are given in Thornber and others (2008b), and additional geochemical data for bulk analyses of 

samples from past 500 years of eruptive activity at Mount St. Helens are summarized in Pallister and others (1992).]
698 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

2004–6 dacite

Average

2004–6 dacite

glassy fragment Roof of spine 1 1986 dacite

Spine No.----- 3–7 1 1

Sample ------- SH300-1A3A SH300 (dense) SH226

INAA (ppm) AVG (5) SD (5) SD AVG (3) SD (3) SD

Rb 35.38 1.22 33.80 1.39 27.60 1.73 33.20 0.80

Sr 463 42 472 19.82 475 18.23 485 14.07

Cs 1.65 0.05 1.51 0.04 0.70 0.13 1.53 0.04

Ba 350 6 357 7.85 284 1.15 329 4.94

Th 2.61 0.07 2.51 0.04 2.17 0.01 2.53 0.03

U 1.23 0.02 1.26 0.05 1.11 0.05 1.20 0.04

La 11.92 0.16 11.60 0.10 10.80 0.00 12.10 0.08

Ce 24.36 0.94 22.90 0.23 23.13 0.15 25.30 0.23

Nd 12.84 0.47 12.80 0.45 12.20 0.20 13.30 0.39

Sm 2.80 0.09 2.78 0.02 2.67 0.02 2.92 0.02

Eu 0.85 0.01 0.89 0.02 0.87 0.02 0.90 0.02

Gd 1.01 1.38 2.56 0.29 2.50 0.06

Tb 0.35 0.02 0.34 0.01 0.37 0.01 0.39 0.01

Ho 0.41 0.02 0.43 0.08 0.46 0.01 0.47 0.06

Tm 0.06 0.09 0.16 0.02 0.18 0.00

Yb 0.99 0.04 0.98 0.03 1.06 0.03 1.12 0.02

Lu 0.15 0.01 0.15 0.004 0.15 0.00 0.16 0.003

Zr 119 15 107 10 115 2.52 143 16.16

Hf 3.10 0.10 2.95 0.04 2.99 0.08 3.28 0.04

Ta 0.35 0.01 0.37 0.01 0.36 0.01 0.40 0.01

W 0.47 0.38 0.77 0.77 0.79 0.16 0.79 0.13

Sc 8.89 0.15 9.61 0.07 10.07 0.31 9.81 0.07

Cr 8.40 0.61 12.20 0.35 12.10 0.98 10.70 0.24

Co 11.44 0.13 12.70 0.11 13.63 0.55 12.40 0.12

Ni 9.96 4.52 8.48 1.26 16.67 0.86 14.60 1.59

Zn 55.08 1.68 58.30 1.69 67.97 1.10 57.90 1.22

As 0.94 0.11 1.07 1.07 0.56 0.08 1.28 0.08

Sb 0.17 0.01 0.19 0.01 0.12 0.02 0.18 0.01

Au, (ppb) 4.84 8.87 0.82 0.49 8.83 5.96 0.56 0.33

Chondrite-normalized REE abundances

La 50.30 0.69 48.95 0.22 45.57 0.00 51.05 0.18

Ce 39.80 1.54 37.42 0.14 37.80 0.25 41.34 0.15

Pr

Nd 27.49 1.00 27.41 0.26 26.12 0.43 28.48 0.24

Sm 18.27 0.57 18.17 0.02 17.45 0.13 19.08 0.03

Eu 14.58 0.20 15.38 0.04 14.94 0.39 15.50 0.04

Gd 4.91 6.72 12.46 0.18 12.18 0.27

Tb 9.36 0.56 8.98 0.03 9.90 0.22 10.45 0.02

Dy

Ho 7.25 0.27 7.60 0.11 8.15 0.15 8.32 0.09

Tm 2.49 3.42 6.27 0.05 7.15 0.08

Yb 5.81 0.24 5.79 0.01 6.24 0.16 6.59 0.01

Lu 5.75 0.20 5.87 0.01 6.05 0.18 6.38 0.01

Table 8. Trace-element analyses by instrumental neutron activation analysis of bulk samples from 2004–2006 

Mount St. Helens lava dome and sample of lava erupted in October 1986.

[Elemental abundances in parts per million (ppm), except Au in parts per billion (ppb). Numbers in parentheses indicate number 

of analyses in average (AVG) and standard deviation (SD). Standard deviation for SH300-1A3A and SH226 are based on multiple 

counts for the same element. Dashes indicate insufficient data to calculate meaningful standard deviations. Normalization to chondrite 

abundance uses chondritic composition reported by Sun and McDonough (1989). Complete analyses of all samples are given in 

Thornber and others (2008b), and additional geochemical data for bulk analyses of samples from past 500 years of eruptive activity at 

Mount St. Helens are summarized in Pallister and others (1992).]
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 699

Grain 

No. Sample No. and eruption date Grain pair size Zoning T, °C Std. dev. fO2

Std. 

dev.

Analyses/ 

grain pairs

SH300-1A3 (1980–86) 1 U 904 4 -11.50 0.08 21/3

SH304-2A9 (10/18/04) U 850 5 -12.29 0.07 65/4

SH304-2Gg1 (10/18/04) U 844 2 -12.43 0.07 61/6

SH305-1 (11/20/04)

1 Average (40 μm) Z (MT) 883 3 -12.11 0.06 5

1 ±3 μm from contact 40 μm Z (MT) 953 -- -11.05 -- 2

2 43 μm from IL 65 μm Z (MT) 898 -- -11.48 -- 2

2 32 μm from IL 65 μm Z (MT) 903 -- -11.41 -- 2

2 26 μm from IL 65 μm Z (MT) 907 -- -11.35 -- 2

2 18 μm from IL 65 μm Z (MT) 921 -- -11.17 -- 2

2 10 μm from IL 65 μm Z (MT) 938 -- -10.95 -- 2

2 4 μm from IL 65 μm Z (MT) 928 -- -11.07 -- 2

2 Average 65 μm Z (MT) 915 13 -11.25 0.15 11

3 Average interior 20 μm Z (MT) 880 3 -12.41 0.02 10

3c ±3 μm from contact 20 μm Z (B) 998 -- -10.77 -- 2

4c weakly zoned 20 μm S (MT) 912 12 -11.71 0.13 6

5 small grain pair 8 μm Z (MT) 1019 83 -10.65 1.07 4

6 Average 40 μm U 908 2 -11.72 0.03 6

7c first MT grain (6 μm) 15 μm U 858 9 -12.79 0.25 5

7c second MT grain (4 μm) 15 μm U 935 -- -11.66 -- 4

8 Average 25 μm U 891 1 -12.09 0.04 10

9 Average 25 μm U 883 1 -12.03 0.01 9

10c Average 40 μm S (MT) 871 10 -12.50 0.19 10

11 Average 40 μm U 857 1 -12.80 0.02 10

SH305-1 (kc) (11/20/04)

1 Average 14 μm S (MT) 873 7 -12.30 0.18 5

2 Average 12 μm U 862 6 -12.53 0.15 6

SH305-2A (11/20/04)

1 Average 30 μm Z (MT) 936 23 -11.42 0.31 16

2 Average 40 μm S (MT) 791 14 -14.91 0.33 8

Table 9. Fe-Ti oxide thermobarometry results for 2004–2005 eruptive products.

[Temperatures and oxygen fugacities calculated using Fe-Ti exchange thermometer and oxygen barometer of Andersen and Lindsay 

(1988) and the solution model of Stormer (1983). Complete oxide analyses, by electron microprobe, are presented in digital appendix. 

Grain number indicates each of successive individual grains analyzed by electron microprobe; c indicates composite grain, typically 

with two or more magnetite crystals attached to single ilmenite. Grain pair size shows maximum dimension of combined grains. Zoning 

coded as follows: U, unzoned; Z (IL), zoned ilmenite; Z (MT), zoned titanomagnetite; Z (B), zoning in both oxide grains; (ZR), reversely 

zoned with lower Ti in magnetite near contact with ilmenite; (S) = slight zoning detected but inadequate to warrant separate calculations 

of temperature and oxygen fugacity (fO2); Ux = unzoned, but partially exsolved; broad beam analyses used. “?” = zoning not determined. 

Eruption dates and related uncertainty described in table 1.

Averages and standard deviation (std. dev.) calculated only for unzoned or weakly zoned grains. Dashes, insufficient data to 

determine standard deviation. For zoned grains, temperature (T) and oxygen fugacity (fO2

) calculated for individual spot analyses of 

zoned phase vs. average composition of the unzoned phase, or for near-contact single ilmenite-magnetite pairs. Column for analyses/

grain pairs refers to total number of analyses used in calculation, followed by number of grain pairs for samples with homogeneous 

distributions of oxide temperatures. Samples SH300 and SH304 yield results that are tightly clustered, reflecting primary magmatic 

equilibration conditions in the Mount St. Helens magma reservoir at depth. All other samples of the 2004–6 dacite have heterogeneous oxide mineral populations that show evidence of disequilibrium, attributed mainly to transient heating.]
700 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

3 Average 30 μm U 858 12 -13.19 0.27 11

4 <20 μm from IL 100 μm Z (MT) 944 8 -11.06 0.10 9

4 Core of MT vs. IL 100 μm “ 842 -- -12.43 -- 7

MSH05JV_1_19 (1/16/05 grains in ash) (1980–86) 1

1 Average 100 μm Ux 884 14 -11.38 0.19 9

3 Average 100 μm U 888 5 -11.61 0.09 8

4 Interior of grains 100 μm S (MT) 912 9 -11.01 0.13 8

4 5 μm from IL 100 μm S (MT) 931 -- -10.78 -- 5

5 Average 130 μm Ux 895 13 -11.46 0.17 6

6 Average 100 μm U 778 6 -13.80 0.15 10

7 Average grain interiors 120 μm Z(B) 938 8 -10.66 0.12 12

7 Average ±15 μm of contact 120 μm “ 830 13 -13.28 0.40 8

8 Average 120 μm U 873 13 -12.09 0.20 18

9 Average 70 μm U 833 6 -12.57 0.12 10

10 Average 35 μm U 750 5 -14.73 0.22 9

MSH05WS_1_19 (1/16/05 grains in ash) (1980–86) 2

9b Average grain interiors 200 μm Z(MT) 833 6 -12.57 0.12 7

9b Average ±8 μm of contact 200 μm “ 929 6 -11.27 0.12 8

SH308-3A (02/22/05)

1 Average 12 μm U 865 6 -13.04 0.01 8

2 Average 20 μm U 909 9 -11.86 0.23 9

3 Average 40 μm S (MT) 891 2 -12.08 0.06 17

4 Average 25 μm U 882 1 -12.05 0.01 11

5 Average 17 μm U 863 6 -12.50 0.10 8

6 Average 40 μm U 882 1 -12.05 0.01 13

7 Average 40 μm U 863 6 -12.50 0.09 12

SH308-3A (CT) (02/22/05)

1 Grain 3-4 40 μm ? 850 -- -13.40 -- 2

2 Grain 7-8 30 μm ? 898 -- -12.01 -- 2

3 Grain 9-10 30 μm ? 844 -- -13.26 -- 2

4 Grain f2 1-2 40 μm ? 850 -- -13.40 -- 2

5 Grain f2 3-4 30 μm ? 857 -- -12.49 -- 2

Average of above 860 22 -12.91 0.63

SH309-1 (CT) (02/22/05)

1 Grain 7-5 50 μm ? 958 -- -10.35 -- 2

2 Grain 8-9 60 μm ? 990 -- -10.07 -- 2

3-08-05 grains in ash (CT) (03/08/05)

1 DRS4g1 1-2 15 μm ? 868 -- -12.86 -- 2

2 DRS4g1 3-4 15 μm ? 887 -- -12.42 -- 2

3 DRS4g1 5-6 30 μm ? 842 -- -13.75 -- 2

4 DRS4g2 1-2 25 μm ? 846 -- -13.31 -- 2

5 DRS4g2 3-4 40 μm ? 825 -- -13.85 -- 2

6 DRS4g2 5-6 50 μm ? 828 -- -13.85 -- 2

Grain 

No. Sample No. and eruption date Grain pair size Zoning T, °C Std. dev. fO2

Std. 

dev.

Analyses/ 

grain pairs

Table 9. Fe-Ti oxide thermobarometry results for 2004–2005 eruptive products.—Continued
30. Petrology of the 2004–2006 Lava Dome—Implications for Magmatic Plumbing and Eruption Triggering 701

7 DRS4f3g1-2 50 μm ? 871 -- -12.66 -- 2

8 DRS4f51 1-2 25 μm ? 846 -- -13.51 -- 2

Average of above 852 22 -13.28 0.56

SH312-1 (03/08/05)

1 Average 40 μm Z (MT) 894 2 -11.86 0.03 11

1 ±5 μm from contact 40 μm Z (MT) 957 -- -10.98 -- 2

2 Average grain interiors 80 μm Z (MT) 871 2 -11.92 0.04 13

2 5 μm from IL 80 μm Z (MT) 898 -- -11.57 -- 7

2 12 μm from IL 80 μm Z (MT) 913 -- -11.38 -- 7

2 18 μm from IL 80 μm Z (MT) 910 -- -11.41 -- 7

2 25 μm from IL 80 μm Z (MT) 898 -- -11.57 -- 7

2 32 μm from IL 80 μm Z (MT) 890 -- -11.68 -- 7

2 40 μm from IL 80 μm Z (MT) 880 -- -11.82 -- 7

3c Average 40 μm U 878 8 -12.42 0.17 20

4c first MT grain (15 μm) 55 μm U 929 5 -11.07 0.07 11

4c second MT grain (10 μm) 55 μm U 921 1 -11.29 0.04 6

5 SH312 gr5 30 μm U 920 17 -11.33 0.33 4

8 SH312 gr5c 25 μm U 857 2 -12.47 0.05 3

6 SH312 gr6 50 μm U 903 4 -11.44 0.02 10

7 SH312 gr7 50 μm U 895 3 -11.85 0.08 12

SH314-2 (04/17/05)

1 Average 20 μm U 794 14 -14.21 0.48 5

2 Average 50 μm U 813 2 -13.51 0.01 8

3 Average grain interiors 50 μm ZR(MT) 872 3 -12.55 0.07 11

3 SH314-2 gr3 ct 50 μm ZR(MT) 840 -13.03 6

3 4 μm from IL 50 μm ZR(MT) 847 -12.92 6

3 8 μm from IL 50 μm ZR(MT) 853 -12.83 6

4 SH314-2 gr4 sm 25 μm U 794 1 -13.51 0.03 5

5 Average 45 μm U 855 11 -12.75 0.16 11

7 Average 40 μm U 764 38 -15.23 1.34 8

8 Average grain interiors 30 μm Z (MT) 736 1 -16.62 0.02 8

8 5 μm from IL 30 μm Z (MT) 815 7 -13.56 0.18 5

SH315-2 (04/01/05)

1 Average 300 μm Z (IL) 739 24 -16.29 0.83 6

1 High Fe IL core vs. avg. MT 300 μm Z (IL) 828 -13.39 4

2 Average grain interiors 60 μm Z (MT) 838 1 -13.79 0.05 8

2 4 μm from IL 60 μm Z (MT) 805 1 -14.34 0.07 9

3 Average 50 μm S (IL) 807 9 -14.23 0.27 7

4 Average 25 μm U 791 33 -14.75 1.14 6

5c Average 30 μm U 742 98 -16.40 3.81 9

6 Average 30 μm U 817 3 -13.86 0.12 7

7 Average 35 μm U 813 40 -14.08 1.22 7

Grain 

No. Sample No. and eruption date Grain pair size Zoning T, °C Std. dev. fO2

Std. 

dev.

Analyses/ 

grain pairs

Table 9. Fe-Ti oxide thermobarometry results for 2004–2005 eruptive products.—Continued

1

 Crater floor debris dragged up by extruding spines but that probably originated by dome growth in the period 1980–86.

2

 Emitted as ash during explosion that disrupted crater-floor debris adjacent to spine 4.
702 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Oxide Analyses used in Thermobarometic Calculations

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

The database for oxide analyses described in this chapter is tabulated in a spreadsheet file.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 31

1

 Geological Sciences Department, Brown University, 324 Brook Street, 

Providence, RI 02912

Magmatic Conditions and Processes in the Storage Zone 

of the 2004–2006 Mount St. Helens Dacite

By Malcolm J. Rutherford1 and Joseph D. Devine III1

Abstract

The 2004–6 eruption of Mount St. Helens produced 

dacite that contains 40–50 volume percent phenocrysts of 

plagioclase, amphibole, low-Ca pyroxene, magnetite, and 

ilmenite in a groundmass that is nearly totally crystallized. 

Phenocrysts of amphibole and pyroxene range from 3 to 5 

mm long and are cyclically zoned, with one to three alternations of Fe- and Al-rich to Mg- and Si-rich layers showing 

little indication of phenocryst dissolution between zones. 

Similar-size plagioclase phenocrysts also contain several 

cyclic zones ranging between ~An68 and An45–35. Textural evidence indicates that amphibole, pyroxene, and ilmenite began 

to crystallize before the most An-rich plagioclase. Magnetite 

and ilmenite phenocrysts are small (less than 100 μm), vary 

somewhat in composition from grain to grain, and are sporadically zoned. Magnetite-ilmenite pairs yield temperatures 

of equilibration ranging from 820oC to 890oC and fO2 values 

of NNO +1 log unit. Magnetite compositions suggest that the 

2004–6 magma was formed by mingling of magmas less than 

5–8 weeks before eruption and that the magma last equilibrated within this temperature range. The amphibole phenocryst zoning involves approximately equal amounts of a 

pressure-sensitive Al-Tschermak molecular substitution and 

a temperature-sensitive edenite substitution in one cycle of 

growth. Hydrothermal experiments done on the natural dacite 

show that crystallization of the Fe- and Al-rich amphibole 

end member requires pressures of 200–300 MPa at temperatures of 900oC, conditions approaching the upper temperature limit of amphibole stability. The dacitic magma crystallizes the An68 plagioclase when the pressure drops to 200 

MPa at 900oC. The magma must cool at this depth to produce 

a complete An68–An40 plagioclase zone and a Mg-rich layer 

on the amphiboles before the magma is cycled back to a high 

pressure, when a new layer of Fe-rich amphibole is acquired. 

The amphibole crystallizing in the dacite experiments at less 

than 200 MPa is lower in aluminum than any compositions 

in the natural cyclically zoned phenocrysts. The outer rim on 

some 2004–6 amphibole phenocrysts appears to have formed 

in the 100–200 MPa range, as do some phenocrysts in the 

May 1980 dacite pumice. Plagioclase rims of An35 in the 

2004–6 magmas indicate that phenocryst growth continued 

until the pressure decreased to 130 MPa and that ascent was 

slow until this depth. Magma then entered the conduit for a 

relatively rapid ascent to the surface as indicated by the very 

thin (less than 5 μm) decompression-induced rims on the 

amphibole phenocrysts.

Introduction

Mount St. Helens began to erupt in 2004 after 18 years 

of relatively shallow seismic activity and no eruptions, aside 

from a few gas explosions during 1989–91 (Mastin, 1994). 

The extrusion of a lava dome onto the 1980 crater floor during 

the 1980–86 time period marked the end of the eruption that 

began in 1980. However, seismic data collected during the past 

two decades (Moran, 1994; Moran and others, this volume, 

chap. 2) showed a continuous, moderate level of small-magnitude seismic activity directly beneath the volcano, particularly in the depth range of 3–12 km. In addition, there were 

significant increases in the level of seismic activity at 3–10-km 

depths in 1995, 1998, and 2002, suggesting the possibility 

that magma was ascending to a depth of 3 to 4 km beneath 

the crater floor at these times. In October 2004, following a 

brief period of steam-and-ash explosions, lava began erupting within the crater at the south margin of the 1980–86 lava 

dome, and a new lava dome has been building slowly from 

a steady eruption of dacite magma (Pallister and Thornber, 

2005; Pallister and others, this volume, chap. 30).
704 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Samples of the new dome collected by the U.S. Geological Survey’s Cascades Volcano Observatory (USGS–CVO) 

scientists in late October 2004 were very oxidized, reddish in 

color, and crystal rich; they are considered likely to be material left in the upper conduit from the 1986 eruption. Samples 

that were collected within 1–2 weeks of eruption in late 

November 2004 (SH304), in early January 2005 (SH305), and 

more recently are interpreted to be samples of new magma. 

Initially, it appeared that this new magma was simply remobilized 1980–86 magma, stored at some level, possibly at 3–4 

km, as suggested by the seismic record. In support of this idea, 

the bulk compositions determined for the new lava (Pallister 

and others, this volume, chap. 30) show it to be only slightly 

more evolved than lava from the 1980–86 eruptions (65 rather 

than 63 weight percent SiO2). However, some characteristics 

of the new lava samples pointed to a new magma source at 

depth, or at least reestablishment of the 1986 magma storage 

system at 7–12 km depth (Scandone and Malone, 1985). These 

characteristics included the presence of apparently stable 

amphibole phenocrysts, the presence of significant amounts of 

glass in many samples, and, most particularly, the identification of magnetite-ilmenite pairs yielding temperatures >900oC 

(Pallister and Thornber, 2005; Rutherford and Devine, 2005). 

If the erupting magma had been stored at pressures less than 

100 MPa (~4 km below the summit; Geschwind and Rutherford, 1995) then amphibole phenocrysts would have developed 

a thick reaction rim according to earlier work (Rutherford 

and Hill, 1993). Amphiboles in the 2004–6 magma generally 

lack evidence of any significant breakdown except for a thin 

(less than 5 μm) rim that is attributed to reaction with melt 

during ascent in the conduit (Rutherford and Devine, 2005). 

The greater than 900oC temperatures from the 2004–6 oxides 

are higher than temperatures (860oC) recorded from samples 

of the 1986 Mount St. Helens eruption (Rutherford and Hill, 

1993), suggesting that hotter magma was involved in the 

2004–6 eruptions.

The observations and the initial analyses of the 2004–6 

lava samples (Pallister and Thornber, 2005) give rise to a 

number of questions that we have attempted to answer in this 

paper: (1) Where did phenocrysts in these magmas last equilibrate with the surrounding melt before entering the conduit 

and undergoing transport to the surface? This question can be 

addressed by studying the rim compositions of the phenocryst 

phases in the natural samples and by doing experiments 

designed to reproduce the phenocryst rim-melt equilibrium at 

different possible preeruption conditions. (2) Can the cyclic 

compositional zoning observed in all of the silicate phenocryst 

phases (plagioclase, amphibole and orthopyroxene) of the 

2004–6 magma (Rutherford and Devine, 2005) be explained 

by crystallization of a dacite magma at different conditions, 

and, if so, what are those conditions? We have studied the 

phenocryst zoning analytically and experimentally to determine the range of conditions experienced by the magma over 

its preeruption crystallization history, and herein we use the 

data to create a model of magma dynamics within this subvolcanic system. (3) Finally, what are the preeruption conditions 

for the 2004–6 magma, and how do they compare with those 

determined for the 1980–86 magma samples (Rutherford and 

others, 1985; Rutherford and Devine, 1988)? What are the 

possible reasons for any differences observed?

Analytical and Experimental Methods

Samples of the new 2004–6 Mount St. Helens dacite 

dome were supplied by members of the USGS–CVO staff 

shortly after it became possible to collect them. Polished 

thin sections were made for mineral-chemistry and textural 

investigations of the different samples, and crushed powders 

were made of some samples for use in hydrothermal experiments. The phases in the natural samples were analyzed using 

a CAMECA SX100 microprobe for the crystalline phases 

and a CAMECA Camebax microprobe for glasses. Analytical methods used are identical to those of Rutherford and 

Devine (2003). Amphibole cation proportions were calculated 

using the method of Holland and Blundy (1994). The method 

of Cosca and others (1991) was used to calculate amphibole 

stoichiometry on the basis of 13 cations (Si, Al, Fe, Mg, Mn, 

and Ti) and to estimate ferric iron contents.

Experiments were performed on sample SH305-1 (erupted 

in November 2004), which was crushed to form a mixture of 

matrix and small (less than 0.2 mm) phenocryst fragments. 

This crushed sample was used in hydrothermal melting experiments and to create a glass-rich starting material for use in 

crystallization experiments. Melting and crystallization experiments were done in adjacent sealed tubes. The objective was 

to reversibly create new phenocryst growth under controlled 

conditions of pressure, temperature, and oxygen fugacity (P, 

T, and fO2

). All of the experiments were done with PH2

O = Ptotal

using previously described methods. The textural data indicate 

that amphibole crystallized before plagioclase in the 2004–6 

magma, and this requires that the melt contain at least 4 weight 

percent dissolved H2O (Merzbacher and Eggler, 1984) at this 

early stage of crystallization. With additional crystallization 

(primarily plagioclase and orthopyroxene), the melt would have 

rapidly approached water saturation for total pressures less than 

300 MPa. The possibility that the 2004–6 magma initially contained some dissolved CO2 cannot be ruled out, however, and is 

considered in the discussion.

Petrology of the 2004–2006 Dacite

General Petrology

The 2004–6 Mount St. Helens dacite is macroscopically 

similar to the material erupted in 1980–86. It is a crystalrich dacite (~45 volume percent phenocrysts on a bubblefree basis) with phenocrysts of plagioclase, orthopyroxene, 

amphibole, Ti-magnetite, and ilmenite in a groundmass that 

is nearly totally crystallized in most of the samples (fig. 1; 
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 705

Pallister and Thornber, 2005; Rutherford and Devine, 2005). 

Phenocrysts range up to 6 mm across and sporadically 

show some preferred orientation owing to magma flow. The 

groundmass and phenocryst assemblages are the same, except 

that amphibole is lacking in the groundmass and, instead, 

there is Ca-rich pyroxene and minor cristobalite and quartz. 

The bulk compositions of the 2004–6 magma samples have 

been remarkably uniform at 65 weight percent SiO2 throughout the eruption (Pallister and others, this volume, chap. 30). 

The erupting magma resembles the magma that erupted in 

1980–86 in that it contains several volume percent (fragments 

and crystal clots) of “gabbroic” material that is rich in plagioclase, pyroxenes, and amphibole (Heliker, 1995). As in the 

earlier eruption, a few of the inclusions in the 2004–6 samples 

contain interstitial melt (glass) that appears to represent partial melting, but many inclusions do not contain a glass phase. 

The smallest inclusions are difficult to distinguish macroscopically from phenocryst clots, but their amphibole differs 

from the phenocrysts compositionally and generally is not 

chemically zoned. Additionally, Ca-rich pyroxene commonly 

is associated with amphibole in the inclusions (Rutherford 

and Devine, 2005).

Plagioclase

The plagioclase phenocrysts in the Mount St. Helens 

2004–6 magma are euhedral, complexly zoned crystals (fig. 

1) as large as 5 mm in diameter, although most are less than 

3 mm. These phenocrysts have a complex core-to-rim cyclic 

zoning easily observable in backscattered electron (BSE) 

images (fig. 2). The composition of the outer plagioclase rim 

ranges from An60 to An35, but the great majority of the rims 

are relatively albite rich (An30–35). Core compositions of the 

plagioclase phenocrysts range up to An68. Plots of composi1 mm

Plg

Hb

Hb

Opx

Plg

Mt

A B

Opx

M.I.

200 μm

Figure 1. Photomicrographs showing thin section of typical Mount St Helens dacite assemblage. Phenocryst phases are 

plagioclase (Plg), amphibole (Hb), low-Ca pyroxene (Opx), all with obvious cyclic zoning, and magnetite (Mt); ilmenite is other

phenocryst, not shown. Groundmass is largely crystallized and slightly vesicular. A, Plane-polarized transmitted light. B, 

Backscattered electron image. M.I., melt inclusion.

tion versus distance (rim to rim) across typical phenocrysts in 

the 2004–6 lava are shown in figure 2, and the compositions 

of a typical core and rim are given in table 1. Most boundaries 

between compositional zones are parallel to growth faces, with 

relatively little evidence of resorption involved in the zoning 

reversals (fig. 2). However, some resorption did occur during 

the plagioclase growth history, and 5 to 10 volume percent 

of the phenocrysts, including both small and large crystals, 

have a strongly sieved core with a relatively thin rim of clear 

plagioclase. A similar population of sieve-cored plagioclase 

phenocrysts was also observed in 1980–86 lava (Rutherford 

and others, 1985). Another characteristic of the plagioclase 

phenocrysts in the new lava is the relatively common presence of included crystals of amphibole, orthopyroxene, and 

Ti-magnetite, in decreasing order of abundance, particularly in 

the outer half of the phenocryst.

Amphibole

Amphibole phenocrysts in the Mount St. Helens dacite are euhedral, mostly 1–3 mm long but infrequently up 

to 5 mm in length, and commonly exhibit an internal cyclic 

compositional zoning parallel to growth faces (fig. 3). Petrographic examination shows some variability in amphiboles 

from samples erupted at different times during the 2004–6 

eruption, particularly different amounts of opacitization due 

to variable syneruptive oxidation (Garcia and Jacobson, 1979; 

Rutherford and Hill, 1993) or decompression-induced rim 

development. The outer margins of some amphiboles appear 

somewhat rounded and possibly eroded or abraded (fig. 3C); 

whereas evidence for rounding or dissolution at internal zone 

boundaries is inconclusive (fig. 3). Inclusions in amphiboles 

are relatively uncommon compared to plagioclase phenocrysts, 

but magnetite and sulfide inclusions do occur. Larger inclu-
706 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Plagioclase phenocrysts

Sample No.--- SH304c SH304r SH305c SH305r SH323r SH323c

SiO2 53.54 60.19 52.16 58.69 58.46 51.86

Al2O3 29.56 25.16 29.86 26.73 26.40 30.81

Fe2O3 0.30 0.24 0.59 0.33 0.32 0.63

CaO 12.21 6.71 13.44 8.04 7.91 13.54

Na2O 4.33 6.85 3.31 6.33 6.46 3.37

K2O 0.11 0.25 0.17 0.25 0.24 0.08

Total 100.05 99.40 99.53 100.37 99.79 100.29

An content 60 35 69 40 39 68

Orthopyroxene phenocrysts

Sample No.--- SH304

Px1

SH304

Px2

SH305

Px1

SH305

Px2

SH323

Opx 1

SH323

Opx 2

SiO2 52.60 52.56 52.17 53.45 52.14 53.09

TiO2 0.22 0.13 0.17 0.12 0.21 0.29

Al2O3 1.56 0.54 0.44 0.56 1.37 1.74

FeO 21.07 23.50 23.28 20.47 22.60 16.97

MgO 22.30 20.80 20.85 23.65 20.40 25.23

CaO 1.16 1.15 1.21 0.86 1.20 0.99

MnO 0.70 0.69 0.63 0.71 0.69 0.41

Cr2O3 0.00 0.01 0.02 0.01 0.01 0.06

Total 99.61 99.38 98.77 99.83 98.62 99.78

Mg# 63 60 60 63.5 61 64

Table 1. Compositions of representative natural plagioclase and orthopyroxene phenocrysts in dacite 

erupted at Mount St. Helens, Washington, 2004–2006.

[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Compositions in weight percent; all Fe as FeO in pyroxene and as Fe2O3 in plagioclase. Magnesium number, Mg# = (Mg×100)/

(Mg+(Fe2+ in octahedral site)).]

sions in amphibole are generally multicrystalline and appear 

to represent trapped melts that have crystallized in what 

appears to have been a hollow-cored crystal (figs. 1, 3A–C). 

The amphiboles that occur in gabbroic inclusions are commonly not in contact with glass, except on the inclusion 

margins. Amphibole crystals in these inclusions are also 

generally anhedral, do not show cyclic compositional zoning, 

and are low in Al2O3 (8–11 weight percent) compared to the 

phenocrysts. Some isolated crystals in the dacite magma have 

similar characteristics.

The cyclic internal zoning present in the majority of 

amphibole phenocrysts in the 2004–6 Mount St. Helens lava 

is present in even the smallest (~100 μm) phenocrysts and is 

observable optically, as well as in BSE images. This zoning 

tends to begin with a bright BSE zone in the center of a phenocryst, involves one or two cycles of alternating bright-dark 

zoning, and ends with a dark, low-FeO zone adjacent to the 

groundmass (fig. 3). A similar zoning occurs in the amphiboles from the 1986 lava dome (fig. 3D), although the outer 

rim in these samples is intermediate to Fe-rich rather than 

Mg-rich. The Mg# (Mg/(Mg+Fe2+)) and VIAl profiles across 

representative amphibole phenocrysts are shown in figure 3, 

and compositions at points across typical cycles of amphibole 

growth are given in table 2.

The cyclic compositional zoning in the Mount St. Helens 

amphiboles involves Fe-, Al-, Na-, and K-enriched bright zones 

(BSE images) that alternate with, and grade into, Mg- and 

Si-rich dark BSE zones (figs. 3, 4). The cation substitutions 

involved in this zoning are similar to those observed in phenocrysts from the 1995–2002 Soufrière Hills andesite erupted 

on Montserrat (Rutherford and Devine, 2003) and in Fish 

Canyon latite (Bachmann and Dungan, 2002), but they differ 

in one significant aspect. A pressure-sensitive Al-Tschermak 

substitution was not inferred for either the Soufrière Hills or the 
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 707

A

B

A

B

Plg1

SH305

A B

SH304

Plg 1

100 μm

30

35

40

45

50

55

60

65

0 40 80 120 160 200 240 280 320

SH305.PLG.1.TRANSECT

DISTANCE, IN MICRONS

A B

20

30

40

50

60

70

0 40 80 120 160

SH304.PLG.1.TRANSECT

DISTANCE, IN MICRONS

A B

200 μm

An CONTENT, IN MOLE PERCENTAn CONTENT, IN MOLE PERCENT

Figure 2. Backscattered electron images from thin section of Mount St. Helens lava, showing typical plagioclase 

phenocryst of 2004–6 lava. Images show cycles of light (Ca-rich) and dark (Na-rich) phenocryst growth; accompanying 

graphs depict An concentrations along the A–B profiles. A, Plagioclase in sample SH304. B, Plagioclase in sample SH305.

Fish Canyon amphibole zoning. In contrast, phenocrysts in the 

2004–6 Mount St. Helens lava are marked by compositional 

variations that are ~50 mol percent Al-Tschermak substitution (M(Mg, Fe) + TSi = MAl + T

Al; fig. 5A), with the remaining 

change in TAl attributed to the temperature-dependent edenite 

(

TSi + Avacancy = T

Al + A(Na+K)) and Ti-Tschermak exchanges 

(2TSi + MMn = 2TAl + MTi), as shown in figures 5B and 5C. In 

going from the Fe-rich to the Mg-rich zones, the decrease in 

Fe appears to be completely balanced by the increase in Mg 

in the octahedral position (fig. 6), as was found to be the case 

in previous studies (Bachmann and Dungan, 2002; Rutherford 

and Devine, 2003). In other words, the increase in VIAl in the 

amphibole structure is accompanied by an increase in Fe2+ at the 

expense of Mg2+ in the octahedral site (fig. 3).

One potentially important aspect of the amphibole 

phenocryst zoning in the 2004–6 Mount St. Helens magma 

is illustrated by the zoning in sample SH323 (fig. 7). Most 

amphibole phenocrysts in this sample have, in contrast to 

2004–6 magma erupted earlier, a relatively bright (Fe-rich) 

rim, but the Al2O3 content is low and the SiO2 content (fig. 4)

is relatively high in this layer. Analytical electron-microprobe 

traverses across such phenocrysts show that the fluorine (F) 

content increases significantly over a distance of 20–40 μm at 

the phenocryst margin as Si increases and Al decreases. The 

core of the amphibole phenocryst contains 960±200 ppm F; 

the very outer margin of the phenocryst contains as much as 

7,700 ppm (fig. 7C; table 2). A similar increase in F and Si 

was noted where the rim-to-rim analytical profile passed close 

to a melt channel in the phenocryst. Another, possibly related, 

observation is that the amphiboles in the gabbroic inclusions 

also are low in octahedral as well as total Al and Si relative 

to the cyclically zoned phenocrysts, and they are somewhat 

enriched in F (2,500 ppm) relative to the cyclically zoned 

phenocrysts (average of 900 ppm F). The presence of a thin, 

decompression-induced reaction rim at the contact of these 

phenocrysts with the groundmass indicates that the fluorine 

enrichment occurs before the final magma ascent and decompression. One possible explanation of the fluorine zoning is 

that these crystals and crystal rims represent a partial recrystallization of amphiboles at conditions of low P and high activity 

of SiO2 just outside the OH-bearing amphibole stability field. 

Similar overgrowths are observed on amphibole phenocrysts 
708 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample No.---

Site------------

Character-----

SH304-1-2

Hb7 rim

Hi-Mg

SH304-1

Hb7 interior

Med.-Al

SH304-1

Hb7 core

Hi-Al

SH305a

Hb rim

Low-Al

SH305b

Core

Hi-Al

Major-element analyses, weight percent

SiO2 44.80 42.86 40.97 43.51 40.40

TiO2 1.95 2.41 2.56 2.72 2.40

Al2O3 11.40 13.09 14.30 11.97 14.86

FeO 12.30 12.90 15.31 11.99 16.05

MgO 15.43 13.92 11.82 14.97 10.54

CaO 11.29 10.67 10.74 11.58 11.39

Na2O 1.98 2.37 2.39 2.39 2.26

K2O 0.20 0.30 0.31 0.32 0.35

MnO 0.16 0.15 0.21 0.10 0.16

Cr2O3 0.02 0.01 0.01 0.00 0.00

Total 99.51 98.92 98.61 99.55 98.41

Cation abundance in structure formula

Si 6.305 6.128 5.942 6.170 5.923

IVAl 1.695 1.872 2.058 1.830 2.077

Ti 0.206 0.259 0.279 0.290 0.265

VIAl 0.196 0.333 0.386 0.171 0.490

Cr 0.007 0.001 0.008 0.000 0.000

Fe3+ 0.893 0.802 0.842 0.725 0.681

Fe2+(Y) 0.468 0.639 0.938 0.649 1.261

Mg 3.237 2.966 2.555 3.164 2.303

Ca 1.702 1.634 1.669 1.760 1.809

Fe2+(X) 0.087 0.101 0.077 0.047 0.026

Mn 0.019 0.018 0.026 0.012 0.020

Na 0.541 0.657 0.673 0.676 0.642

K 0.036 0.055 0.057 0.058 0.065

Mg# 87.37 82.27 73.14 82.98 64.62

Fe3+/ΣFe 0.617 0.520 0.453 0.510 0.346

Table 2. Representative natural amphibole compositions in dacite erupted at Mount St. Helens, Washington, 

2004–2006.

[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Compositions 

in weight percent; all Fe as FeO; n.d., not determined. Structure formula calculated after Holland and Blundy (1994). Magnesium 

number, Mg# = (Mg×100)/(Mg+(Fe2+ in Y site)).]

erupted later in 2005 (SH324) and in 2006 (SH328), when the 

magma mass-eruption rate was also relatively low (Pallister 

and others, this volume, chap. 30).

Low-Ca Pyroxene (Orthopyroxene)

The orthopyroxene phenocrysts present in the 2004–6 

magma are commonly cyclically zoned parallel to growth surfaces, similar to the zoning observed in amphibole. Examples 

where the zoning stands out as alternating light and dark bands 

are present in figures 1B, 3A, and 3B. Compositional profiles 

across a typical zoned phenocryst are shown in figure 8 for 

the pyroxene that is present in figure 1B. Like the amphibole, pyroxene phenocrysts can have as many as four cycles 

of light and dark zones rich in Fe and Mg, respectively, but 

many have only two. The smaller number of cycles visible in 

BSE and compositional profiles across orthopyroxene phenocrysts might be related to their smaller size, although some 

100-μm-diameter crystals have as many as three light-dark 

growth cycles (fig. 3A). The trend from cycle to cycle is similar 

in both amphibole and orthopyroxene, in that the transition from 

dark, Mg-rich to bright, Fe-rich bands is sharp, whereas the 

bright to dark transition is gradational (figs. 1, 3). This similar-
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 709

Sample No.------

Site---------------

Character--------

SH308-1

#1 rim

Low-Al

SH308-1

#3 core

Med.-Al

SH323

Rim

Low-Al

SH323

Near rim

Med.-Al

SH323

27 core

Hi-Al

Major-element analyses, weight percent

SiO2 43.89 41.53 43.88 41.15 42.08

TiO2 2.43 2.27 2.56 2.88 2.22

Al2O3 11.38 13.31 10.25 13.88 14.11

FeO 12.48 14.97 14.69 14.65 14.79

MgO 15.15 12.12 13.69 12.23 12.65

CaO 10.71 11.02 11.07 11.40 10.74

Na2O 2.32 2.25 2.46 2.96 2.83

K2O 0.23 0.35 0.31 0.27 0.27

MnO 0.08 0.23 0.21 0.14 0.23

Cr2O3 0.06 0.01 0.01 0.00 0.00

F n.d. n.d. 0.77 0.102 0.096

Total 98.73 98.06 99.90 99.66 100.02

Cation abundance in structure formula

Si 6.251 6.054 6.375 5.957 6.017

IVAl 1.749 1.946 1.625 2.043 1.983

Ti 0.260 0.249 0.280 0.314 0.239

VIAl 0.162 0.341 0.130 0.325 0.395

Cr 0.007 0.001 0.001 0.000 0.000

Fe3+ 0.827 0.807 0.716 0.584 0.777

Fe2+(Y) 0.528 0.968 0.909 1.139 0.893

Mg 3.216 2.633 2.964 2.639 2.696

Ca 1.634 1.721 1.723 1.768 1.645

Fe2+(X) 0.131 0.049 0.160 0.051 0.098

Mn 0.010 0.028 0.026 0.017 0.028

Na 0.641 0.636 0.693 0.831 0.785

K 0.042 0.065 0.057 0.050 0.049

Mg# 85.89 73.12 76.53 69.86 75.12

Fe3+/ΣFe 0.556 0.442 0.401 0.329 0.440

Table 2. Representative natural amphibole compositions in dacite erupted at Mount St. Helens, Washington, 

2004–2006.—Continued

[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Compositions 

in weight percent; all Fe as FeO; n.d., not determined. Structure formula calculated after Holland and Blundy (1994). Magnesium 

number, Mg# = (Mg×100)/(Mg+(Fe2+ in Y site)).]
710 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

HB#7

Opx

Plg

Mt A

B

SH304 SH305

HB-1

B

A

Opx

A B

30

40

50

60

70

80

90

100

0.2

0.4

0.6

0.8

1

0 40 80 120 160

SH304.HB7.13CT

DISTANCE, IN MICRONS

Mg#

VIAl

A B

30

40

50

60

70

80

90

0

0.2

0.4

0.6

0.8

1

1.2

0 40 80 120 160

SH305-HB1.13CT.

DISTANCE, IN MICRONS

Mg#

VIAl 

A B

M.I.

M.I.

M.I.

200 μm 100 μm

Mt

VIAl, IN Y SITE

Mg# = Mg/(Mg+Fe2+), IN Y SITE

Mt—Magnetite

Plg—Plagioclase

Opx—Orthopyroxene

EXPLANATION

M.I.—Melt inclusion

A

B

SH-308

Hb-1

SH-205

Hb-1

A B

C D

SH205

 HB 1 SH308

 HB 1

60

65

70

75

80

85

90

0

0.2

0.4

0.6

0.8

1

0 20 40 60 80 100

SH308.HB1

DISTANCE, IN MICRONS

Mg#

VIAl

A B

40

50

60

70

80

90

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0 40 80 120 160

SH205-1

DISTANCE, IN MICRONS

Mg#

VI A Al

B

M.I.

VIAl, IN Y SITE

Mg# = Mg/(Mg+Fe2+), IN Y SITE

Figure 3. Backscattered electron images of typical amphibole and orthopyroxene phenocrysts; 

accompanying graphs show compositional profiles along lines A–B in each image. Note growth cycles 

of alternating Fe-rich (bright) and Mg-rich (dark) zoning. Compositional profiles show the Fe-rich zones 

are also rich in octahedral-coordinated Al (VIAl). A, Sample SH304. B, Sample SH305. C, Sample SH308. D, 

Sample SH205 from 1986 dacite lava dome.
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 711

10

11

12

13

14

15

9 10 11 12 13 14 15 16

MgO, IN WEIGHT PERCENT

SH305

SH304

SH323

C

39

40

41

42

43

44

45

SiO2, IN WEIGHT PERCENT FeO, IN WEIGHT PERCENT Al2O3, IN WEIGHT PERCENT

A

10

11

12

13

14

15

16

17

18

2004–6 amphibole

2004–6 amphibole

2004–6 amphibole

B

0.15

0.2

0.25

0.3

0.35

0.4

0.45

A

0.2

0.22

0.24

0.26

0.28

0.3

0.32

B

Ti, IN Y SITE Na+K, SUM IN A SITE

0.35

1.5 1.6 1.7 1.8 1.9 2.0

0.4

0.45

0.5

0.55

SH305.HB.1

Lines 1 and 2, 13CT

C

IVAl, IN Z SITE

VIAl, IN Y SITE

Figure 4. Chemical variation diagrams showing rim-to-rim 

compositional zoning in typical amphibole phenocrysts from 

three samples of the 2004–6 eruption. A, MgO vs. SiO2. B, MgO 

vs. FeO. C, MgO vs. Al2O3. Crystal rims in SH323 (five points) with 

anomalously high SiO2 (panel A) and low Al2O3 (panel C) are 

discussed elsewhere in the paper.

Figure 5. Plots showing position of chemical species in 

crystallographic sites versus tetrahedrally coordinated Al (IVAl) in 

Z site for typical, cyclically zoned amphibole in lava samples from 

Mount St. Helens 2004–6 eruption. A, Octahedrally coordinated Al 

(

VIAl) in the Y site. B, Ti in the Y site. C, Sum of Na+K in the A site. 

Note that approximately equal amounts of the IVAl increase are 

balanced by increases in VIAl and the sum of Na and K.

ity in compositional zoning implies that the two minerals were 

co-crystallizing as P–T conditions in the magma went through a 

series of cycles. As in the amphibole, the Al and Ti concentrations in orthopyroxene are slightly higher in the Fe-rich zones, 

but the correlation is not strong (fig. 8). The lack of a correlation 

of Fe with Al in pyroxene is probably due in part to the fact that 

the Al2O3 content is low (0.5–1.5 weight percent). Analyses of 

representative pyroxene phenocrysts are given in table 1.

Fe-Ti Oxides

Subhedral to rounded Fe-Ti oxide crystals (<150 μm rim 

to rim) are present in all samples studied. Titanomagnetite 

phenocrysts in dome lavas characterized by relatively rapid 

volumetric extrusion rates (6.5 m3/s; Pallister and others, this 

volume, chap. 30) are commonly homogeneous in composition from core to rim, although TiO2 contents may vary 
712 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0.5

0.6

0.7

0.8

0.9

1

2.6 2.7 2.8 2.9 3.0 3.1 3.2

SH305.HB.5

Lines 1 and 2

Mg2+, IN Y SITE

Fe2+, IN Y SITE

1σ error

Figure 6. Plot of Mg2+ versus Fe2+ for the same amphibole 

phenocryst data plotted in figure 5, illustrating 1:1 correlation of Fe 

and Mg in the Y site of cyclically zoned 2004–2006 phenocrysts.

Figure 7. Typical amphibole phenocryst in sample SH323 

erupted in September 2005 when the mass eruption rate had 

dropped to about 1 m3

/s (Pallister and others, this volume, chap. 

30). A, Backscattered electron image showing compositional 

traverse as line A–B. B, C, Compositional traverses, showing 

presence of a rim less than 30 μm thick with low Al and Mg (panel 

B) and high F (panel C). Profile for SiO2 across same phenocryst 

(fig. 4) shows the rim is also enriched in Si.

500 μm

50

60

70

80

90

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.8

0 50 100 150 200

SH323-2.HB1

DISTANCE, IN MICRONS

VIAl

Mg#

B

A B

0

0.2

0.4

0.6

0.8

1

1

1.2

1.4

1.6

1.8

2

2.2

0 50 100 150 200

SH323-2.HB1

DISTANCE, IN MICRONS

F 

IVAl

C

A

B

A

Mg#, IN Y SITE F, IN WEIGHT PERCENT

IVAl, IN Z SITE VIAl, IN Y SITE

substantially from grain to grain, even within the same hand 

sample (fig. 9). There is no apparent correlation between TiO2 

content and grain size (fig. 9). Lava samples characterized by 

relatively low volumetric extrusion rates have titanomagnetite 

crystals that contain lamellae of titanohematite that occur in 

trellis-like intergrowths with titanomagnetite that is higher in 

TiO2 content than the presumed parent. Titanohematite grains 

in the groundmass of such oxidized samples impart a reddish 

color to the rocks.

Subhedral to anhedral ilmenite grains (typically less 

than 50 μm), though rare, are present in every sample. Many 

grains appear to be embayed to a greater or lesser extent. The 

TiO2 contents of the ilmenite grains in rapidly extruded lavas 

are essentially bimodally distributed, with one group somewhat higher in TiO2 (~44 weight percent) and lower in FeO* 

(~48 weight percent, all iron as FeO) than the other (TiO2

~42 weight percent; FeO* ~50 weight percent). Analyses of 

representative phenocrysts are included in table 3. Individual 

grains are essentially homogeneous in composition except for 

the outermost rim, which may be somewhat higher in TiO2

than the cores. In some slowly extruded samples, near-surface 

oxidation processes have resulted in decomposition of the 

ilmenite and its replacement by veinlets of titanohematite, 

interspersed with rare small patches of a TiO2 polymorph 

(greater than 83 weight percent TiO2), likely rutile.

Rare ilmenite-titanomagnetite in-contact pairs are present 

in all samples examined. Concentration gradients for TiO2 in 

titanomagnetite crystals in contact with ilmenite are commonly observed in these oxide pairs, with high TiO2 adjacent 

to the ilmenite. Analyses of coexisting oxide pairs are given 

in table 3, together with the temperature and oxygen fugacity 

indicated by the coexisting compositions.

Preeruption Magma Conditions 

Inferred from the Phenocrysts

The compositions of the coexisting Fe-Ti oxides in the 

different samples of the 2004–6 Mount St. Helens dacite have 

been used to obtain estimates of temperature and fO2 of the 

magma during the period of last magnetite-ilmenite equilibration. Methods of conversion are those described by Devine 
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 713

0.12

0.16

0.2

0.24

0.28

0.4

0.6

0.8

1

1.2

1.4

1.6

0 10 20 30 40 50 60 70

SH304.OPX13

B

TiO2

Al2O3

20

21

22

23

24

20

21

22

23

24

0 10 20 30 40 50 60 70

SH304.OPX13

A

FeO

MgO

DISTANCE, IN MICRONS

MgO, IN WEIGHT PERCENT

FeO, IN WEIGHT PERCENT

TiO2, IN WEIGHT PERCENT

Al2O3, IN WEIGHT PERCENT

8

9

10

11

12

13

14

0 20 40 60 80

SH315 TITANOMAGNETITE PHENOCRYSTS

DISTANCE RIM TO RIM, IN MICRONS

TiO2, IN WEIGHT PERCENT TiO2, IN WEIGHT PERCENT

B

7

8

9

10

11

12

13

14

0 20 40 60 80

SH305 TITANOMAGNETITE PHENOCRYSTS

A

Figure 8. Compositional profiles (rim to rim) across typical zoned 

orthopyroxene in sample SH304 (phenocryst imaged in fig. 1B). 

Note tendency of Ti and Al to be correlated with Fe in the pyroxene 

growth cycles as they are in amphibole (figs. 5, 6). A, MgO and 

FeO. B, TiO2 and Al2O3.

Figure 9. Composition (TiO2) profiles across representative 

titanomagnetite phenocrysts in samples SH305 (A) and SH315 

(B) from the 2004–6 eruption, showing range of relatively 

homogeneous crystals present.

and others (2003) and utilize the algorithm of Andersen and 

Lindsley (1988) as amended by Andersen and others (1993). 

However, meaningful geothermometry based on analysis of 

ilmenite-titanomagnetite (two-oxide) pairs in these samples is 

hampered by the common presence of TiO2 concentration gradients in the titanomagnetite grains, especially near the contact 

with ilmenite. Similar occurrences in the 1995–2003 Soufrière 

Hills volcano (Montserrat) andesite dome lavas were interpreted to be caused by replacement of the ilmenite grains by 

the in-contact encroaching titanomagnetite grain (Devine and 

others, 2003). The concentration gradients were interpreted 

as arising from submicron-scale mixtures of encroaching titanomagnetite upon islands of relict ilmenite. It was concluded 

in that case that temperature estimates based on analysis 

of the titanomagnetite crystal in the gradient zone (likely a 

two-phase mixture) were spurious. In some cases, two-oxide 

geothermometry calculations for the gradient zones yielded 

temperature estimates greater than 1,000oC, far above the 

experimentally determined upper stability limit (855oC) of the 

amphibole crystals observed to be stable in the natural magma 

(Rutherford and Devine, 2003). The two-oxide geothermometry calculations produced spurious results for zoned oxide 

pairs. Similar high-Ti magnetite zones adjacent to ilmenite in 

the Mount St. Helens 2004–6 magma are interpreted to give 

spurious temperature estimates for the same reasons.

In addition to using pairs of titanomagnetite and ilmenite (table 3), estimates of preeruptive magma temperatures 

also were attempted by analyzing the cores of large individual 

titanomagnetite and ilmenite grains in the 2004–6 lava samples. 

The titanomagnetite analyses in SH305 fall into three groups, 

with high, intermediate, and low TiO2 contents (fig. 9). The 

range of TiO2 contents observed in the titanomagnetite crystals 

might, in part, reflect incorporation of xenolithic material from 

the conduit walls. Temperature estimates based on the combined 

analytical transects range from 816oC to 889oC for sample 

SH305 and from 823oC to 884oC for sample SH315. A significant part of the variance in the estimates is due to the wide range 

of compositions observed in the respective phases. Temperature estimates based on analyses of titanomagnetite grains in 

the intermediate-TiO2 group range from 823oC to 846oC when 

used in combination with the high-TiO2 ilmenite; the range of 

temperature estimates is 854–875oC when the low-TiO2 ilmenite 

is used in the calculations. Oxygen fugacities generally are less 

than one log unit above the NNO synthetic buffer (range +0.6 to 

+1.2 log units) at the calculated temperatures in all cases.
714 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample No.----

Date erupted--

Relation--------

SH304-1b2

10/18/04

Oxide pair in plag

SH305-1-8

11/20/04

Oxide pair in proximity

SH305-1-8

11/20/04

Oxide pair, small grains

Oxide ILM MT ILM MT ILM MT 

FeO 51.39 82.85 49.94 80.67 50.34 74.88

TiO2 42.98 8.54 44.02 11.13 44.36 12.73

Al2O3 0.21 2.33 0.17 1.79 0.21 1.74

Cr2O3 0.01 0.02 0.02 0.03 0 0.02

MgO 1.83 0.92 2.04 1.09 1.75 0.84

MnO 0.54 0.4 0.62 0.43 0.62 0.41

Total 96.96 95.06 96.82 95.15 97.29 90.63

Analyses 14 5 4 11 3 3

ToC 847 879 922

Log10 fO2

−11.82 −11.10 −11.59

∆NNO +1.16 +0.79 +0.50

Sample No.----

Date erupted---

Relation--------

SH305-1-8

11/20/04

Oxide pair in contact

SH315-1a

4/1/05

Oxide pair in proximity

SH315-1a

4/1/05

Oxide pair in proximity

Oxide ILM MT ILM MT ILM MT

FeO 51.90 78.65 49.24 82.63 51.27 82.29

TiO2 41.75 13.77 45.07 9.57 43.22 7.86

Al2O3 0.22 1.48 0.14 1.47 0.16 2.80

Cr2O3 0.02 0.03 0.02 0.06 0.02 0.02

MgO 1.90 1.20 1.74 0.91 1.80 1.23

MnO 0.56 0.52 0.63 0.44 0.44 0.37

Total 96.35 95.65 96.84 95.09 96.91 94.57

Analyses 5 7 14 6 5 7

ToC 963 828 832

Log10 fO2

−10.22 −12.57 −12.07

∆NNO +0.70 +0.79 +1.21

Table 3. Chemical composition of ilmenite and titanomagnetite occurring as pairs, with calculated 

temperature and fO2 of last equilibration geothermometry and oxygen geobarometry, Mount St. Helens, 

Washington, 2004–2006.

[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Compositions 

in weight percent; all Fe as FeO. Structural formula, temperature, and oxygen fugacity calculated after Devine and others (2003). 

Oxide phases abbreviated as follows: ILM, ilmenite; MT, magnetite.]

In spite of the difficulty of obtaining meaningful temperature estimates from in-contact Fe-Ti oxide pairs in the Mount 

St. Helens 2004–6 magma, the presence of the TiO2 concentration gradients is evidence of a recent heating event (Nakamura, 

1995; Venezky and Rutherford, 1999). However, the general 

lack of rim-to-core TiO2 diffusion gradients in titanomagnetite phenocrysts suggests that heating of the crystals must 

have taken place some weeks before eruption, judging from 

experimental work on similar samples (Rutherford and Devine, 

2003; Devine and others, 2003) and from diffusion data on Ti 

in magnetite (Venezky and Rutherford, 1999). It appears likely 

that the variability of TiO2 contents in titanomagnetite is due 

in part to commingling of parcels of magma that experienced 

different P–T histories prior to eruption. This interpretation is 

consistent with observations of cyclic variations of amphibole 

and orthopyroxene chemistry described above.

Two other important intensive parameters that can 

occasionally be determined for a preeruption magma using 

phenocryst and melt compositions are the pressure (depth) and 

the PH2

O (for example, Rutherford and others, 1985). The phenocryst assemblage in the 2004–6 magma suggests an intermediate depth (greater than 3.5–4 km) because the stability of 
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 715

amphibole in a melt requires that the water content of the melt 

be at least equal to H2O saturation at 100 MPa (Merzbacher 

and Eggler, 1984). The compositions of coexisting amphibole 

and plagioclase can also potentially yield both pressure and 

temperature information (Holland and Blundy, 1994), but the 

complex compositional zoning in both these phases makes it 

difficult to determine what amphibole coexists with a given 

plagioclase. In an attempt to circumvent this problem, the 

margins of amphibole crystals included in plagioclase were 

analyzed along with the immediately adjacent plagioclase for 

the 2004–6 magma samples. The results showed a complete 

range of amphibole compositions in contact with a given plagioclase and, thus, no decipherable indication of pressure or 

temperature. Similarly, the analysis of glassy melt inclusions 

in phenocrysts have yielded important data on the volatile content of the preeruption melt for other eruptions (for example, 

Rutherford and others, 1985; Wallace, 2005), but the relatively 

slow ascent and eruption of the 2004–6 dacite appears to 

have caused cracking and/or crystallization of almost all melt 

inclusions in the new Mount St. Helens magma. Analyses of 

2004–6 samples indicate dissolved water contents less than 

3 weight percent (J. Blundy, oral commun., 2006), and CO2

below FTIR detection limits (less than 40 ppm).

The observation that amphibole crystallized before plagioclase in the 2004–6 samples indicates that this dacite contained 

at least 4 weight percent dissolved H2O when the crystallinity was very low (less than 4–8 volume percent), as discussed 

above. This means that when the phenocryst content of the 

dacite reached the preeruption 40–50 volume percent as the 

result of decompression and/or cooling, the dissolved H2O in the 

remaining melt would have almost doubled. In fact, the residual 

melt probably would have reached H2O saturation where the 

storage-region pressure was less than 300 MPa. These arguments assume that all magma entering the magma storage zone 

was similar in composition, including dissolved volatiles. The 

fact that the erupted dacite has been uniform in composition and 

contains no mafic melt-bearing enclaves (Pallister and others, 

this volume, chap. 30) suggests that this assumption is justified.

Experimental Phase Equilibrium 

Constraints

In order to determine the P–T conditions where the compositions of the phenocrysts observed in the 2004 Mount St. 

Helens magma are stable, hydrothermal experiments were performed on a crushed powder of SH305. Experiments were done 

using water-saturated conditions, a decision justified by arguments presented in the previous section. Most experiments were 

at a log fO2

 close to NNO+1 because of the fO2 conditions determined from the Fe-Ti oxide phenocrysts, but higher and lower 

oxidation states also were investigated using solid-state buffers. 

Sample and buffer were contained in Ag or Ag-Pd tubes. Details 

of the experiments are given in table 4, and the results are plotted in figure 10. The main objective of these experiments was 

to determine the composition of the amphibole and plagioclase 

that is stable (grows in experimental charges previously equilibrated at somewhat higher or lower PH2

O and T) at each point in 

the range of P–T conditions where amphibole is stable in the 

magma. Our interest in the amphibole stability in this dacite 

magma stems from the observation that the main amphibole 

population appears to have crystallized at a range of conditions 

within the amphibole stability field prior to the final magma 

ascent to the surface. The final magma ascent is represented by 

the 0- to 5-μm-thick reaction rims that occur at the amphibolemelt contact (Rutherford and Devine, 2005).

As expected, the general hydrothermal phase equilibria 

of the 2004–6 dacite are similar to those determined for the 

slightly less evolved 1980 lava (63 versus 65 weight percent 

SiO2). The amphibole stability field at the fO2 conditions 

(NNO +1.0±0.5) of the many reconnaissance experiments is 

essentially unchanged by the composition difference, with the 

breakdown at 200 MPa occurring at 910±8oC (fig. 10). Above 

200 MPa, the reaction to form amphibole with decreasing temperature does not involve Ca-rich pyroxene, as it does at lower 

pressures. Orthopyroxene, plagioclase, magnetite, and ilmenite 

appear almost simultaneously on the liquidus (920±5oC at 200 

MPa) of the 2004–6 dacite for H2O-saturated conditions. Melt 

and phenocryst abundances were determined for representative 

experiments using the bulk-rock and melt compositions (table 

5), along with phenocryst compositions from each experiment, 

in a mass-balance calculation (Wright and Doherty, 1971). 

Just within the amphibole stability field at 200 MPa (900oC), 

the magma contains 8 weight percent crystals (4 percent 

plagioclase, 2 percent each of orthopyroxene and amphibole, 

and less than 1 percent of magnetite and ilmenite). At 850oC 

and 200 MPa, the crystals make up 32 weight percent of the 

magma (M-27; table 4), and at 850oC and 100 MPa, they make 

up 52 weight percent of the magma.

The composition of the plagioclase that is stable at any 

set of P–T conditions is contoured in figure 10 primarily on 

the basis of long-duration crystallization experiments, but 

reversals also were achieved. New-growth plagioclase is readily identifiable morphologically in higher temperature experiments (900±25oC), and it was analyzed. Very long duration 

(15 to 30 days) crystallization experiments used a starting 

material created at 235 MPa to determine the stable plagioclase-composition contours at lower pressures, all at 850oC. 

Plagioclase compositions used to create the contours in figure 

10 are recorded in table 4.

The compositions of the amphibole produced are plotted in figure 11 for the experiments where new growth could 

be identified optically and in BSE images (table 6). Plotted for 

comparison are the data from representative natural phenocrysts 

in sample SH305. Six experiments run at 850oC or 870oC and 

pressures of 100–230 MPa all have amphiboles with 9–10.5 

weight percent Al2O3, well outside the range observed in the 

normal cyclically zoned natural amphiboles (11.5–15 weight 

percent). The range of MgO (and FeO) in these low-temperature 

amphiboles is partly a function of the differences in fO2 in this 

group of experiments, as illustrated by the most oxidized experi-
716 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Run

No.

Starting

material

Pressure

(MPa)

Temperature 

(oC)

Buffer 

(Δ NNO)

Time

(days) Products

M-45a M-41 260 915 1 1.0 L

M-63a 305pdr 260 850 -1 1.0 Plg, A, Opx, Mt, I, L (25%)

M-63b M-16 260 850 -1 1.0 An53, A, Opx, Mt, I, L (25%)

M-73 305pdr 250 895 0 1.5 Plg, A, Opx, Mt, I, L

M-35 M-31 235 900 1 2 An68, A, M, L (3%)

M-20 305pdr 235 850 1 1.5 An50, A, Opx, Mt, I, L (32%)

M-49 M-17 225 860 1 5 An51, Opx, A, Mt, L

M-62 M-58+s 210 900 -1 2 An64, Opx,A, Mt, L

M-59a M-56a 205 900 -2 1 An62, A, Opx, Mt, L

M-59b M-19 205 900 -2 1 An62, A, Opx, Mt, L

M-60b M-16 205 900 -2 1.5 Plg, A, Opx, Mt, L

M-58 305pdr 200 930 1 1 L

M-16 305pdr 200 922 1 2 Opx, Mt, [Plg], L

M-24 305pdr 200 920 1 1.8 An71, Opx, Mt, M, L

M-19 M-17 200 915 1 0.5 Plg, Opx, M, L (4%)

M-36 M-17 200 915 1 2 An70, Opx, Mt, L (2%)

M-41a 305pdr 200 910 1 2 An67, Opx, A, Cpx, Mt, L

M-41b M-24 200 910 1 2 Plg, Opx, A, M, L

M-50b M-24 200 900 2 2 An62, Opx, A, Mt, L (8%)

M-51a M-41a 200 900 -1 1 Plg, Opx, A, Mt, I, L

M-51b 305pdr 200 900 -1 1 Plg, Opx, A, Mt, I, L

M-56a 305pdr 200 900 -2 0.7 Plg, A, Opx, Mt, L

M-25 M-24 200 885 1 2 An58, A, Opx, Mt, I, L

M-26 M-24 200 870 1 3 An52, A, Opx, Mt, I, L

M-66 305pdr 200 850 1 2 An42, A, Opx, Mt, I, L

M-27 M-24 200 850 1 6 An44, A, Opx, Mt, I, L

M-17 305pdr 200 840 1 4 Plg, A, Opx, Mt, I, L

M-21 M-20 180 850 1 28 An40, A, Opx, Mt, I, L

M-48 M-34 150 905 1 2 Plg, A, Opx, Mt, I, L

M-34 305pdr 150 880 1 2 An58, A, Opx, Mt, I, L

M-32 M-31 150 860 1 6 An42, A, Opx, Mt, I, L

M-28 M-24 150 860 1 6 Plg, A, Opx, Mt, I, L

M-22 M-20 140 850 2 28 An37, A, Opx, Mt, I, L

M-54 305pdr 125 940 1 2 An70, Opx, Mt, I, L

M-44 M-39 125 930 1 3 An68, Opx, Mt, I, L

M-39 305pdr 125 920 1 1 An58, A, Opx, Mt, I, L

M-38 305pdr 125 910 1 2 An54, Opx, Cpx, Mt, I, L

M-31 305pdr 125 875 1 4 An44, Opx, Cpx, [A], Mt, I, L

M-46 M-30 125 864 1 6 An40, A, Opx, Mt, I, L

M-53 M-38 125 860 1 6 An39, A, Opx, Mt, I, L

Table 4. Hydrothermal experiments on Mount St. Helens 2004 dacite.

[Total of 48 runs, arranged in order of decreasing pressure, from 260 to 70 MPa. Temperature ranges from 940oC to 840oC. Starting material 

was powder derived from sample SH305 (dome lava erupted November 2004) for 20 runs (shown as 305pdr); whereas starting material for other 

runs was product of an intervening run, shown listed. Oxygen fugacity buffered within 2 log units of nickel-nickel oxide (NNO). Crystallization 

products, listed in order of decreasing abundance (except glass), are plagioclase, Plg; amphibole; A; orthopyroxene, Opx; Cpx, clinopyroxene; 

magnetite, Mt; ilmenite, Il; and glass, L. Product listed in square brackets indicates phase is breaking down, is not in contact with melt, and 

is considered unstable at conditions of experiment; number in brackets refers to rim thickness on amphibole. Weight percent crystals shown 

parenthetically for some runs, determined by mass balance methods. Anorthite content of plagioclase, specified for 60 percent of runs, ranges 

from An71 to An34.]
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 717

50

100

150

200

250

300

840 860 880 900 920 940

TEMPERATURE, IN DEGREES CELSIUS

PH2O, IN MEGAPASCALS

Amph

Cpx

1980

Opx Plag + Mt

An42

An38

An35 Phase equilibria of

SH305 dacite

fO2

 = NNO+1

An50

An60

An70

+Ilm

. 8 32

51

M-54a M-44 120 912 1 2.5 An56, Opx, Cpx, Mt, I, L

M-54b M-38 120 912 1 2.5 Plg, Opx, Cpx, Mt, I, L

M-52 305pdr 120 845 1 21 An56, Opx, A, Mt, I, L (51 %)

M-47 305pdr 115 890 1 3 Plg, Opx, Cpx, Mt, I, L

M-30 305pdr 100 865 1 4 Plg, Opx, Cpx, Mt, I, [A, 8 μm], L

M-23 M-20 100 850 1 30 An34, Opx, A, Mt, I, L

M-12 305pdr 82 850 1 10 Plg, Opx, Cpx, Mt, I, [A, 10 μm]

M-11 M-20 70 850 1 6 Plg, Opx, Cpx, Mt, I, [A, 8 μm]

Run

No.

Starting

material

Pressure

(MPa)

Temperature 

(oC)

Buffer 

(Δ NNO)

Time

(days) Products

Figure 10. Pressure-temperature phase diagram determined 

for the Mount St. Helens 2004–6 dacite at water-saturated 

conditions and oxygen fugacity equal to NNO+1 log units. 

Symbols for individual experiments (table 4) point in the 

direction of approach to the final P and T plotted; solid 

symbols indicate amphibole is stable. Solid heavy lines 

mark upper stability limit for a given phenocryst phase: Cpx, 

Ca-pyroxene; Amph, amphibole; Plag, plagioclase; Opx, 

orthopyroxene; Mt, magnetite; Ilm, ilmenite. Medium-weight 

lines are contours labeled with plagioclase anorthite content 

(for example, An70), on basis of data in table 4. Boxed numbers 

(8, 32, 51) indicate weight percent crystals present at P–T

conditions occupied by box, as determined by a mass balance 

calculation (see text). Solid bar labeled 1980 shows estimated 

conditions in preeruption 1980 magma (Rutherford and others, 

1985; Venezky and Rutherford, 1999).

Table 4. Hydrothermal experiments on Mount St. Helens 2004 dacite.—Continued

[Total of 48 runs, arranged in order of decreasing pressure, from 260 to 70 MPa. Temperature ranges from 940oC to 840oC. Starting material 

was powder derived from sample SH305 (dome lava erupted November 2004) for 20 runs (shown as 305pdr); whereas starting material for other 

runs was product of an intervening run, shown listed. Oxygen fugacity buffered within 2 log units of nickel-nickel oxide (NNO). Crystallization 

products, listed in order of decreasing abundance (except glass), are plagioclase, Plg; amphibole; A; orthopyroxene, Opx; Cpx, clinopyroxene; 

magnetite, Mt; ilmenite, Il; and glass, L. Product listed in square brackets indicates phase is breaking down, is not in contact with melt, and 

is considered unstable at conditions of experiment; number in brackets refers to rim thickness on amphibole. Weight percent crystals shown 

parenthetically for some runs, determined by mass balance methods. Anorthite content of plagioclase, specified for 60 percent of runs, ranges 

from An71 to An34.]
718 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample ---

T (oC) ------

P (MPa)---

fO2 (ΔNNO)

SH-305

1300

0.1

+1

M-62

900

210

−1

M-59a

900

205

−2

M-51a

900

200

−1

M-27

850

200

+1

M-17

850

200

+1

M-23

850

100

+1

M-34

890

150

+1

Major-element analyses, weight percent

SiO2 64.80 (24) 62.48 (28) 62.48 (26) 61.59 (39) 65.30 (98) 67.69 (52) 72.83 (24) 67.95 (32)

TiO2 0.64 (5) 0.60 (5) 0.54 (7) 0.52 (6) 0.32 (9) 0.25 (3) 0.25 (4) 0.34 (4)

Al2O3 17.23 (14) 15.85 (13) 15.94 (19) 16.06 (4) 14.46 (28) 13.91 (18) 12.38 (12) 13.68 (10)

FeO 4.35 (10) 3.24 (12) 3.22 (8) 3.15 (16) 2.10 (52) 2.11 (17) 1.42 (11) 2.40 (9)

MgO 1.91 (4) 1.26 (5) 1.24 (7) 1.45 (11) 0.68 (51) 0.46 (3) 0.21 (2) 0.53 (11)

CaO 4.68 (8) 3.93 (7) 3.93 (7) 4.28 (19) 1.96 (44) 2.13 (11) 1.20 (7) 2.27 (6)

Na2O 4.58 (9) 4.84 (10) 4.84 (10) 4.64 (19) 5.60 (25) 4.77 (19) 4.60 (14) 4.66 (17)

K2O 1.45 (7) 1.53 (6) 1.53 (6) 1.33 (11) 2.40 (14) 1.83 (9) 2.43 (11) 1.97 (4)

MnO 0.10 (3) 0.02 (3) 0.02 (3) 0.05 (2) 0.06 (5) 0.04 (3) 0.02 (4) 0.04 (3)

Total 99.74 93.75 93.68 93.07 92.89 93.19 95.34 93.84

Analyses 9 6 6 6 6 7 6 6

VBD 0 6.25 6.32 6.93 7.11 6.81 4.66 6.16

Table 5. Bulk and experimental glass compositions, Mount St. Helens, Washington, 2004–2006.

[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Each analysis is an average; total number of 

analyses for each sample listed in row “Analyses.” Compositions in weight percent; all Fe as FeO. Parenthetical numbers, standard deviation on average for the 

last places given. VBD, dissolved volatiles (H2O) by difference. Column 1 (all glass), SH305 bulk composition.]

ment, M-22 (fig. 11A). The externally imposed fO2 in the 900oC 

experiments ranges from 2 log units below to 2 log units above 

NNO, bracketing the fO2 indicated by the Fe-Ti oxides (NNO +1) 

in the natural sample.

The variation in Al2O3 content (including Altotal and VIAl) 

of amphibole that crystallizes from the 2004–6 dacite (table 6) 

should be a function of temperature and pressure as discussed 

earlier in this paper, and the experimental data (fig. 11) show 

this to be the case. It is notable, however, that only experiments done at 200 MPa and higher pressures have compositions that plot in the zone defined by the cyclically zoned 

natural phenocrysts. Additionally, none of the experiments 

in the pressure range 200–250 MPa produced amphiboles 

comparable to the highest Al- and Fe-rich zones in the natural 

amphibole. Experiments at 200 MPa, 900oC, and a log fO2 

in the range NNO+2 to NNO-2 produce the low and intermediate Al2O3 compositions in the cyclically zoned natural 

phenocrysts. Experiments at 900oC and pressures of 235 MPa 

(NNO+1) and 250 MPa using the NNO buffer did not produce 

an amphibole with higher Al. Both the Al2O3 and VIAl composition data indicate that a pressure somewhat greater than 

250 MPa is required to stabilize the highest Al- and Fe-rich 

amphibole end members of the natural phenocrysts.

Depth of the Preeruptive 2004–2006 

Magma System

The question of preeruptive magma-storage depth is of 

considerable interest and importance for understanding the 

ongoing Mount St. Helens eruption and for predicting how 

the eruption is likely to progress. Is the magma erupting from 

the same storage zone determined for the 1980–86 eruptions 

(Scandone and Malone, 1985; Rutherford and others, 1985), or 

has the system changed? The seismic record over the decade 

before the new eruption suggests that magma may have moved 

to a relatively shallow depth (3–4 km) below the crater floor, 

particularly at one or two times during this period (Moran, 

1994; Moran and others, this volume, chap. 2). Seismicity accompanying the 2004 eruption appears to have been 

largely in the region from 2 km to the surface, consistent with 

magma erupting from storage below this depth. Theoretically, 

the abundance and composition of phenocrysts, particularly 

the rim compositions, are dependent on the final P and T of 

phenocryst-melt equilibration in the erupting magma. We have 

attempted to use this knowledge and the phenocryst data from 

the natural sample and experiments to determine the depth of 

the magma storage upper boundary.

As described in the previous section, most of the amphibole-phenocryst growth in the 2004–6 dacite was at pressures 

of 200 MPa or higher and at temperatures greater than 850oC. 

The phenocryst abundance in the dacite at these high pressures is less than 32 volume percent at temperatures greater 

than 850oC (fig. 10). However, given that much of the magma 

erupted at 850oC with ~45 volume percent phenocrysts (Pallister and others, this volume, chap. 30), plagioclase phenocryst 

growth must have continued to a much lower pressure (fig. 

10). Analytical data from experiments designed to achieve a 

good approach to plagioclase phenocryst rim-melt equilibrium 

as a function of pressure at 850oC are compared with rimcomposition measurements in the natural samples (figure 12). 
Sample No.---P (MPa)-------T (oC)----------fO2 (∆NNO)--

M-62

210900

-1

M-59b

205900

-2

M-59a

205900

-2

M-51a

200900

-1

M-49

225860

1

Major-element analyses, weight percent

SiO2 42.63 43.68 43.58 42.07 44.44

TiO2 2.48 2.28 2.62 2.72 1.89Al2O3 13.38 12.54 11.31 13.00 10.31FeO 13.13 10.72 11.48 14.07 14.95MgO 13.30 15.21 14.90 13.14 13.15CaO 10.95 11.07 11.06 11.58 10.56Na2O 2.25 2.37 2.26 2.39 1.89

K2O 0.27 0.33 0.26 0.32 0.26

MnO 0.19 0.11 0.21 0.10 0.26Total 98.58 98.31 97.68 99.39 97.71

Cation abundance in structure formula

Si 6.119 6.224 6.275 6.041 6.452

AlIV 1.881 1.776 1.725 1.959 1.548

Ti 0.268 0.244 0.284 0.294 0.206AlVI 0.383 0.331 0.195 0.241 0.216Fe3+ 0.728 0.654 0.701 0.745 0.823

Fe2+(Y) 0.776 0.536 0.622 0.908 0.909

Mg 2.846 3.231 3.198 2.812 2.846Ca 1.684 1.690 1.706 1.782 1.643Fe2+(X) 0.072 0.088 0.059 0.037 0.083Na 0.627 0.609 0.631 0.665 0.532Mg# 78.57 85.78 83.71 75.60 75.80Fe3+/ΣFe 0.462 0.512 0.507 0.441 0.454

Table 6. Representative experimental amphibole compositions and structure formulas, Mount St. Helens, Washington, 2004–2006.[Chemical analyses determined by electron microprobe at Brown University, Providence, R.I.; J.D. Devine, analyst. Compositions in weight percent; all Fe as FeO. Structure folated after Holland and Blundy (1994). Magnesium number, Mg# = (Mg×100)/(Mg+(Fe2+ in Y site)).]

Sample No.---P (MPa)-------T (oC)----------fO2 (∆NNO)--

M-32

150200

1

M-28

150860

1

M-20

234850

1

M-27

200850

1

Major-element analyses, weight percent

SiO2 44.72 45.19 45.59 42.92 

TiO2 2.46 1.73 1.68 2.14 Al2O3 9.29 10.26 9.48 11.49 FeO 14.58 13.12 14.21 15.80 

MgO 13.07 14.04 13.71 12.19 

CaO 11.14 11.28 11.32 10.97 

Na2O 1.97 1.98 1.79 2.23 

K2O 0.27 0.28 0.26 0.32 

MnO 0.14 0.18 0.28 0.21 

Total 97.64 98.06 98.32 98.27 

Cation abundance in structure formula

Si 6.540 6.515 6.575 6.264 

AlIV 1.460 1.485 1.425 1.736 

Ti 0.271 0.188 0.182 0.235 

AlVI 0.141 0.259 0.187 0.240 

Fe3+ 0.558 0.625 0.698 0.745 

Fe2+(Y) 1.182 0.911 0.986 1.129 

Mg 2.849 3.017 2.947 2.652 

Ca 1.745 1.77 1.749 1.715 

Fe2+(X) 0.043 0.046 0.030 0.055 

Na 0.558 0.554 0.500 0.631 

Mg# 70.68 76.81 74.93 70.14 

Fe3+/ΣFe0 0.313 0.395 0.407 0.386 
720 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

8

9

10

11

12

13

14

15

16

9 10 11 12 13 14 15 16 17

EXPERIMENTAL VERSUS SH305 AMPHIBOLES

MgO, IN WEIGHT PERCENT

A

0.05

0.15

0.25

0.35

0.45

1.5 2 2.5

Altotal = IVAl + VIAl

20

32

28

22

49

27

73

59a

62

51

59b

B

VIAl, IN Y SITE Al2O3, IN WEIGHT PERCENT

Figure 11. Compositions of amphiboles produced in 

hydrothermal experiments (table 4) compared with those of 

natural cyclically zoned phenocrysts in SH305, a representative 

phenocryst in 2004–6 magma. Black solid dots, natural 

amphiboles; open symbols, experimental results at temperatures 

from 840oC to 870oC and pressures less than 200 MPa. Solid 

red symbols represent experiments at 900oC and pressures of 

200–260 MPa. The fO2

 of the 900oC experiments ranged from 

NNO-2 to NNO+2; others were all at NNO+1 log units. Numbered 

symbols indicate amphibole compositions (table 6) produced in 

experiments; numbers correspond to sample numbers in table 4. 

A, Al2O3 versus MgO. B, VIAl versus Altotal.

The results of the plagioclase-growth experiments indicate that 

if the 2004–6 magma was stored for more than 30 days at 180 

MPa before experiencing a fairly rapid ascent to the surface, 

the rim compositions would be approximately An40. The similarity of the 14- and 28-day experimental results suggest that 

much longer storage at this depth would not change the most 

albite-rich composition observed, as long as the temperature 

and PH2

O remained unchanged. At 140 and 100 MPa, the plagioclase in equilibrium with the residual melt at 850oC is An35

and An32, respectively. Comparing these experimental plagioclase compositions to the most Na-rich rims on natural phenocrysts in SH304 and SH305 suggests that the phenocrysts in 

the 2004–6 magma were crystallizing at a pressure as low as 

120–140 MPa before the final ascent. The natural phenocrysts 

with more Ca-rich rims (fig. 12) are interpreted to have experienced little or no growth in the decompression accompanying ascent to 120 MPa. This lack of new plagioclase growth 

is expected in an ascending magma when the phenocryst 

content is high, because the melt immediately adjacent to 

some phenocrysts is limited in volume by other phenocrysts 

that are close or touching. Additionally, because the ascent 

from the storage zone to the surface did not produce significant reaction rims on amphiboles, we conclude that measurable additions to the plagioclase phenocrysts are unlikely to 

be added during the final (100 to 0.1 MPa) ascent. We make 

this conclusion because H2O loss from the melt in ascending 

magma is the main cause of both amphibole breakdown and 

plagioclase crystallization. The amphibole rims are less than 5 

μm thick in these samples, indicating a relatively rapid magma 

ascent (Rutherford and Hill, 1993; Rutherford and Devine, 

2003). Although nucleation and growth of plagioclase microlites occurs readily in such decompressions (Geschwind and 

Rutherford, 1995), the experimental data indicate that additions to associated phenocrysts would be minor and too thin to 

analyze by electron microprobe.

To summarize, the plagioclase rim compositional data 

indicate that the 2004–6 magma could have been stored at a 

pressure as low as 120 MPa, or a depth of 4–5 km below the 

crater floor for a relatively prolonged period before eruption. When these data are considered together with the depth 

range indicated by seismic data during the eruption (Moran 

and others, this volume, chap. 2), the 4–5-km depth appears 

to represent the top of the magma storage zone. However, it is 

also obvious that the preeruptive magma storage extended to 

much greater depths, and the Na-rich plagioclase phenocryst 

rims may have developed during the slow, final ascent process 

that began somewhere below 4 km.

The amphibole phenocryst compositions and their very 

thin decompression-induced reaction rims do not significantly 

constrain estimates of the preeruptive magma storage-zone 

depth, except to indicate that it was greater than 100 MPa (4–5 

km) for magma at 850oC. If the storage zone pressure had been 

less than 100 MPa, there would have been thick, coarse-grained 

reaction rims on all amphiboles. For example, 10-μm-thick 

rims were produced in experiments at 85 and 70 MPa on the 

2004–6 dacite in as little as 7 days (experiments M-11 and 

M-12; table 4). However, as discussed in the following section, the crystallization of the cyclically zoned amphiboles 

must have occurred at pressures more than twice the 100 MPa 

estimated for the final pressure of phenocryst-melt equilibration. No attempt was made to use the Al-in-hornblende 

geobarometer to estimate the pressure because the phenocryst 

phase assemblage in the 2004–6 dacite is not silica saturated as 

required by the geobarometer calibration (Johnson and Rutherford, 1989). Using the geobarometer (Thornber and others, this 

volume, chap. 32) should give an upper-pressure limit, particularly for the early crystallization of the Al-rich amphiboles.
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 721

30

35

40

45

50

55

60

Mount St. Helens

phenocryst rims

PRESSURE, IN MEGAPASCALS

Naturally

occurring rims

Experimentally produced rims, 850oC

0 50 100 150 200

PLAGIOCLASE An CONTENT,

IN MOLE PERCENT 

Figure 12. Comparison of plagioclase phenocryst-rim 

compositions in samples from 2004–2006 Mount St. Helens 

eruption (SH304 and SH305) with those produced in long-duration 

crystallization experiments at 850oC and pressures of 100 to 

180 MPa (table 4). Squares, 15-day runs; circles, 30-day runs. 

Similarity between different run durations indicates experiments 

were sufficiently long to approach chemical equilibrium. Arrow 

indicates pressure (about 120 MPa) where these experiments 

produce new plagioclase that is compositionally equal to most 

Na-rich rims on the natural phenocrysts. More Ca-rich rims in 

natural plagioclase are interpreted as rims where there was 

limited late-stage growth.

Origin of the Cyclically Zoned 

Phenocrysts: Preeruptive Magma 

History

A second objective of this project was to determine the 

origin and significance of the cyclic compositional zonation 

observed in 2004–6 magma phenocrysts. Is it possible to 

crystallize the cyclically zoned plagioclase, amphibole, and 

orthopyroxene phenocrysts by simply changing conditions in 

the dacite magma, or are injections of mafic magma required? 

For example, the presence of cyclically zoned phenocrysts in 

the silicic andesite magma erupted recently (1995–2002) at 

Soufrière Hills, Montserrat, was clearly the result of recrystallization following injections of a more mafic basaltic 

andesite (Rutherford and Devine, 2003). Vesicular blobs of 

the mafic magma were carried up in the erupting magma, 

and small, blade-shaped crystals of high-An plagioclase and 

pargasitic amphibole from the mafic magma are present in the 

andesite groundmass. However, there is no evidence of recent 

mafic magma injections into the 2004–6 dacite (Pallister and 

others, this volume, chap. 30). Thus, we have concentrated on 

determining how the 2004–6 phenocryst compositions could 

have been formed from dacite magma by trying to determine 

the range of pressure-temperature conditions where the various phenocryst compositions would crystallize.

Figure 13 is a model of the Mount St. Helens 2004–6 

magma storage zone that uses the experimental data described 

above to explain the cyclic zoning of the natural amphibole, 

plagioclase, and orthopyroxene phenocrysts. Because there is 

no evidence, such as mafic enclaves, to indicate other magma 

compositions were mixed or mingled with the 2004–6 dacite, 

our model assumes a dacite magma input. The timing of this 

input is not known; it could have been associated with the 

beginning stages of the 2004–6 eruption, or it could have 

been associated with past (1986–98) peaks in deep seismic 

activity (Moran and others, this volume, chap. 2). Given that 

the preeruptive magmatic temperatures decreased from 1980 

to 1986 (Rutherford and Hill, 1993), it seems unlikely that 

the hot 2004–6 magma was emplaced during the 1980–86 

eruptions, although this possibility cannot be ruled out. The 

homogeneity of the 2004–6 dacite suggests that the magma 

composition entering the base of the storage zone is similar 

to the magma that has been erupting. Uranium-series dating 

of mineral separates from the 2004–6 magma indicate there 

is an old component present in the erupting magma (Cooper 

and others, this volume, chap. 36), but this may be explained 

by the xenolithic material, and it does not appear to affect our 

conclusions about magma entering the 2004–6 system.

If the incoming dacite magma had a high dissolved water 

content, as the phenocryst phase assemblage suggests, it 

would crystallize Fe- and Al-rich amphibole before or during 

its invasion of the storage zone at about 900oC (fig. 13). This 

temperature for the incoming magma is suggested by the highTiO2 magnetite phenocrysts in some of the samples (fig. 9), by 

the lack of plagioclase inclusions in amphibole phenocrysts 

that indicate that amphibole crystallized before plagioclase, 

and by the anorthite content (An68) in the cores of cyclically 

zoned plagioclase phenocrysts. Small amounts of orthopyroxene, magnetite, and ilmenite crystallized along with the early 

amphibole as indicated by inclusions in amphibole and the 

experimental phase equilibria. The pressure on the magma as 

it enters the storage zone is interpreted to be ~300 MPa (12 km 

depth) on the basis of the experiments (fig. 11) that indicate 

such pressures are required to stabilize the high-Al zones in the 

amphibole phenocrysts. Importantly, An68 plagioclase is not 

stable above 200 MPa at 880–900oC, according to the watersaturated experiments (fig. 10), and would be lacking in the 

phenocryst-poor magma just after it entered the storage zone.

As the water-rich magma moved higher in the storage 

zone, the high-An (An68) zones in plagioclase would begin 

to crystallize at 200 MPa and temperatures near 900oC. At 

this point (about 7–8 km depth), much of the magma must 

stagnate and cool in the storage zone in order to form the more 

Ab-rich zones in the plagioclase and the more Mg-rich zones 

in the amphiboles. The plagioclase phenocrysts require that 

the magma cools to 850oC in order to stabilize and grow the 

albite-rich layer (An45–50) in the first growth cycle. We interpret 

this cooling as occurring in a lateral flow (fig. 13). If cooling 

did not occur, the original phenocrysts would tend to dissolve 

during convection to depth, and textures within phenocrysts 

indicate dissolution was not extensive. The second cycle of 
722 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

8

4

12

DEPTH, IN KILOMETERS (DATUM IS CRATER FLOOR)

Mount St. Helens

Magma System

2004–6

Ascent of magma 

into conduit; some 

low-pressure 

recrystallization

Minimum P for

stabilizing 

high-alumina 

amphibole

Dacitic magma enters reservoir; 

temperature 900oC and oxygen 

fugacity at NNO buffer.

Amphibole stable, but not 

plagioclase

Magma enters conduit; 

amphibole rim development 

begins, and measurable 

growth of plagioclase 

phenocrysts ends

Magma crystallizes 

An-rich plag, cools, 

and is recycled to 

depth

Cycles of 

phenocryst 

growth

Figure 13. Simplified cross section showing model of the Mount 

St. Helens magma storage zone and transport system based on 

petrologic data for the 2004–6 samples discussed in this paper 

(see also Pallister and others, this volume, chap. 30) and the 

seismic depths associated with magma movement before and 

during the 2004 eruption (Moran and others, this volume, chap. 2). 

Depth scale is relative to the crater floor.

amphibole growth begins with an Fe- and Al-rich material 

that is stable only at pressures substantially greater than 200 

MPa (fig. 11); the amphibole data require the convection 

system shown in figure 13. Sharp chemical transitions—those 

between Na-rich and Na-poor zones in plagioclase (200 MPa 

and 850oC) and those between Mg-rich and Fe-rich zones in 

amphibole—are consistent with the convection process, in that 

crystallization would not occur in the sinking magma.

The forces to drive the convection certainly are present. The melt-rich character, relatively high temperature, and 

high dissolved H2O content would tend to make the incoming 

magma buoyant in the storage zone occupied by older, cooler, 

and more crystal-rich magma. During ascent from the 12-km 

depth, the buoyancy of this magma would decrease as a result 

of decompression-induced crystallization. Cooling, crystallization, and loss of released gas at the top of the convective cell would further reduce the buoyancy of the magma, 

apparently to the point where the combination of density and 

viscosity cause it to sink in the chamber. Using the data of 

Geschwind and Rutherford (1995) for the densities of the 

phases, the density of the magma would increase from 2,310 

kg/m3

, as magma with 5 volume percent crystals (amphibole 

and orthopyroxene) enters the storage zone, to 2,450 kg/m3 at 

850oC where magma contains 30 volume percent phenocrysts 

(in the proportions plag:amph:opx:mt = 23:3:3:1 on a volume 

percent basis). This represents a 6-percent density increase 

of the crystal-rich magma relative to the incoming melt-rich 

magma, assuming gas released in the upwelling process is 

able to escape during the crystallization. Given that viscosity 

of the H2O-rich melt will be relatively low, the viscosity of 

the phenocryst-poor magmas should facilitate the convection. 

Sharp increases in viscosity when the phenocryst content goes 

above 30 volume percent (Marsh, 1981) may indicate why the 

cycle ends and why there are only two zoning cycles, on average, in amphibole.

Following an average of two cycles of phenocryst 

growth and convective overturn, some fraction of the magma 

would separate and rise toward the conduit. Although there 

were additions to the plagioclase phenocrysts in this stage 

of magma ascent (fig. 12), there was almost no additional 

amphibole crystallization, because the interstitial H2Osaturated melt at 200 MPa and 850oC has essentially no 

amphibole component, and there is no Ca-rich pyroxene in 

the H2O-rich dacite. However, amphibole recrystallization to 

form Fe-rich, Al-poor amphibole rims, as observed in SH323 

(fig. 7) may occur in magma that spent a somewhat longerthan-average time in the less-than-200-MPa pressure range 

during this ascent.

A variation of the above model involving an incoming magma that has a somewhat lower initial dissolved H2O 

content cannot be ruled out. Rather than H2O-saturated, 

the incoming magma could have a PH2O as low as 0.7 Ptotal, 

in which case the amphibole stability would not have been 

affected, but plagioclase phenocrysts would have been stabilized to a higher temperature at a given pressure. This is 

equivalent to proposing that the magma was saturated with 

a CO2-bearing fluid when it entered the storage zone. As 

discussed above, however, the initial melt-rich magma must 

have contained at least 4 weight percent dissolved H2O in 

order to stabilize amphibole, and this limits the CO2 content 

of the coexisting fluid (Newman and Lowenstern, 2002) to 

~600 ppm. The maximum temperature of the magma coming 

into the storage zone at 260–300 MPa is still required to be 

less than 910oC in order to stabilize the high-Al amphibole. 

However, in this case, the An68 plagioclase cores could have 

formed at pressures greater than 200 MPa, possibly as early 

as the amphibole. The absence of plagioclase inclusions in 

the amphibole cores suggests this variant of the storage zone 

model is unlikely.

Comparison of 2004–2006 and 1980–

1986 Preeruption Magmas

Some additional data collected for the 1986 lava dome 

(sample SH205) and for the 1980 pumice (SH084) during the 

course of this study indicate that a set of convective processes similar to those described above was operating in the 

magma storage zone during the 1980–86 eruptions. Magma 
31. Magmatic Conditions and Processes in the Storage Zone of the 2004–2006 Mount St. Helens Dacite 723

erupted in 1986 last equilibrated at 860±10oC (Rutherford 

and Hill, 1993) and contained amphiboles with cyclic zoning 

generally indistinguishable from that in the 2004–6 magma. 

A plot of Al2O3 versus MgO for amphiboles (fig. 14) shows 

essentially total overlap between typical amphiboles from 

SH205 (1986) and SH323 (2005; table 2 and fig. 4). A cyclical set of phenocryst growth conditions similar to that outlined for the 2004–6 magmas (fig. 13) appears to be required 

for the 1986 dome-forming lava. The lack of decompressioninduced rims on amphibole in the 1986 sample (fig. 3D) is 

one significant example of how the 1986 magma differs from 

that erupted in 2004–6. This observation seems to require an 

even more rapid ascent for the 1986 magma over the final 

4–5 km. Outer rims on the 1986 amphibole that are higher 

in FeO, relative to those in the 2004–6 magma, might also 

be explained by a rapid ascent of magma from depth. This 

mineralogical evidence of amphibole crystallization over a 

pressure range suggests a process similar to one proposed 

recently (Blundy and Cashman, 2005; Cashman and McConnell, 2005) as operating in the Mount St. Helens magma system, on the basis of melt-inclusion data. The magma ascent 

history, however, appears to have been more complicated 

than can be determined from the melt inclusions.

8

9

10

11

12

13

14

15

11 12 13 14 15 16

1980,1986, and

2005 amphiboles

MgO, IN WEIGHT PERCENT

SH205 (10/1986)

SH323 (9/2005)

SH084 (5/1980)

 

Al2O3, IN WEIGHT PERCENT

(all open symbols, x, and +)

Figure 14. Compositions of eight amphibole phenocrysts (rim to rim transects) 

in the 1980 Mount St. Helens pumice (SH084) compared with similar amphibole 

transects for a typical sample of the 2004–6 lava dome (SH323) and 1986 dome 

(SH205). Lines are fit to the eight SH084 amphibole profiles. Zoning in the 1980 

pumice phenocrysts is similar to that in the 2004–6 phenocrysts except that 

it occurs at lower Al2O3 in most crystals. Note that rims on SH323 amphiboles 

spread over the low-Al2O3 compositions found in the May 1980 samples, whereas 

the zoning profile in SH205 (1986 lava) is indistinguishable from 2004–6 samples 

such as SH304 and SH305 (fig. 4).

Compositional zoning in representative amphibole phenocrysts from the May 18, 1980, white pumice is shown in 

figure 14. The Al2O3 versus MgO zoning within individual 1980 

phenocrysts defines line segments that are parallel to the 2004–6 

trend, but they range down to significantly lower Al2O3. The 

lower-Al2O3 phenocrysts are similar in composition to the Alpoor rims analyzed on amphiboles in SH323 (fig. 7) and other 

more recently erupted 2004–6 samples. In a previous section we 

suggested that the composition of the anomalous Al-poor rims 

on SH323 amphiboles are best explained by recrystallization at 

pressures of 100–200 MPa, possibly even just outside (below) 

the stability limit of OH-amphibole. This would help explain the 

fluorine enrichment of these rim compositions. These observations suggest that the convective process described in the model 

for the 2004–6 magma storage zone also was working in 1986 

and 1980. However, at least some of the 1980 magma seems to 

have experienced a significant part of the amphibole crystallization at a pressure in the 150–200 MPa range at the 880oC temperature (Rutherford and Devine, 1988; Venezky and Rutherford, 1999) of this system. The 1980 phenocrysts generally are 

higher in F (~1,200 ppm) than are the cores of 2004–6 samples 

(900 ppm) but are not as high as the 2004–6 rims (SH323), 

consistent with this lower-pressure origin.
724 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Conclusions

In summary, and in answer to the questions posed in the 

introduction to this paper, the following conclusions can be 

made: (1) The phenocryst assemblage that occurs in the 2004–6 

Mount St. Helens dacite lava-dome samples can be attributed to 

crystallization from dacite magma; no mafic magma injection 

is required. (2) The compositions of the Fe-Ti oxides, magnetite and ilmenite, although somewhat variable, indicate that the 

temperature of equilibration was in the range of 820–890oC and 

the fO2 was NNO + 1 log unit. The existence of a range of TiO2

in different, relatively homogenous titanomagnetite phenocrysts 

also suggests there was mingling of different batches of dacite 

with different thermal histories to create each batch of erupting magma. (3) Cyclic zoning observed in amphiboles in the 

2004–6 magma consists of Al- and Fe-rich zones alternating 

with Mg- and Si-rich zones and can be explained as approximately equal amounts of two substitutions: a pressure-sensitive 

Al-Tschermak substitution and a temperature-controlled 

edenite substitution. Significantly, different amounts of the AlTschermak end member, in different zones of each amphibole 

phenocryst, require crystallization over a range of total pressure. 

Hydrothermal experiments on the dacite composition show 

that the Fe- and Al-rich end member of the amphibole zoning 

must have crystallized at pressures of ~300 MPa, whereas more 

Mg- and Si-rich amphibole zones crystallized in the 200-MPa 

range. (4) Cyclical zoning in natural plagioclase phenocrysts, 

ranging from An68 to An35, also requires the magma to cycle 

through a range of crystallization conditions. The hydrothermal 

experiments indicate that An68 plagioclase would crystallize 

at ~200 MPa and 900oC, well after the initial crystallization of 

amphibole in an ascending dacite magma. Cooling to ~850oC is 

required to develop the first layer of Ab-rich plagioclase. (5) In 

order to explain the cyclic zoning in the phenocrysts, each batch 

of magma must have experienced convective overturn, going 

back to near the original high pressure at the lower temperature 

(850oC), where it interacted with new incoming dacite magma, 

as illustrated in figure 13. (6) Rims on plagioclase phenocrysts, as 

sodic as An35, indicate that plagioclase phenocryst growth continued in ascending magma until the pressure was ~130 MPa. This 

pressure is interpreted to represent the top of the magma storage 

zone at 4–5 km depth. (7) Compositional zoning in the 1980 and 

1986 amphibole phenocrysts indicates that convective conditions 

also existed during phenocryst growth in that magma system.

Acknowledgments

We acknowledge the financial support of the National Science Foundation and field support from the Cascades Volcano 

Observatory staff for providing samples. We thank the members 

of the Mount St. Helens petrology working group for assistance 

and ideas and for providing access to manuscripts during their 

preparation. Finally, we acknowledge productive discussions 

with John Pallister and Carl Thornber and insightful reviews by 

Jon Blundy and Wes Hildreth.

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eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

Nakamura, M., 1995, Continuous mixing of crystal mush and 

replenished magma in the ongoing Unzen eruption: Geology, v. 23, p. 807–810.

Newman, S., and Lowenstern, J.B., 2002, Volatilecalc; a silicate melt-H2O-CO2 solution model written in Visual Basic 

for exel®: Computers and Geosciences, v. 28, no. 5, p. 

597–604, doi:10.1016/S0098-3004(01)00081-4.

Pallister, J.S., and Thornber, C.R., 2005, Is the 2004–05 eruption of Mount St. Helens tapping new dacite from the deep 

crust? [abs.]: Eos (American Geophysical Union Transactions), v. 86, no. 52, Fall Meeting Supplement, Abstract 

V52B–05.

Pallister, J.S., Thornber, C.R., Cashman, K.V., Clynne, M.A., 

Lowers, H.A., Mandeville, C.W., Brownfield, I.K., and 

Meeker, G.P., 2008, Petrology of the 2004–2006 Mount St. 

Helens lava dome—implications for magmatic plumbing 

and eruption triggering, chap. 30 of Sherrod, D.R., Scott, 

W.E., and Stauffer, P.H., eds., A volcano rekindled; the 

renewed eruption of Mount St. Helens, 2004–2006: U.S. 

Geological Survey Professional Paper 1750 (this volume).

Rutherford, M.J., and Devine, J.D., 1988, The May 18, 1980, 

eruption of Mount St. Helens, III; stability and chemistry of 

amphibole in the magma chamber: Journal of Geophysical 

Research, v. 93, no. B10, p. 11949–11959.

Rutherford, M.J., and Devine, J.D., 2003, Magmatic conditions and magma ascent as indicated by hornblende phase 

equilibria and reactions in the 1995–2002 Soufrière Hills 

magma: Journal of Petrology, v. 44, p. 1433–1454.

Rutherford, M.J., and Devine, J.D., 2005, Crystallization of 

and conditions in the MSH 2004–05 dacite magma as indicated by phenocryst compositions and experiments [abs]: 

Eos (American Geophysical Union Transactions), v. 86, no. 

52, Fall Meeting Supplement, Abstract V52B–05.

Rutherford, M.J., and Hill, P.M., 1993, Magma ascent rates 

from amphibole breakdown; an experimental study applied 

to the 1980–1986 Mount St. Helens eruptions: Journal of 

Geophysical Research, v. 98, no. B11, p. 19667–19685.

Rutherford, M.J., Sigurdsson, H., Carey, S., and Davis, A., 

1985, The May 18, 1980, eruption of Mount St. Helens, 1; 

melt composition and experimental phase equilibria: Journal 

of Geophysical Research, v. 90, no. B4, p. 2929–2947.

Scandone, R., and Malone, S.D., 1985, Magma supply, 

magma discharge and readjustment of the feeding system 

of Mount St. Helens during 1980: Journal of Volcanology 

and Geothermal Research, v. 23, nos. 3–4, p. 239–262, 

doi:10.1016/0377-0273(85)90036-8.

Thornber, C.R., Pallister, J.S., Lowers, H.A., Rowe, M.C., 

Mandeville, C.W., and Meeker, G.P., 2008, Chemistry, mineralogy, and petrology of amphibole in Mount St. Helens 

2004–2006 dacite, chap. 32 of Sherrod, D.R., Scott, W.E., 

and Stauffer, P.H., eds., A volcano rekindled; the renewed 

eruption of Mount St. Helens, 2004–2006: U.S. Geological 

Survey Professional Paper 1750 (this volume).

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Fe-Ti oxide reequilibration of the 1991 Mount Unzen mixed 

magma: Journal of Volcanology and Geothermal Research, 

v. 89, p. 213–230, doi:10.1016/S0377-0273(98)00133-4.

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concentrations and fluxes based on melt inclusion and 

volcanic gas data: Journal of Volcanology and Geothermal Research, v. 140, nos. 1–4, p. 217–240, doi:10.1016/j.

jvolgeores.2004.07.023.

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and least squares method for solving petrologic mixing 

problems: Geological Society of America Bulletin, v. 81, 

p. 1995–2008.
726 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 32

1

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

2

 U.S. Geological Survey, Box 25046, Denver Federal Center, Denver, CO 

80225

3

 Department of Geosciences, 104 Wilkinson Hall, Oregon State University, 

Corvallis, OR 97331; now at Department of Geoscience, 121 Trowbridge Hall, 

University of Iowa, Iowa City, IA 52242

4

 American Museum of Natural History, Central Park West at 79th Street, 

New York, NY 10024

Chemistry, Mineralogy, and Petrology of Amphibole in 

Mount St. Helens 2004–2006 Dacite

By Carl R. Thornber1, John S. Pallister1, Heather A. Lowers2, Michael C. Rowe3, Charles W. Mandeville4, and 

Gregory P. Meeker2

Abstract

Textural, compositional, and mineralogical data are 

reported and interpreted for a large population of clinoamphibole phenocrysts in 22 samples from the seven successive 

dacite spines erupted at Mount St. Helens between October 

2004 and January 2006. Despite the uniformity in bulk composition of magma erupted since 2004, there is striking textural 

and compositional diversity among amphibole phenocrysts 

and crystal fragments that have grown from, partly dissolved 

in, or been accidentally incorporated in the new dacite. This 

study demonstrates that magma erupted throughout the current 

dome-building episode is the end product of small-scale, thorough mixing of multiple generations of crystal-laden magma. 

The mixed amphibole population provides important clues to 

magma conditions within the dacite magma reservoir prior to 

ascent and, to some extent, the dynamics of mixing and ascent.

The predominant amphibole in new dome rock ranges 

from moderate- to high-alumina tschermakite and magnesiohastingsite compositions. As substantiated by major- and 

trace-element geochemistry and barometry calculations, 

this compositional range of crystals, along with plagioclase, 

orthopyroxene, and iron-titanium oxide, is likely to have 

precipitated from dacite magma over a range of pressures and 

temperatures consistent with experimentally determined phase 

relations (~900°C to ~800°C between 100 MPa and ~350–400 

MPa or ~4-km and 13.5–15-km depth). Along with traceelement characteristics, textural and compositional data help 

to distinguish some low-alumina magnesiohornblende crystals 

as xenocrysts. The diverse range in composition of amphibole 

in all samples of 2004–6 dacite, and the complex zonation 

observed in many phenocrysts, suggests a well-mixed source 

magma with components that are subjected to repeated heating and (or) pressurization within this pressure-temperature 

window. Amphibole textural and compositional diversity 

suggest dynamic conditions in the upper-reservoir zone, which 

has been tapped steadily during ~2 years of continuous and 

monotonous eruption. This well-mixed crystal mush is likely 

to have been subjected to repeated injection of hotter magma 

into cooler crystal-laden magma while simultaneously assimilating earlier generations of dacitic roof material and surrounding gabbroic rock.

Decompression-related reaction rims around subhedral, 

rounded, resorbed, and fragmented amphibole phenocrysts, 

regardless of composition, indicate that this mixed-crystal 

assemblage was being broken, abraded, and dissolved in 

the magma as a result of mechanical mixing before and 

during early stages of ascent from conduit roots extending 

into a mushy cupola of the shallow reservoir. In the earliest 

lava samples (October 2004), amphiboles with <3-μm rims 

associated with a glassier matrix than later samples suggest a 

slightly faster ascent rate consistent with the relatively high 

eruptive flux of the earliest phases of dome extrusion. Reaction rim widths of ~5 μm on amphibole in all subsequently 

extruded lava result from a steady influx and upward transport 

of magma from 3.5–2.5-km to ~1-km depth at rates of ~600 

to ~1,200 m/day, through a conduit less than 10 m in radius. 

Slower ascent rates inferred from volumetric-flux and matrixcrystallization parameters are explained by a widening of the 

conduit to greater than 60 m radius within 1 km of the surface.
728 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Introduction

The cooled and degassed dacite magma that has been 

tapped continuously by the 2004–6 eruption at Mount St. 

Helens is a mixture of 45 percent phenocrysts and 55 percent 

dacite melt (Pallister and others, this volume, chap. 30, table 

4). As shown by Pallister and others (this volume, chap. 30), 

the whole-rock composition of lava erupted from 2004 to 

2006 is strikingly uniform (64.9±0.09 weight percent SiO2), 

and the dense-rock-equivalent (DRE) volume proportion 

of plagioclase (30–40 percent), amphibole (4–8 percent), 

orthopyroxene (3–5 percent), and iron-titanium-oxide (1–2 

percent) phenocrysts is consistent throughout the eruption. 

Similarity of bulk-rock and bulk-matrix compositions, as well 

as the experimental phase equilibria of Rutherford and Devine 

(this volume, chap. 31), suggest that the dominant phenocryst 

assemblage shares a dacite parentage. However, the isotopic, 

mineralogical, and textural diversity of phenocrysts leaves 

little doubt that the bulk compositional uniformity is the end 

result of repeated and thorough mixing of phenocryst-laden 

magma with preexisting crystalline material.

Petrologic models of magmatic processes for this eruption rely, in part, upon mineral-melt criteria that are clouded 

by a genetically mixed population of phenocrysts. The 

phenocrysts, including variably resorbed, reacted, and fragmented crystals, are discrete pieces of a complicated “dacite 

puzzle.” Some crystal populations look alike but did not 

necessarily crystallize together at the same depth, temperature, 

magma composition, and time. Detailed studies of phenocryst 

1 cm

Figure 1. Photograph showing freshly broken surface of SH325-1A, a typical sample of Mount St. Helens 2004–6 dacite lava.

morphology and chemistry provide petrologic context for 

establishing the genetic affinities of the crystals. This chapter 

provides detailed documentation of the potpourri of amphibole 

phenocrysts in the 2004–6 lava at Mount St. Helens.

All amphibole found in 1980–86 and 2004–6 Mount 

St. Helens dome lava is calcium-rich, monoclinic amphibole 

(clinoamphibole). In fresh gray pieces of plagioclase-rich dacite (fig. 1), amphibole phenocrysts are conspicuous as darkbrown to black, lustrous to semitranslucent flecks, chunks, 

and, locally, flow-aligned acicular blades.

The ubiquitous amphibole phase in 1980–86 and 2004–6 

Mount St. Helens dome rock incorporates all of the majorelement components of a water-saturated dacite melt and, 

thus, records changes in the pressure-temperature (P-T) paths 

and the compositional variations of host magmas at depths 

in excess of ~ 4 km, where amphibole is stable (~100 MPa, 

Rutherford and Devine, this volume, chap. 31). Clinoamphibole phenocrysts dissolve in magmas undergoing decompression and degassing during ascent. The nature and extent of 

amphibole breakdown reactions record the rates and paths of 

magma ascent from depth (Rutherford and Hill, 1993; Browne 

and Gardner, 2006; Buckley and others, 2006). In this study 

of newly erupted Mount St. Helens lava, we establish and 

correlate variations in texture, chemistry, and mineralogy 

among amphibole crystals, and we document and interpret 

the characteristics of amphibole reaction rims. Aluminum-inamphibole thermobarometry is applied to infer the relative P-T

conditions for the range of amphiboles likely to have crystallized from host dacite or similar magma. We aim to provide an 
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 729

“amphibolic” perspective toward understanding the character 

and disposition of a well-mixed, cooled, and degassed magma 

beneath Mount St. Helens during its near-steady-state transport to the surface from 2004 to 2006.

Samples Used in this Study

This investigation is based on microbeam studies of 22 

Mount St. Helens dacite samples erupted during the October 

2004 to January 2006 interval of spine emergence and dome 

growth (table 1). Aside from the samples derived unequivocally 

from newly extruded material (including the March 8, 2005, 

ash-fall sample, MSH05DRS3/8-4), the suite described here 

includes two unusual glassy dacite fragments (SH300-1A3, 

SH304-2G) dredged from debris shed by the emergent dome in 

early October 2004. Because these two samples are of uncertain origin and are uniquely glassy, their data, along with data 

for lithic inclusions of nonjuvenile andesite and dacite found 

in lava samples SH315-4, SH325-1A, and SH321-1, are 

discussed separately. Petrologic attributes of several samples 

reported here are also reviewed elsewhere in this volume; for 

example, the petrology of the early glassy fragment (SH3

04-2G) and the small, glassy andesite inclusion in SH315-4 

are discussed by Pallister and others (this volume, chap. 30). 

An overview and complete catalog of samples collected during 

the 2004–6 eruption, including sampling methods, locations, brief descriptions, and chemistry, is presented for dome 

samples by Thornber and others (2008) and for tephra samples 

by Rowe and others (2008).

Three older dome samples from the Cascades Volcano 

Observatory (CVO) archive were also analyzed to assess differences between clinoamphibole chemistry from the current 

eruption suite and the 1980–86 dacite (table 1). Compositional 

data on amphiboles in two gabbro xenoliths from the 1986 

dome rock (791-8a, 791-8b samples of Pallister and others, 

1991) provide a basis for identifying xenolith-derived crystal 

fragments among the diverse population of amphiboles in the 

2004–6 dacite.

Overview of Amphibole Data

Amphiboles in polished thin sections were selected for 

analysis without bias to size or morphology. Microbeam analyses and imagery used in this study were accomplished at three 

different laboratories (details of microbeam methods are in 

appendix 1). All amphibole analyses used in this investigation 

fit the mineralogical criteria established for calcic amphiboles 

by Leake and others (1997). The stoichiometric calculation 

used to assess cation site occupancy and proper nomenclature 

is discussed in a separate section below. Compositional data 

acquired include 458 major-element analyses of 399 individual 

crystals with 373 interior and 85 outer (20 μm) rim analyses. 

These data include 26 analyses of 17 crystals in lithic inclusions of andesite or dacite and 49 analyses of 28 crystals in 

glassy lithic samples. In addition, rare-earth elements were 

analyzed for 80 amphibole cores in nine samples spanning the 

2004–6 eruption. No amphibole trace-element data are currently available for 1980–86 lava or xenolithic inclusions. All 

chemical analyses of 2004–6 amphibole phenocrysts compiled for this investigation are tabulated in appendix 2, which 

appears only in the digital versions of this work (in the DVD 

that accompanies the printed volume and online at http://pubs.

usgs.gov/pp/1750). For purposes of the comparison with current eruptive products, major-element analyses of amphibole 

in 1980–86 Mount St. Helens dacite and in gabbroic xenoliths 

in the 1980–86 dome material also were compiled and are 

provided in appendix 2.

Maximum length and width measurements of 391 (98 

percent) of the analyzed phenocrysts were made using scaled 

scanning electron microscope (SEM) images. Although 

providing a general assessment of amphibole crystal sizes 

and aspect ratios, this method does not consider variations 

resulting from crystal orientation within the plane of the thin 

section. Measurements of the average width of amphibolebreakdown rims resulting from decompression (Rutherford 

and Hill, 1993; Browne and Gardner, 2006) were made on 195 

crystals having rims in contact with the 2004–6 dacite matrix. 

Care was taken to exclude any anomalous rim widths for crystals sectioned unevenly along their outer edges.

Amphibole Morphology and Reaction 

Textures

In typical new-dome material, amphibole grains range 

in size from phenocrysts <4 mm long to microphenocrysts 

ranging from 100 to 10 μm. An amphibole size distribution based upon approximate surface areas (length-by-width 

measurements) shows that nearly all of the smallest crystals 

observed (<1×103 μm2) are fragments with broken or irregular 

edges and that the amphibole population as a whole is more 

normally distributed about a range of 10–100×103 μm2

 (fig. 

2). Distinctly resorbed edges are observed among broken and 

unbroken crystals. A 62-percent majority of the amphibole 

phenocryst population is equant to slightly acicular (0–10 

percent acicularity) and 23 percent is distinctly acicular, with 

length/width ratios ranging from 2.5:1 to 5:1 (20–30 percent 

acicularity, fig. 3).

Morphologic variants of isolated amphibole grains in 

2004–6 dacite-lava samples include (1) subhedral, equant to 

acicular crystals with well-defined faces and typically, but 

not always, multiple, compositionally distinct growth zones 

(fig. 4); (2) angular fragments, variably rounded or embayed 

and commonly broken along cleavage planes that intersect 

normal growth zones (fig. 5); (3) well-rounded and seemingly 

abraded grains lacking substantially embayed or cuspate grain 

boundaries (fig. 6); and (4) irregularly shaped, rounded and 

embayed crystals or crystal fragments, commonly with fine 
Sample No. Spine No. Collection date Eruption date1 Sample type Rock ty

SH100 9/8/1981 9/6/1981 Dome talus Dacite SH133 6/5/1985 2/1–3/31/1983 Dome rock Dacite SH187 6/6/1989 5/30/1985 Dome rock Dacite SH300-1A3 1 10/20/2004 10/15/2004 Crater debris, either 1985 or Oct 2004 (?) Dense dark-gray glassy daciSH300-1C 1 10/20/2004 10/15/2004 Crater debris, either 1985 or Oct 2004 (?) Vesicular dark-gray dacite frSH302-1A 1 or 2 10/27/2004 10/14/2004 Crater debris, either1985 or Oct 2004 (?) Vesicular dark-gray dacite frSH304-2A 3 11/4/2004 10/18/2004 Dome talus and debris Vesicular “hot-pink” dacite SH304-2G 3 11/4/2004 10/18/2004 Dome talus and debris Dense dark-gray glassy daciSH305-1 3 1/3/2005 11/20/2004 Spine margin Vesicular gray dacite SH306-A 4 1/14/2005 12/15/2004 Dome talus Vesicular pink-gray dacite SH308-3 4 2/22/2005 1/21/2005 Dome talus Vesicular gray dacite SH309-1C 4 2/22/2005 1/13/2005 Spine margin Vesicular gray-pink dacite, fSH311-1B 4 1/19/2005 1/16/2005 Ballistic block, 1/16/2005 explosion Vesicular gray dacite MSH05DRS_3_9_4 4 3/9/2005 3/8/2005 Coarse fragments in tephra-fall deposit from 3/8/2005 explosion Vesicular gray-pink dacite SH312-1 4 4/10/2005 3/8/2005 Ballistic block, 3/8/2005 explosion Vesicular gray dacite SH314-1A1 5 4/19/2005 4/17/2005 Gouge fragment Vesicular dacite fragment SH315-3 4 4/19/2005 4/1/2005 Dome talus, 4/18/2005 collapse Dense gray dacite SH315-4 4 4/19/2005 4/1/2005 Dome talus, 4/18/2005 collapse Dense gray daciteSH317-4 5 6/15/2005 5/1/2005 Dome talus Dense pink-gray daciteSH319-1 5 7/13/2005 5/15/2005 Dome talus Dense gray-pink dacite, foliSH321-1 6 8/19/2005 8/10/2005 Dome talus, 8/19/2005 collapse Dense gray-pink dacite, “pinSH323-2, -3 6 10/18/2005 9/10/2005 Dome talus, 8/19/2005 collapse Dense gray dacite

SH324-1A 7 12/15/2005 12/5/2005 Spine margin Vesicular gray-pink daciteSH324-3 7 12/15/2005 12/5/2005 Spine margin Dense gray daciteSH325-1 7 2/7/2006 12/20/2005 Dome talus Dense gray dacite

Table 1. Summary of Mount St. Helens 2004–2006 samples used in this investigation.

1 Estimate of when sampled lava was extruded from vent, whose altitude is assumed to be that of the crated floor below Crater Glacier.
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 731

0

20

40

60

80

100

120

140

<0.1

>0.1

 1.0

TO

1

10

TO

10

50

TO

50

100

TO

100

500

TO

500

1,000

TO >1,000

AMPHIBOLE SURFACE AREA, IN μm2×103 

NUMBER OF GRAINS COUNTED

0

10

20

30

40

50

60

70

0 to 10 >10 to 20 >20 to 30 >30 to 40 >40 

AMPHIBOLE ACICULARITY, IN PERCENT

NUMBER OF GRAINS COUNTED

Figure 3. Histogram of acicularity of amphibole phenocrysts in 

Mount St. Helens 2004–6 dacite lava. Percent acicularity = [100 × 

length/(length + width)] – 50. (Equant crystals have zero-percent 

acicularity).

Figure 2. Histogram of size distribution of amphibole 

phenocrysts in Mount St. Helens 2004–6 dacite lava presented 

as approximate surface area. The size distribution of fragmented 

crystals, shown in gray, is compared to that of the overall 

population, shown in black.

to coarse cuspate margins indicative of magmatic resorption 

(fig. 7). The patterns of compositional zoning apparent in 

SEM backscattered electron images of many crystals or fragments typically define subparallel growth bands. In numerous 

cases, regardless of grain morphology, zoning patterns have 

an irregular or patchy appearance. This characteristic has been 

attributed to late-stage igneous (near-solidus) diffusion (Hammarstrom and Zen, 1992) but could result from patchy dissolution and regrowth (M.J. Rutherford, written commun., 2006).

Many amphibole crystals or fragments in the new Mount 

St. Helens dacite have glass inclusions (fig. 4) or partially to 

wholly crystallized melt inclusions (glass±opx±plg±oxides; 

for example, fig. 5B), suggestive of variable cooling or reaction rates among a mixed phenocryst population. Amphibole 

phenocrysts of subhedral and resorbed morphology also are 

present in glomerophyric clusters, coupled during growth with 

various combinations of plagioclase, orthopyroxene, or oxide 

crystals (fig. 8). Although some of the more pristine-looking 

glomerophyric aggregates appear as though they grew in the 

dacite host, those cored by resorbed amphiboles may have 

resulted from prolonged amphibole-melt reaction, as discussed 

further below.

Because of similarities in texture and degree of reaction, there might be no definitive textural distinction between 

cognate and xenocrystic aggregates, such as those derived 

from gabbroic xenolith fragmentation. The latter could appear 

relatively pristine (unreacted or unresorbed) if exposed to the 

degassed host melt during xenolith fragmentation at nearsolidus temperatures and shallow depths.

The edges of all amphibole crystals in contact with 

matrix in 2004–6 dacite lava are armored by an aureole of 

acicular microlites of plagioclase and orthopyroxene (±Fe-Ti 

oxide). As demonstrated by experiments, such rims result from 

reaction of clinoamphibole with dacite melt during decompression and degassing (Rutherford and Hill, 1993; Browne 

and Gardner, 2006). Rim widths are a function of the rate 

and path of ascent of the host magma during transport to the 

surface from a minimum PH2

O of 100 MPa where amphibole is 

stable (>3.5–4 km depth at Mount St. Helens).

Essentially all of the amphibole crystals measured have 

decompression reaction rims between 3 and 8 μm thick; 

average reaction-rim width is 5±1 μm. Typical reaction rims 

pseudomorphically replace the outer edges of subhedral, 

rounded, or resorbed crystals and fragments (fig. 9). Exceptions to the normal reaction-rim thickness are observed in 

the earliest dome sample (SH304-2A, fig. 9B), which has 

<3-μm-thick rims, and the glassy lithic fragment (SH304-2G, 

fig. 9A), in which amphibole crystals lack reaction rims. 

Atypically thick rims (30–200 μm) are found among isolated 

amphibole crystals in 2004–6 dacite. These rims are defined 

by relatively coarse-grained plagioclase-orthopyroxene-oxide 

intergrowths that surround embayed (resorbed) amphibole 

grains (fig. 10).

An unusual but recurrent texture observed in SH314-1 

and DRS_3_9_4 indicates amphibole reacting with the host 

melt to form second-generation amphibole + liquid (fig. 11). 

Such amphibole-to-amphibole reaction reflects exposure to a 

hotter or chemically incompatible melt under amphibole saturation conditions. As discussed further within the context of 

the compositions of these crystals, this amphibole-to-amphibole reaction is likely to occur at depths near the low-pressure 

limits of amphibole stability (100 MPa).

Amphibole phenocrysts in lithic inclusions (SH315-4inc, 

SH325-1Ainc, and SH321-1inc) have reaction-rim characteristics that differ from those of the host 2004–6 dacite. The 

SH315-4 glass-bearing inclusion has distinctly acicular amphibole intergrown with plagioclase and minor clinopyroxene, 

and it has an andesitic bulk composition (Pallister and others, 
732 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 4. SEM backscattered electron images of analyzed 

subhedral amphibole grains (AM) in Mount St. Helens 

2004–6 dacite lava. Glass inclusions (GL INC) indicated 

where confirmed by microbeam analysis. The sample and 

image numbers shown on the lower right side in each image 

correspond to analysis numbers in appendix 2. A, Sample 

SH305-1. B, Sample SH315-3. C, Sample SH325-1.

Figure 5. SEM backscattered electron images of analyzed 

broken amphibole grains (AM) in Mount St. Helens 2004–6 dacite 

lava. Orthopyroxene and glass inclusions (OPX + GL INC) are 

indicated where confirmed by microbeam analysis. The sample 

and image numbers shown on the lower right side in each 

image correspond to analysis numbers in appendix 2. A, Sample 

SH304-2A. B, Sample SH304-2A. C, Sample SH319-1.

SH305-1 MR4-3

AM

GL INC

A

SH325-1 F3

SH315-3 F5

AM

GL INC

AM

GL INC

B

C

100 μm 

200 μm 

100 μm 

SH319-1 MR2-1

AM

OPX + GL INC

100 μm

AM

SH304-2A MR1-1

SH304-2A MR1-2

AM

A

B

C

100 μm

100 μm
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 733

Figure 6. SEM backscattered electron images of analyzed 

rounded or abraded amphibole grains (AM) in Mount St. Helens 

2004–6 dacite lava. The sample and image numbers shown 

on the lower right side in each image correspond to analysis 

numbers in appendix 2. A, Sample SH304-2A. B, Sample SH305-1. 

C, Sample SH324-3.

Figure 7. SEM backscattered electron images of resorbed 

amphibole crystals (AM) with embayed and cuspate margins 

in Mount St. Helens 2004–6 dacite lava. The sample and 

image numbers shown on the upper right side in each image 

correspond to analysis numbers in appendix 2. A, Sample 

SH304-2A. B, Sample SH306A-1. C, Sample SH319-1.

100 μm

A

B

C

SH304-2A MR4-1

SH305-1 MR4-8

SH324-3 MR1-2

AM

AM

AM

100 μm

50 μm

SH304-2A MR5-2

SH306A-1 MR2-2

SH319-1 MR2-3

AM

AM

AM

A

B

C

100 μm

100 μm

100 μm
734 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 8. SEM backscattered electron images of 

amphibole grains (AM) intergrown with plagioclase (PLG) in 

Mount St. Helens 2004–6 dacite lava. The sample and image 

numbers shown on the upper right side in each image 

correspond to analysis numbers in appendix 2. A, Sample 

SH305-1. B, Sample SH315-3. C, Sample SH315-4.

Figure 9. SEM backscattered electron images showing reaction 

rims around amphibole phenocrysts in Mount St. Helens 2004–6 

dacite lava. The sample and image numbers shown on the lower 

right side in each image correspond to analysis numbers in 

appendix 2. A, No reaction rims present on amphiboles in glassy 

lithic fragment, SH304-2G. B, Rim widths are ~3 μm in October 

2004 spine lava, SH304-2A. C, Typical 5-μm rims in all subsequent 

2004–6 spine samples (SH305-1 shown here) are composed of an 

acicular microcrystalline intergrowth of plagioclase (dark colored), 

orthopyroxene (lighter colored), and Fe-Ti oxide (white).

PLG

SH315-4 MR2-4

SH315-3 F4

SH305-1 MR4-1

PLG

PLG

AM

AM

AM

A

B

C

100 μm

100 μm

100 μm

C

A

B

SH304-2G MR1-1B

SH304-2A F1B

SH305-1A 4C

50 μm

50 μm

10 μm
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 735

Figure 10. SEM backscattered electron images 

showing rare coarse-grained reaction rims (plagioclase + 

orthopyroxene + Fe-Ti oxide) around amphibole phenocryst 

remnants (AM) in Mount St. Helens 2004–6 dacite lava. The 

sample and image numbers shown on the lower right side in 

each image correspond to analysis numbers in appendix 2. 

A, Sample SH315-3. B, Sample SH317-4. C, Sample SH323-1.

Figure 11. SEM backscattered electron images showing 

examples of amphibole-to-amphibole+melt reaction in 

Mount St. Helens 2004–6 dacite lava. The sample and image 

numbers shown on the upper right side in each image 

correspond to analysis numbers in table 2 and appendix 2. A, 

In sample SH314-1, amphibole with ~12.5 percent Al2O3 (AM1) 

is surrounded by aggregate of amphibole crystals (AM2) 

with ~7.2 percent Al2O3 and high fluorine (~0.18 percent) and 

glass (dark gray). B, In sample DRS_3_9_4, amphibole with 

~12.2 percent Al2O3 (AM1) is surrounded by an aggregate of 

amphibole crystals (AM2) with ~6.9 percent Al2O3 and high 

fluorine (~0.25 percent) and glass (dark gray). 

A

B

C

100 μm

200 μm

200 μm

SH315-3 F9

SH317-4 F7

SH323-1 JP6

AM

AM

AM

B

SH314-1 F3

DRS-3/9-4

AM2 (LOW-AL)

AM1 (MED-AL) 

AM1 (MED-AL) 

AM2 (LOW-AL)

A

100 μm

250 μm
736 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

this volume, chap. 30). Reaction rims ~5-μm wide are found 

on the inclusion edge, where amphibole crystals are in contact 

with the dacite matrix, but no such rims occur along interior 

contacts between amphibole and glass. Inclusion SH325-1Ainc 

has a bulk texture similar to that of SH315-4inc, but the matrix 

lacks a significant glass component. Reaction rims are similarly 

restricted to grains in contact with the host dacite. These are the 

only andesite inclusions found among the well-scrutinized suite 

of 2004–6 dacite samples. They are small (<1.5 cm) and are 

likely accidental fragments derived from conduit walls. We do 

not consider them evidence of mafic-magma mixing.

Dacite inclusions such as SH321-1inc contain resorbed 

amphibole crystals with broad reaction rims similar to those 

found in the bimodal population in 1980–86 dacite (Rutherford and Hill, 1993; Cashman and McConnell, 2005). This 

inclusion also is distinguished from the 2004–6 host lava 

by the presence of clinopyroxene microphenocrysts in the 

matrix around amphibole, which is suggestive of isobaric 

heating of a cooler crystal-laden magma or of slow decompression of a newer hotter one (see Rutherford and Devine, 

this volume, chap. 31, fig. 10). It is theoretically possible that 

such fragments could have originated from eruption-related 

magmatic processes. However, like the andesite inclusions, 

these rounded fragments are more likely to be inconsequential 

xenoliths eroded from the shallow conduit walls.

We discern some fundamental characteristics of the 

2004–6 dacite magma from the textures of its amphibole phenocrysts. The well-dispersed population of texturally diverse 

amphibole described above is a strong indication that the 

uniform composition of the dacite is inherited from a process of small-scale mingling and intermixing of crystal-laden 

magmas at depth. The consistent presence of variably zoned, 

broken, abraded, and resorbed amphibole crystals suggests 

that a number of temporally (and spatially) distinct crystallization environments are represented. Because similar 5-μm 

reaction rims are observed around all morphologic variants 

of amphibole, crystal fragmentation, abrasion, and resorption must have occurred at depth after assembly of the diverse 

crystal population and during initial stages of decompression 

with upward transport. Further discussion of amphibole-melt 

reaction textures and implications for dynamics of magma 

ascent is presented later in this chapter.

In the following sections, the chemistry and mineralogy 

of this textural medley of amphibole phenocrysts are reviewed 

to ascertain genetic distinctions and to gain insight into magmatic conditions associated with the 2004–6 eruption.

Amphibole Major-Element Chemistry

Amphibole phenocrysts analyzed to date in 2004–6 

Mount St. Helens dacite (fig. 12; table 2; appendix 2) are all 

clinoamphibole with a calcium content of ~11.0 weight percent (11.0±0.3 percent CaO). Their silica and alumina contents 

span a continuous range that varies inversely between average 

end members with high silica and low alumina (47.5 percent 

SiO2, 6.4 percent Al2O3) and low silica and high alumina (41.8 

percent SiO2, 15.0 percent Al2O3) (fig. 13A). The prevalent 

38

40

42

44

46

48

50

5

10

15

20

0

1

2

3

4

5

0

0.2

0.4

0.6

0.8

SiO2 FeTO Al MgO 2O3 CaO TiO2 Na2O K2O

WEIGHT PERCENT

MnO

Range

Average

±1σ

Figure 12. Average, range, and standard deviation of major elements analyzed in Mount St. Helens 2004–6 amphibole 

phenocrysts (excluding data from crystals in lithic inclusions and glassy lithic fragments). FeTO is all iron calculated as Fe2+. 
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 737

40

42

44

46

48

50

12

14

16

18

10

12

14

16

18

5 7 9 11 13 15

Al2O3, IN WEIGHT PERCENT

FeTO, IN WEIGHT PERCENT MgO, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT 

A

Low-Al Med-Al High-Al

C

B

Figure 13. Alumina variation diagrams for Mount St. Helens 

2004–6 amphibole phenocrysts (excluding data from crystals 

in lithic inclusions and glassy lithic fragments); average and 

1σ standard deviation in red. Curves are high-order polynomial 

fits. A, Silica. B, Magnesia. C, Iron oxide (FeTO calculated as 

100 percent ferrous iron). Reversals in FeTO and MgO with 

increasing Al2O3 define low-Al (6 to 9 percent), medium-Al (9 to 

12 percent), and high-Al (>12 percent) compositions.

2004–6 amphibole compositional range, defined as ±1σ variation from the mean, is from 42.1 to 43.8 percent SiO2 and 11.4 

to 9.7 percent Al2O3.

Three compositional trends are evident within the silica/

alumina range of amphibole crystals in dome dacite (and domederived ash) by the patterns of Mg and Fe variation with Al 

(figs. 13B, 13C). The low-alumina range (low-Al), from 6 to 

~9 percent Al2O3, has a sympathetic increase in FeTO (all iron 

calculated as Fe2+) averaging from 12 to 15.5 percent associated 

with decreasing MgO from 16 to 12 percent. Over the midalumina range (medium-Al), from ~9 to 12 percent Al2O3, average FeTO decreases to 12 percent and MgO increases to 12–16 

percent. High-alumina (high-Al) amphibole ranges from ~12 to 

15 percent Al2O3, with a corresponding average FeTO increase 

to 16 percent and a MgO decrease to ~11 percent.

Amphibole phenocrysts of high-Al, medium-Al, and lowAl compositions are present and distributed similarly in nearly 

all 2004–6 samples examined, as well as in 1980–86 dacite 

samples (fig. 14A). Our data reveal that amphibole in 2004–6 

and 1980–86 dacites have the same range of major-element 

compositions. The low-Al amphiboles in old and new dacite 

overlap the phenocryst compositions in gabbroic xenoliths 

(fig. 14A). Amphibole crystals in the SH304-2G glassy lithic 

fragment are chiefly medium-Al in composition but display 

a low- to high-Al range similar to that of the 2004–6 population (fig. 14B). The high-Al amphibole compositions of lithic 

inclusions (SH315-4inc and SH325-1Ainc) are similar to those 

5 7 9 11 13 15

Al2O3, IN WEIGHT PERCENT

10

12

14

16

FeTO, IN WEIGHT PERCENT FeTO, IN WEIGHT PERCENT 

18

20

10

12

14

16

20

18

Lithic Inclusions

SH321-1 INC

SH315-4 INC

SH325-1A INC

Gabbroic

xenocrysts

MSH 2004–2006 MSH 1980–1986

A

B Glassy Fragments

SH300-1A3

SH304-2G

Figure 14. Iron oxide versus alumina in Mount St. Helens 

amphibole phenocrysts. A, 2004–6 amphibole phenocrysts 

(open circles and gray-shaded area) compared to those in 

1980–86 Mount St. Helens dacite and gabbroic xenocrysts. 

B, Amphiboles in lithic inclusions and glassy lithic fragments; 

shaded area corresponds to field of 2004–6 amphibole 

phenocrysts shown in A.
Sample No.-- SH304-2A SH305-1 SH306-A DRS_3_9_4 SH314-1 SH315-3 SAnalysis No. ------------ OS- MR4-3 AM- G24C AM- G24R OS- MR3-2 OS- MR3-3 OS- MR3-4 OS- MR2-2 OS- MR2-5 GS- F3P13 GS- F3P9 GS- F3P1 GS- F3P5 GS- F16G1 P1 GS- F16G1 P2 GS- F16G1 P3 GS- F1G

P1 

Position------- core core rim core core core core core core rim core rim inZ core rim core 

SiO2 48.4 42.0 44.0 43.1 42.9 46.9 42.6 47.8 43.1 47.9 42.6 47.1 43.7 44.0 45.1 42.2 

TiO2 1.54 2.53 2.39 2.37 2.24 2.09 1.97 1.46 2.21 1.46 2.37 1.48 1.46 1.61 0.88 2.3Al2O3 7.4 12.8 12.5 14.2 14.0 8.6 13.1 7.9 12.2 6.9 12.4 6.9 9.4 8.5 8.0 12.1 FeO 13.7 11.3 10.7 14.6 12.0 13.7 14.8 13.4 10.8 12.7 10.9 13.1 16.3 16.2 16.3 10.4 MnO 0.3 0.1 0.1 0.12 0.13 0.16 0.20 0.28 0.12 0.23 0.12 0.24 0.35 0.29 0.35 0.1MgO 15.2 14.7 15.5 11.6 14.5 14.4 12.0 15.0 14.8 14.7 14.9 15.5 12.4 12.7 13.3 15.3 CaO 11.0 11.2 11.4 10.4 10.3 11.3 11.2 10.6 10.8 10.8 11.5 11.2 10.4 10.4 9.8 11.1 Na2O 1.5 2.4 2.4 2.4 2.6 1.8 2.1 1.7 2.5 1.5 2.8 1.8 1.8 2.1 1.6 2.9 K2O 0.28 0.30 0.32 0.38 0.29 0.34 0.35 0.23 0.30 0.22 0.29 0.22 0.26 0.31 0.20 0.3F 0.02 0.08 0.10 0.00 0.00 0.01 0.01 0.02 0.06 0.25 0.07 0.08 0.22 0.00 0.14 0.1Cl 0.05 0.02 0.01 0.03 0.02 0.05 0.03 0.04 0.02 0.04 0.01 0.04 0.02 0.02 0.02 0.0O = F,Cl 0.00 0.04 0.05 0.01 0.00 0.01 0.00 0.00 0.03 0.11 0.03 0.04 0.10 0.00 0.06 0.0Total 99.19 97.37 99.38 99.19 99.11 99.33 98.36 98.39 96.94 96.63 97.99 97.60 96.20 96.16 95.56 96.9T-Site Si+4 6.869 6.074 6.209 6.176 6.088 6.690 6.189 6.830 6.242 6.969 6.149 6.862 6.490 6.587 6.717 6.1T-Site IVAl+3 1.131 1.926 1.791 1.824 1.912 1.310 1.811 1.170 1.758 1.031 1.851 1.138 1.510 1.413 1.283 1.8C-Site VIAl+3 0.105 0.250 0.284 0.576 0.437 0.141 0.432 0.165 0.331 0.153 0.258 0.041 0.128 0.089 0.114 0.2C-Site Ti+4 0.165 0.275 0.253 0.255 0.239 0.224 0.215 0.156 0.241 0.160 0.258 0.162 0.163 0.181 0.099 0.2C-Site Fe+3 0.620 0.772 0.660 0.643 0.704 0.585 0.673 0.667 0.642 0.571 0.591 0.370 0.963 0.621 0.896 0.6C-Site Mg+2 3.212 3.161 3.272 2.487 3.070 3.052 2.604 3.184 3.187 3.192 3.202 3.355 2.752 2.828 2.946 3.3C-Site Fe+2 0.901 0.527 0.522 1.036 0.546 0.999 1.077 0.826 0.599 0.925 0.691 1.072 0.995 1.281 0.946 0.5B-Site Fe+2 0.102 0.066 0.080 0.070 0.179 0.049 0.043 0.110 0.071 0.049 0.036 0.153 0.064 0.125 0.187 0.0B-Site Mn+2 0.033 0.011 0.010 0.015 0.015 0.019 0.025 0.034 0.014 0.028 0.015 0.029 0.044 0.037 0.044 0.0B-Site Ca+2 1.674 1.734 1.720 1.595 1.571 1.722 1.742 1.615 1.682 1.678 1.778 1.741 1.660 1.674 1.560 1.7B-Site Na+ 0.192 0.189 0.191 0.319 0.235 0.209 0.190 0.242 0.233 0.245 0.171 0.076 0.232 0.164 0.209 0.1A-Site Na+ 0.220 0.471 0.465 0.341 0.470 0.285 0.403 0.224 0.481 0.191 0.605 0.438 0.275 0.444 0.247 0.6A-Site K+ 0.051 0.056 0.058 0.070 0.053 0.061 0.064 0.042 0.055 0.041 0.053 0.041 0.049 0.060 0.038 0.0Total 15.27 15.53 15.52 15.41 15.52 15.35 15.47 15.27 15.54 15.23 15.66 15.48 15.32 15.50 15.28 15.6

Table 2. Selected electron microprobe analyses of amphibole phenocrysts in Mount St. Helens 2004–2006 dacite.[Laboratories identified by first two letters in analysis number: AM, American Museum of Natural History; GS, U.S. Geological Survey, Denver; OS, Oregon State University. of analyzed spot: inZ, inner zone; outZ, outer zone. Nomenclature from Leake and others (1997): mghst, magnesiohastingsite; tscher, tschermakite; mghb, magnesiohornblendeparg, pargasite. See appendix 2, table A2, for all analytical data.]
Sample No-- SH319-1 SH321-1 SH323-2 SH323-3 SH324-3 

Analysis No.----------- OS-MR2-1

OSMR2-3 OS- MR3-1 GS- F0P3 GS- F2P3 OS- MR1-3 OS- MR2-3 OS- MR3-2 GS- G1-2 GS- G1-7 GS- G1-6 GS- G1-8 GS- F0P3 GS- F0P8 OS- MR1-1 OS- MR1-

Position------- core core core core core core core core core inZ outZ rim core core core core 

SiO2 44.3 45.2 43.4 43.7 43.5 44.8 46.4 42.8 43.6 43.7 44.2 43.8 42.7 45.6 44.5 43.5 

TiO2 2.24 2.13 2.38 2.48 2.74 2.55 1.31 2.79 2.19 2.20 2.17 2.78 2.42 1.11 2.51 2.27 

Al2O3 11.7 12.3 13.2 10.8 9.8 10.3 9.2 14.2 10.9 12.3 11.7 10.1 11.6 9.4 11.0 13.4 

FeO 12.3 10.6 11.7 11.1 13.7 13.7 17.0 13.5 13.6 11.6 10.5 12.9 14.8 15.6 13.1 12.7 

MnO 0.17 0.12 0.10 0.14 0.15 0.19 0.27 0.13 0.19 0.15 0.12 0.19 0.20 0.26 0.16 0.12 

MgO 14.6 15.6 14.5 15.0 13.8 13.9 12.5 12.6 13.2 14.5 15.2 13.7 12.3 13.5 13.8 13.9 

CaO 11.2 10.9 11.1 11.1 10.7 11.0 11.0 10.9 10.8 10.8 10.9 10.5 10.8 10.0 11.4 11.1 Na2O 2.2 2.3 2.4 2.3 2.7 2.0 1.9 2.4 2.3 2.3 2.3 2.3 2.6 1.9 2.1 2.4 

K2O 0.29 0.29 0.32 0.25 0.29 0.31 0.24 0.36 0.25 0.30 0.28 0.30 0.29 0.23 0.31 0.34 F 0.00 0.00 0.00 0.02 0.03 0.00 0.00 0.01 0.69 0.16 0.10 0.19 0.02 0.04 0.01 0.00 

Cl 0.02 0.01 0.02 0.18 1.55 0.04 0.02 0.02 0.02 0.01 0.01 0.02 0.11 0.16 0.02 0.02 

O=F,Cl 0.00 0.00 0.00 0.05 0.35 0.01 0.00 0.00 0.29 0.07 0.04 0.08 0.03 0.05 0.00 0.00 

Total 98.88 99.53 99.16 97.17 98.59 98.72 99.83 99.71 97.34 98.08 97.31 96.61 97.85 97.82 98.91 99.64 

T-Site Si+4 6.312 6.343 6.165 6.341 6.395 6.430 6.665 6.097 6.388 6.261 6.360 6.431 6.270 6.609 6.381 6.16

T-Site IVAl+3 1.688 1.657 1.835 1.659 1.605 1.570 1.335 1.903 1.612 1.739 1.640 1.569 1.730 1.391 1.619 1.83C-Site VIAl+3 0.268 0.369 0.369 0.192 0.087 0.172 0.213 0.473 0.267 0.346 0.334 0.172 0.277 0.210 0.240 0.40

C-Site Ti+4 0.240 0.225 0.254 0.270 0.303 0.276 0.142 0.298 0.242 0.238 0.234 0.307 0.267 0.120 0.271 0.24

C-Site Fe+3 0.700 0.615 0.671 0.613 0.492 0.665 0.687 0.623 0.620 0.685 0.620 0.610 0.585 0.884 0.573 0.68C-Site Mg+2 3.092 3.273 3.074 3.233 3.017 2.968 2.666 2.681 2.877 3.108 3.251 2.990 2.682 2.917 2.958 2.92C-Site Fe+2 0.697 0.514 0.631 0.673 1.100 0.916 1.292 0.926 0.995 0.623 0.560 0.921 1.190 0.864 0.959 0.74B-Site Fe+2 0.063 0.117 0.081 0.060 0.091 0.068 0.060 0.064 0.055 0.083 0.077 0.052 0.043 0.148 0.045 0.08B-Site Mn+2 0.021 0.014 0.012 0.017 0.019 0.024 0.033 0.016 0.024 0.018 0.015 0.023 0.025 0.032 0.020 0.01B-Site Ca+2 1.709 1.645 1.689 1.724 1.692 1.686 1.691 1.667 1.690 1.665 1.676 1.657 1.700 1.559 1.746 1.67B-Site Na+ 0.207 0.224 0.218 0.200 0.198 0.222 0.217 0.253 0.231 0.234 0.232 0.267 0.232 0.260 0.190 0.22A-Site Na+ 0.391 0.391 0.446 0.449 0.563 0.343 0.324 0.399 0.425 0.413 0.399 0.384 0.513 0.270 0.399 0.42

A-Site K+ 0.053 0.053 0.059 0.047 0.054 0.057 0.045 0.066 0.047 0.054 0.051 0.056 0.053 0.042 0.056 0.06

Total 15.44 15.44 15.50 15.50 15.62 15.40 15.37 15.46 15.47 15.47 15.45 15.44 15.57 15.31 15.46 15.49 

Name tscher tscher mghst tscher mghst tscher mghb tscher tscher tscher tscher tscher mghst mghb tscher tsche

Table 2. Selected electron microprobe analyses of amphibole phenocrysts in Mount St. Helens 2004–2006 dacite.—Continued[Laboratories identified by first two letters in analysis number: AM, American Museum of Natural History; GS, U.S. Geological Survey, Denver; OS, Oregon State University. Rof analyzed spot: inZ, inner zone; outZ, outer zone. Nomenclature from Leake and others (1997): mghst, magnesiohastingsite; tscher, tschermakite; mghb, magnesiohornblende; parg, pargasite. See appendix 2, table A2, for all analytical data.]
740 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

CORES INNER ZONES OUTER ZONES RIMS

0

10

20

30

40

50

60

70

80

Al2O3, IN WEIGHT PERCENT

NUMBER OF ANALYSES

0

10

0

10

0

10

20

0

10

0

10

0

10

0

10

20

0

10

Oct. 2004–Dec. 2005

Dec. 2004–Jan. 2005

Oct.–Nov. 2004

Mar.–Apr. 2005

May–Aug. 2005

Sep.–Dec. 2005

Intergrown with plagioclase

0

10

20

Fragmented

Resorbed

Glass inclusion bearing

0

10

Intergrown with ilmenite

6 to 7 7 to 8 8 to 9 9 to 10 10 to 11 11 to 12 12 to 13 13 to 14 14 to 15 15 to 16

NUMBER OF ANALYSES NUMBER OF ANALYSES

A

B

C

Figure 15. Histograms showing the frequency distribution of Al2O3 in cores, inner zones, 

outer zones, and rims of amphibole crystals in Mount St. Helens 2004–6 dacite lava (excluding 

crystals in glassy lithic fragments and lithic inclusions). A, For all 2004–6 amphiboles. B, For 

consecutive time intervals. C, For phenocrysts with specific textural attributes.
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 741

crystallized from andesite at high pressures (for example, 

Allen and Boettcher, 1978; Grove and others 2003). High-Al 

amphibole compositions, along with differences in reaction 

textures, are consistent with these inclusions originating from 

an andesite magma (Pallister and others, this volume, chap. 

30), but there is no clear evidence in support of a comagmatic 

relation between the andesitic magma and the host dacite.

The one dacitic inclusion studied (SH321-1inc) has a 

range of amphibole chemistry similar to that of 1980–86 and 

2004–6 lava. As inferred from the presence of clinopyroxene 

and the distinctly thick amphibole reaction rims (see reactionrim discussion), this inclusion could be an accidental piece of 

shallow 1980s conduit material.

The frequency distribution of alumina concentrations in 

cores, inner zones, outer zones, and rims of amphiboles in the 

2004–6 dacite indicate an approximately normal distribution 

about the mean value of 11.4±1.7 percent Al2O3 (fig. 15). A 

similar distribution is observed for a smaller population of 

1980–86 amphiboles, but amphiboles in gabbroic xenoliths 

overlap the low end of the 2004–6 population (fig. 14A).

Correlations between alumina content and size of 

amphibole crystals reveal a normal distribution about the mean 

alumina content of ~11 percent and about a median size range 

within 10–100×103 μm2

 (fig. 2), suggestive of compositional 

and morphologic consistency of the mean population. In 

contrast, most of the crystal fragment population (gray bars 

in fig. 2) has low-Al compositions, supporting the likelihood 

of a xenocrystic origin. There is no clear correlation between 

alumina content and acicularity. Both equant and relatively 

acicular populations have nearly similar distributions about the 

alumina average.

The abundance of low- to high-Al amphiboles, just as the 

overall amphibole modal abundance and bulk-dacite composition, has not changed throughout the duration of the 2004–6 

eruption (fig. 15). Furthermore, the same overall distribution 

of alumina contents is apparent among amphiboles of similar 

texture or mineral associations (fig. 15).

Multiple analyses from core to rim were obtained for 

numerous amphibole crystals. The compositional variation of 

amphibole rims within and among samples is similar to the 

broad range of core compositions and the amphibole population as a whole (fig. 16). Detailed analyses of amphiboles 

with multiple growth zones reveal some systematic tendencies 

within limited populations of crystals. For example, analyses 

of cores, inner zones, and outer zones of several SH323-3 

amphiboles vary widely and are without obvious correlations 

of core-to-rim progressions between crystals, but the rims all 

trend toward an average 10.8±0.6 percent Al2O3 (fig. 16A). 

This convergence of rims toward the amphibole composition to last equilibrate with the host magma is within the 1σ

uncertainty of the 11.4 percent average of Al2O3 in all 2004–6 

amphiboles. Such is not the case for core-to-rim analyses of 

crystals in SH325-1A (fig. 16B), which are similar to the overall broad dispersion of rim data. The variable rim compositions must reflect changing conditions in the magma reservoir 

and during ascent from it.

Amphibole Rare-Earth-Element (REE) 

Chemistry

The REE concentrations in all 2004–6 dacite amphibole phenocrysts display middle rare-earth-element (MREE) 

enrichment patterns typical of igneous clinoamphiboles (fig. 

17; appendix 2). Similar MREE-enriched patterns for different 

crystals in each sample are distinguished by varying degrees 

of overall REE enrichment, accompanied by progressive Eu 

depletion. Like other amphibole characteristics discussed 

above, there are no clear distinctions of REE patterns among 

samples over time or, in most cases, with textural attributes. 

However, a distinct trend of overall REE enrichment (typified by La) and Eu depletion (relative to Nd) with decreasing 

alumina content in medium- to low-Al amphibole crystals is 

clearly observed (fig. 18).

SH323-3 

10

12

14

16

FeTO, IN WEIGHT PERCENT FeTO, IN WEIGHT PERCENT 

18

20

5 7 9 11 13 15

Al2O3, IN WEIGHT PERCENT

10

12

14

16

20

18

CORES INNER ZONES OUTER ZONES RIMS

SH325-1A 

A

B

Figure 16. Iron oxide (FeTO) versus alumina within individual 

amphibole phenocrysts in Mount St. Helens 2004–6 dacite 

lava. Arrows connect core, inner-zone, outer-zone, and rim 

compositions. Shaded area is the field of 2004–6 amphibole 

phenocrysts from figure 14A. A, Interior compositions of 

amphiboles vary widely, but rim compositions converge 

toward an average of 10.8±0.6 percent Al2O3. B, Example of 

amphiboles that do not show consistent variations.
742 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Concentration patterns for the REE-enriched and Eudepleted (Eu/Nd<1) amphibole crystals suggest that some are 

not comagmatic (fig. 19). The three most extreme patterns 

(from samples SH304-2A, SH305-1, SH311-1B) are all lowAl amphiboles (7.4–8.9 percent Al2O3) with patchy zoning 

and variable resorption (for example, fig. 4A). They must be 

derived from a strongly Eu-depleted melt that has undergone 

significant Eu3+-enriched plagioclase crystallization in an 

oxidizing environment. Such distinctive REE patterns suggest 

that these low-Al amphiboles are xenocrysts. New trace-element data for gabbroic xenoliths, in which amphibole is a late 

cumulus phase, would help to test this hypothesis.

There are no textural distinctions among the group of 

slightly LREE-enriched, low- to medium-Al amphiboles with 

Eu/Nd <1 (fig. 18B). The slight REE enrichment reflects a 

normal magmatic fractionation trend from one or more dacite 

magmas. Low-Al amphiboles with slight Eu depletion may 

have been derived from earlier generations of solidified dacite accumulated at the top of the Mount St. Helens magma 

reservoir or by shallow low-temperature crystallization 

within the host magma. For the medium- to high-Al amphiboles, the most prevalent of our 2004–6 samples, LREE 

concentrations are consistently low, and Eu is not depleted 

relative to Nd.

Amphibole Stoichiometry and 

Nomenclature

The amphibole nomenclature used here, as recommended 

by the International Mineralogical Association, is that of 

Leake and others (1997). Nomenclatorial guidelines are based 

upon occupancies of the A, B, C, and T sites within the amphibole mineral structural formula,

A[Na+K] BCa2-x [x=Mg, Fe2+, Mn, Mg] CM1,M2,M3 Mg5-y

[y=Fe2+,Mn, VIAl, Ti, Fe3+] TSi8-z [z=IVAl] O22 OH2-n [n=F, Cl]).

Herein, the Leake and others (1997) nomenclature was 

applied to the amphibole stoichiometry determined using the 

Holland and Blundy (1994) method. This method was chosen 

because it allows for each amphibole to be normalized to the 

best stoichiometric constraint for ferric iron determination 

and gives the most reasonable amphibole formula for calculating pressure using the Anderson and Smith (1995) Alin-hornblende barometer. Leake and others (1997) propose 

a method for calculating amphibole formula from electron 

microprobe analysis that is similar to that of Holland and 

Blundy (1994). Amphibole classification using the two methods typically results in the same name except for species that 

are sensitive to ferric iron determinations, such as pargasite 

and magnesiohastingsite.

The complete 2004–6 amphibole data set, plotted using 

the classification diagram of Leake and others (1997), is 

shown in figure 20. High-silica, low-Al amphiboles in dacite 

and gabbroic xenoliths are magnesiohornblende, but they also 

include some alkaline, edenite compositions. Medium- to 

high-Al ranges of the data set fall in tschermakite to magnesiohastingsite fields. The medium-Al group, consisting 

mostly of tschermakite, gives way to magnesiohastingsite with 

increasing alumina and alkali contents, extending toward rare 

pargasite compositions. Aside from mineral classification, the 

main purpose of calculating amphibole stoichiometry is to 

evaluate crystallographically controlled cationic variations that 

are sensitive to pressure and temperature.

As is widely documented in studies of clinoamphibole 

thermobarometry of igneous and metamorphic rocks, the 

extent of Tschermak’s molecule substitution is a function 

of relative pressure and temperature of equilibration (Helz, 

1982; Hammarstrom and Zen, 1986; Anderson and Smith, 

1995; Bachmann and Dungan, 2002). At higher pressures and 

temperatures, the amount of Al in octahedral (C) sites (VIAl) 

increases together with the amount in tetrahedral Al (IVAl) 

sites. Most of this coupled alumina substitution occurs at 

the expense of divalent cations (for example, Mg2+ and Fe2+) 

in the octahedral C-site and silica in tetrahedral sites in a 

pressure-sensitive Al-Tschermak exchange (C(Mg,Fe) + TSi ↔ IVAl+VIAl). At higher temperatures, increased concentrations 

of alkali cations (Na+ and K+) and Ti4+ are accommodated with 

additional IVAl in the amphibole lattice and form edenite (Avacancy + TSi ↔ A(Na+K) + IVAl) and Ti-Tschermak (BMn+2TSi 

↔ CTi+2IVAl) substitutions. There is a broad, continuous range 

of pressure-sensitive Al-Tschermak component substitution 

exhibited by the complete array of amphibole compositions 

in 2004–6 Mount St. Helens dacite (fig. 21A). In addition, a 

broad range of temperature-sensitive Ti-Tschermak and edenite cationic exchanges also are apparent (figs. 21B–D).

The amphibole calculation optimizes ferric iron (Fe3+) in 

the stoichiometric scheme and is a means of “ironing out” the 

secondary deviations in total iron and magnesium with alumina (figs. 13B, 13C). The exchange of Fe2+ and Mg2+ in the 

C site is consistently linear in high- and medium-Al amphiboles (fig. 22). At higher alumina contents and consistent 

with Al-Tschermak substitution, the absolute concentrations 

of Mg and Fe2+ are lower along lines of constant Mg number 

(Mg#, 100*(Mg/(Mg+ Fe2+)). The Mg# variation in low-Al 

amphibole is more erratic. The Mg# increases with increasing 

proportion of Fe3+ (fig. 23A), suggesting that the Mg# trends 

are at least partly affected by differences in the oxidation state 

of host magmas. The variations of Ti and Mn with Mg# reveal 

distinctly different trends of low-Al magnesiohornblendes and 

the prevalent medium- to high-Al tschermakites and magnesiohastingsites (figs. 23B, 23C). Inverse trends of Ti versus 

Mg# (or Fe3+) exhibited by the range of common amphibole 

compositions (±1σ population trend in fig. 23) are consistent 

with growth at variable temperature and fO2

 under buffered 

conditions, as suggested by temperature-fO2 ranges calculated 

for Fe-Ti oxide pairs in 2004–6 dacite (Pallister and others, 

this volume, chap. 30, fig. 13).

Within the dataset there is no clear correlation between 

the Mg# and the relative partitioning of VIAl and IVAl (Al-
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 743

SH311-1

5

15

25

35

45

La Ce Pr Nd Sm Eu Gd Dy Er Yb

SH306-1A

0

5

10

15

20

SH305-1

5

15

25

35

45

SH304-2A

0

10

20

30

40

50

60

70

0

5

10

15

20

25

SH321-1

0

5

10

15

20

0

5

10

15

20

0

5

10

15

20

SH324-3

SH319-1

SH315-1

RARE-EARTH-ELEMENT CONCENTRATION, IN PARTS PER MILLION (NORMALIZED TO PRIMITIVE MANTLE)

La Ce Pr Nd Sm Eu Gd Dy Er Yb

n=9

n=8 n=9

n=8 n=7

n=9

n=10 n=9

ELEMENT ELEMENT

Figure 17. Rare-earth-element (REE) concentrations of amphibole phenocryst cores in eight consecutive 

Mount St. Helens 2004–6 dacite lava samples (normalized to “primitive mantle” of Sun and McDonough, 1989). 

Spidergrams show middle REE enrichment patterns that are typical of clinoamphibole and distinguished by 

varying degrees of REE enrichment and Eu depletion. Number of analyzed grains indicated (for example, n=8).
744 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

2

4

6

8

10

12

14

16

18

0.2

0

0.4

0.6

0.8

1.2

1.0

1.4

6 8 10 12 14 16

Al2O3, IN WEIGHT PERCENT

La, IN PARTS PER MILLLION Eu/Nd RATIO

A

B

SH304-2A

SH305-1

SH306-A

SH311-1B

SH315-1

SH324-3

SH323-2

SH321-1

SH319-1

0

10

20

30

40

50

60

70

La Ce Pr Nd Sm Eu Gd Dy Er Yb

RARE-EARTH-ELEMENT CONCENTRATION,

IN PARTS PER MILLION

(NORMALIZED TO PRIMITIVE MANTLE)

SH306-1A 1-5

SH315-1 2

SH319-1 5-1

SH304-2A 4-3 SH311-1B 3-5

SH305-1 4-3

SH311-1B 3-2

SH324-3 2-2

SH321-1 (1-1)

SH305-1 4-2

SH321-1(1-3)

SH324-3 1-3

ELEMENT

Figure 18. Rare-earth 

element (REE) versus 

alumina among amphibole 

phenocryst cores in 

consecutive Mount St. 

Helens 2004–6 dacite 

lava samples, showing 

trends reflective of overall 

REE enrichment and Eu 

depletion with decreasing 

alumina in medium- to 

low-Al amphiboles. A, La 

versus alumina. B, Eu/

Nd versus alumina; Eu/Nd 

values <1 are Eu depleted.

Figure 19. Rareearth-element (REE) 

concentrations in 

Eu-depleted 2004–6 

amphibole phenocrysts 

(normalized to 

“primitive mantle” of 

Sun and McDonough, 

1989). Samples with 

the three most extreme 

REE patterns, likely 

xenocrysts, have low Al 

(7.4–8.9 percent Al2O3), 

patchy zoning, and 

resorbed edges.
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 745

0.0

0.5

Magnesiohastingsite

Hastingsite Ferro-edenite

Edenite

0.5

Magnesiohornblende

Ferrotschermakite Hornblende

Tschermakite

0.0

6.0 6.5 7.0

6.5 6.0

Ferro-edenite

Edenite Pargasite

5.5

CaB≥1.50; (Na + K)A>0.5; Al< Fe V I 3+

CaB≥1.50; (Na + K)A<0.5

CaB≥1.50; (Na + K)A≥0.5; Al> Fe V I 3+

Si IN FORMULA

Si IN FORMULA

Mg/(Mg+Fe2+) IN C SITE

Figure 20. Mount St. Helens 2004–6 

amphibole phenocryst compositions plotted 

on classification diagrams of Leake and 

others (1997). Symbols used in each of the 

compositionally defined fields carry through to 

figures 21, 23, and 25.

Tschermak exchange; fig. 23D), indicating that Mg-Fe 

exchange is independent of relative pressure of crystallization. Also, no consistent differences are observed between 

core and rim Mg# and VIAl or IVAl. These observations and 

interpretations are in contrast to the inverse relation of Mg# 

and IVAl documented within individual 2004–6 amphibole 

crystals by Rutherford and Devine (this volume, chap. 31). 

The covariance of Mn, Ti, and Fe2+ in the medium- to high-Al 

amphibole suite likely reflects differences in the abundance 

and composition of coexisting Fe-Ti oxides in the crystallizing assemblage, which are affected by changes in temperature 

and the oxidation state over a range of pressures reflected by 

Al-Tschermak substitution.

Amphibole Thermobarometry

Rutherford and Devine (this volume, chap. 31) conducted 

phase-equilibria experiments on 2004 dacite that establish 

lower PH2O and upper T limits of amphibole+plagioclase+

orthopyroxene+Fe-Ti-oxide saturation. Their results reproduce the observed phenocrysts of the dacite between pressures 

of 130 and 300 MPa (~5–12 km depth) at 850°C to 920°C, 

respectively (Rutherford and Devine, this volume, chap. 31, 

fig. 10). They observed increased Tschermak component of 

Mount St. Helens amphibole equilibrated at 200–300 MPa 

and 900°C in comparison with that at or below 200 MPa at 

~860–850°C. These data suggest that the Tschermak and 

edenite substitutions, exhibited by the predominant 2004–6 

tschermakite and magnesiohastingsite compositions, ought 

to constrain the relative P–T path of dacite magmas that have 

contributed to the phenocryst mix observed in new lava.

The Al-in-hornblende thermobarometer of Anderson 

and Smith (1995) is used here to estimate relative equilibration pressures over the ~850–900°C temperature range 

established by iron-titanium oxide thermometry and experimental petrology of new dacite lava (Pallister and others, this 

volume, chap. 30; Rutherford and Devine, this volume, chap. 

31). As expressed by Anderson and Smith (1995, p. 554) and 

debated by others (for example, Blundy and Holland, 1990; 

Rutherford and Johnson, 1992; Hammarstrom and Zen, 

1992), this algorithm may not yield accurate results at temperatures in excess of experimental calibration (800°C) and 

without all of the Tschermak-buffering phases in the barometric assemblage. We assume that most amphiboles in the 

2004–6 dacite were crystallized from similar dacite magmas 

at temperature <920°C and that all were cosaturated with 

fluid, plagioclase, orthopyroxene, and oxide. The covariance 

of alumina and silica in these phases is not fixed for every 

amphibole that crystallizes, but the Tschermak molecular 

substitution should be constrained by this same five-phase 

assemblage throughout the pressure range in which the 

amphibole crystallized, thus providing a relative indication 

of crystallization pressures and temperatures.

Calculated pressures for amphiboles equilibrated under 

known P and T conditions help validate application of Al-inhornblende thermobarometry to Mount St. Helens dacite (fig. 

24). The Anderson and Smith (1995) algorithm is accurate 

for the 200- to 800-MPa range of experimental amphiboles at 
746 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0.0

0.1

0.2

0.3

0.4

0.5

0.6

0.1

0.2

0.3

0.4

0.5

0.01

0.02

0.03

0.04

0.05

0.2

0.3

0.4

0.5

0.6

0.7

0.8

1.0 1.2 1.4 1.6 1.8 2.0

IVAl, IN T SITE

Na+K, IN A SITE Ti, IN C SITE Mn, IN C SITE VIAl, IN C SITE

A

C

D

B

0.6

1.0

1.4

1.8

2.2

2.0 2.2 2.4 2.6 2.8 3.0 3.2 3.4 3.6

Mg#=70

Mg#=80

Mg#=60

Medium-Al Low-Al

EXPLANATION

High-Al

Fe2+ IN C SITE

Mg2+ IN C SITE

Figure 21. Site-specific cation concentrations versus 

tetrahedrally coordinated Al (IVAl) in Mount St. Helens 2004–6 

amphibole phenocrysts grouped according to amphibole type, 

with symbols as defined in figure 20. Magnesiohornblendes 

of 1986 gabbroic xenoliths, shown with black squares, are 

included for comparison. Lines shown represent best-fit 

linear regressions. A, Increasing octahedral alumina (VIAl) 

concentrations roughly coincide with increasing tetrahedral 

alumina (IVAl), reflecting a range of pressure-sensitive 

Al-Tschermak substitution. B and C, Antithetic variations of 

Mn and Ti with increasing IVAl demonstrate a continuous range 

of temperature-sensitive Ti-Tschermak substitution. D, Total 

alkali concentrations (Na+K in A site) increase with increasing 

IVAl in an edenite exchange, also indicative of increasing 

temperature of crystallization.

Figure 22. Fe2+ versus Mg2+ variation diagram for low-Al, 

medium-Al, and high-Al groups of Mount St. Helens 2004–6 

amphiboles. At higher alumina contents, concentrations of 

Mg and Fe2+ (in C sites) are lower along lines of constant 

Mg# (=Mg/(Mg+Fe2+), shown as dashed black lines. A 

well-balanced Fe-Mg C-site exchange is observed among 

medium-Al to high-Al amphiboles. The dashed red line shows 

best-fit linear regression of predominant 2004–6 amphibole 

population (defined as ±1σ from average composition).
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 747

0.0

0.1

0.2

0.3

0.4

0.5

VIAl, IN C SITE Ti, IN C SITE

0.1

0.2

0.3

0.4

Mn, IN B SITE

0.01

0.02

0.03

0.04

Mg2+/(Mg2++Fe2+) IN C SITE

Fe3+/(Fe3++Fe2+), IN C SITE

0.1

0.2

0.3

0.4

0.5

0.6

0.7

0.5 0.6 0.7 0.8 0.9

A

D

B

C

Figure 23. Site-specific cation concentrations versus 

Mg# in C site (Mg2+/(Mg2++Fe2+) in Mount St. Helens 

2004–6 amphibole phenocrysts, grouped according to 

amphibole type, with symbols as defined in figure 20. 

Dashed lines represent best-fit linear regressions of 

predominant 2004–6 amphibole population (defined as ±1σ

from average). Magnesiohornblendes in 1986 gabbroic 

xenoliths, shown as black squares, are presented for 

comparison. A, Mg# increases with increasing proportion 

of Fe3+ relative to total iron (Fe3+ + Fe2+), suggesting that 

Mg# is largely affected by magmatic oxidation state. B

and C, Covariance of Mn and Ti with Mg# in predominant 

amphibole population indicates Mg# varies independently 

of temperature-sensitive Ti-Tschermak exchange (figs. 

21B, 21C). D, Octahedral alumina (VIAl) concentrations 

show no systematic correlation with Mg#, indicating that 

Fe2+-Mg exchange is independent of pressure-sensitive 

Al-Tschermak substitution.

J&R 750°C

R&D 850°C

R&D 860°C

R&D 900°C

100

300

500

900

700

100 300 500 700 900

MEASURED PRESSURE, IN MEGAPASCALS

CALCULATED PRESSURE, IN MEGAPASCALS

Figure 24. Plot of experimentally measured pressures 

versus pressures calculated by Al-in-hornblende barometer 

of Anderson and Smith (1995), using temperatures and 

amphibole compositions from experimental runs. J&R, 

Johnson and Rutherford (1989) experiments at 750°C on 

dacite from Fish Canyon Tuff (64.4 percent SiO2) and Round 

Valley, Idaho (65.2 percent SiO2), in which amphibole coexists 

with melt, fluid, quartz, biotite, plagioclase, sphene, and Fe-Ti 

oxides. R&D, Rutherford and Devine (this volume, chap. 31) 

experiments on 2004 Mount St. Helens dacite at 850°C, 860°C, 

and 900°C in which amphibole coexists with melt, fluid, 

plagioclase, orthopyroxene and Fe-Ti oxides.
748 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

740°C in three calc-alkaline magma compositions (64.4–65.2 

percent SiO2, Johnson and Rutherford, 1989) for which it was 

calibrated. In those experiments, the coexisting phase assemblage (melt, fluid, quartz, biotite, plagioclase, alkali-feldspar, 

sphene, and Fe-Ti oxides) provides a complete Tschermakbuffering capacity (Anderson and Smith, 1995). In experiments 

on 2004 Mount St. Helens dacite at pressures of 100–200 MPa 

and temperatures of 850–860°C and 900°C (Rutherford and 

Devine, this volume, chap. 31), amphibole coexists with a 

fluid+melt+plagiclase+orthopyroxene+oxide assemblage, 

similar to that of newly erupted lava. Overall, barometric calculations for amphibole equilibrated in the silica undersaturated 

dacite (analyses from table 6 of Rutherford and Devine, this volume, chap. 31) yield results consistent with the <200 MPa run 

conditions. Most of the estimates are low by more than 50 MPa, 

but there is closer agreement for the highest-Al amphiboles in 

runs at both 900°C and 850°C.

Barometric calculations for the predominant 2004–6 

amphibole population (±1σ range, fig. 25) define a depth range 

for crystallization from 13.5 to ~4 km depth (350 MPa to 100 

MPa) at temperatures between 850°C and 900°C. The overall 

average of phenocryst compositions yields barometry results 

of 220 MPa at 850°C. As restricted by the minimum pressure 

of amphibole stability (100 MPa), our calculations show that 

4

0

8

12

16

20

24

28

100

300

500

700

900 850 800 750

APPROXIMATE DEPTH,

IN KILOMETERS BELOW SURFACE

TEMPERATURE USED FOR BAROMETRIC CALCULATIONS,

 IN DEGREES CELSIUS

PRESSURE, IN MEGAPASCALS

AMPHIBOLE OUT

+1σ

–1σ

X

the bulk of amphibole crystallization occurred at temperatures 

less than ~875°C. At ~900°C, only high-Al amphiboles yield 

tenable calculated pressures, ranging from 100 to ~250 MPa, 

and perhaps to 300 MPa. Judging from the apparent error in 

applying Al-barometry to experimentally produced amphiboles at <200 MPa (fig. 24), it may be that actual equilibration 

pressures are higher by about 50–100 MPa. At 850°C this correction would increase the maximum depth range of average 

amphibole crystallization from ~8.5 km (220 MPa) to ~10.5 

or perhaps to 12.5 km (270–320 MPa) and would increase the 

maximum depth range of the more aluminous tschermakite 

and magnesiohastingsite to >15 km (400 MPa).

The relative calculated pressures for the majority of 

amphibole compositions suggests two P-T scenarios for the 

2004–6 magma and for prior generations of hostlike, precursory magma that are preserved in the crystal-laden mixture: 

(1) high-Al amphiboles formed at relatively low temperature 

and great depth (>15 km at <850°C), and the subsequent bulk 

of amphibole crystallization occurred during nearly isothermal 

ascent, or (2) crystallization during cooling between 900°C 

and 800°C at shallower depths within the magma reservoir 

(from 4 km to ~8 or possibly ~10 km), or a combination of (1) 

and (2). In either case, the range of calculated thermobarometric conditions reflected by tschermakite to magnesiohastingFigure 25. Plot showing results of Al-in-hornblende barometry applied to 2004–6 amphiboles at 750°C, 800°C, 

850°C, and 900°C. Symbols correspond to amphibole types, as shown in figure 20. Inferred pressure (depth) versus 

temperature window for the average (X) and predominant amphibole population (±1σ range) is indicated.
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 749

site compositions in new lava reflects a well-mixed population 

of crystals formed in similar dacite magma over a likely depth 

range of 4–13.5 km, or possibly as deep as 15 km.

Our calculations suggest that low-Al magnesiohornblendes require low crystallization temperatures (750–850°C) to 

yield minimum pressures of amphibole stability (>100 MPa, 

fig. 25). Their presence in the mix suggests entrainment of 

near-solidus, and perhaps older, dacite mush (along with 

gabbroic material) in hotter 2004–6 magma. Those low-Al 

amphiboles that are not xenocrysts may have crystallized at 

low pressures and temperatures from host dacite melt. The 

highest Al-tschermakite, magnesiohastingsite, and pargasite 

could be derived from a mafic source (magma or wall rock) or 

could have crystallized from dacite under high P-T conditions 

above plagioclase saturation. In either case, magma crystallizing the end-member, high-Al amphiboles may not be close to 

quartz-saturation or to five-phase saturation conditions. In this 

case, barometry calculations would yield untenably low pressures over the temperature range of interest.

Amphibole Decompression Reactions, 

Magma Ascent Dynamics, and Conduit 

Geometry

The nature and extent of amphibole dissolution and reaction associated with variable decompression paths of dacite 

magma are documented in the recently published experimental 

study of Browne and Gardner (2006). Data from their series of 

isothermal decompression runs, performed on similar compositions and initially equilibrated under conditions analogous 

to those of the Mount St. Helens 2004–6 magma (840°C and 

150 MPa), help to constrain paths of magma ascent. Reaction 

rims similar to the typical 5-μm rims on subhedral, resorbed, 

or fragmented amphibole crystals in the 2004–6 Mount St. 

Helens dacite were reproduced by isothermal decompression 

from 60 to <30 MPa (~2.5- to <1-km depth) over an interval 

of 3 to 4 days (compare textures in our fig. 7 with those in fig. 

4 of Browne and Gardner, 2006). Such 5-μm rims must have 

formed by the time magma reached the uppermost conduit 

because amphibole ceases to react at depths less than ~0.6 km 

in a highly viscous, nearly solidified, and completely degassed 

host (Browne and Gardner, 2006).

Textures of resorbed amphibole crystals seen in the 

2004–6 Mount St. Helens dacite (fig. 7) could be produced if 

the host magma stalled for ~2 days at ~3.5 to 2.5-km depth (90 

to 60 MPa) during ascent (Browne and Gardner, 2006). Under 

such conditions, amphibole dissolves without forming pseudomorph reaction rims, which develop subsequent to resorption 

during further ascent. The disappearance of amphibole, as 

documented by resorption textures and decreased modal abundance in experimental products, is accompanied by nucleation 

and growth of orthopyroxene and plagioclase (±Fe-Ti oxide) 

microphenocrysts in the vicinity of dissolved amphibole. 

Such microphenocrysts are likely to have contributed to the 

microcrystalline matrix, which further develops during ascent 

(Blundy and Cashman, 2001; Cashman and others, this volume, chap. 19).

Amphiboles with <3-μm rims in early dome samples 

with a glassier matrix (figs. 9A, 9B; and Pallister and others, 

this volume, chap. 30) suggest faster ascent rates through the 

60–30-MPa pressure range than later dacite, consistent with 

the relatively high eruptive flux during the earliest phases of 

dome extrusion (Schilling and others, this volume, chap. 8; 

Reagan and others, this volume, chap. 37).

The sporadic occurrence of wider reaction rims (30–200 

μm) may be attributed to the haphazard entrainment of 

amphibole phenocrysts subjected to extensive reaction while 

occluded to walls of the lower conduit. Alternatively, some 

phenocrysts with wide reaction rims could be xenolithic fragments of 1980s dacite, which is known to have two reactionrim populations (10–12 and 30 μm; Rutherford and Hill, 

1993). Composite phenocrysts with resorbed and patchy cores 

surrounded by subhedral zoned overgrowths and amphiboleto-amphibole reactions probably reflect xenocryst assimilation 

at depth. It is evident from less commonly observed examples 

of resorbed medium-Al amphibole crystals with low-Al 

(and high fluorine) amphibole overgrowths (for example, in 

SH314-1 and DRS_3_9_4; fig. 11, table 2) that this style of 

amphibole dissolution and reaction has occurred in magma 

stored at depths near to and perhaps slightly above the limits 

of amphibole stability. Such cases would seem to reflect the 

disruption of a crystal mush by a hotter or chemically incompatible melt. Rutherford and Devine (this volume, chap. 31) 

suggest that fluorine has the effect of stabilizing amphibole at 

depths above that of H2O-only amphibole saturation. Thus, it 

is possible that late-stage fluorine enrichment could promote 

the occurrence of such amphibole-to-amphibole+melt reactions in the lower conduit.

The persistent 5-μm amphibole rims in all post-November 2004 spine lava reflect steady influx and upward transport 

of magma through the lower conduit, possibly on the order 

of 2–4 days, from 3.5–2.5-km depths to as shallow as ~1 km. 

Thus, an amphibole-based rate of magma ascent for this depth 

interval is between ~600 and ~1,200 m/day. As constrained 

by a first-year erupted volume of about 70×106

 m3 (Schilling 

and others, this volume, chap. 8), such rates infer a conduit 

radius between 10 and 5 m. In contrast, ground deformation 

and eruptive flux infer a significantly larger average conduit 

radius of ~35 m (down to 5 km depth) (Dzurisin and others, 

this volume, chap. 14).

A similar discrepancy existed between conduit dimensions inferred from amphibole-rim and volumetric-flux modeling of the 1980s dome-building eruption (Cashman, 1992; 

Pallister and others, 1992). For the current eruption, both interpretations might be valid, but for different depth ranges. Taken 

together, these analyses suggest that the conduit may have a 

wineglass shape. At ~3.5–1-km depth, the conduit might be a 

narrow pipe (or short dike) that accommodates relatively rapid 

transit of dacite magma, as recorded by the thin amphibole 
750 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

reaction rims. Above ~1 km, where matrix crystallization of 

magma occurs (Cashman and others, this volume, chap. 19; 

Pallister and others this volume, chap. 30) and amphibole 

reactions cease, a cooling, solidifying, and sluggishly moving 

magma is accommodated by a wider conduit (>60 m radius). 

Such a wineglass shape for the conduit also might help 

explain (1) the apparent localization of deformation to the 

crater and to a deep region whose centroid lies at depth of ~8 

km (Dzurisin and others, this volume, chap. 14; Lisowski and 

others, this volume, chap. 15) and (2) the lack of geophysical 

or petrologic evidence for magma storage between ~4–5-km 

and ~1-km depths.

The Amphibole Perspective: from 

Source to Surface

Our study of amphibole systematics shows the near-solid 

magma recently extruded from Mount St. Helens is a wellblended, crystal-rich dacite mush that originated from depths 

of at least ~4 km and ascended to within 1 km of the surface 

at rates in the range ~0.6–1.2 km/day. The compositional 

ranges of amphibole phenocrysts in the 2004–6 dacite reflect 

multiple generations of (primarily) dacite magma crystallization in different P-T environments. Amphibole morphology 

and texture reveal a dynamic magma-mixing regime in which 

a genetically disparate, amphibole phenocryst assemblage 

is variably dissolved, overgrown, broken, and abraded. The 

same ranges of amphibole compositions and textures are 

found in all thin sections of dome material. An obvious 

implication of the well-dispersed amphibole variability is that 

the monotonous bulk chemistry of new lava is the result of 

fine-scale (less than a cubic centimeter) blending of multiple, 

crystal-laden dacite magmas.

Petrologic constraints, including complex zoning in some 

amphibole phenocrysts, have spawned models in which this 

latest generation of Mount St. Helens lava evolved physically 

and chemically to its preeruptive condition by convective 

recycling of crystal-laden dacite magma through a contiguous 

P-T regime (such as that depicted in the simplified magmareservoir diagrams of Pallister and others, this volume, chap. 

30, fig. 22; Rutherford and Devine, this volume, chap. 31, fig. 

13). Our thermobarometric calculations place amphibole crystallization in a P-T window between 100 and ~350–400 MPa 

(~4- and 13.5–15-km depth) and ~900°C and ~800°C. This 

interpretation of the P-T regime for eruption-related magmatism is within that constrained by experimentally determined 

phase relations (Rutherford and Devine, this volume, chap 31) 

and is consistent with the seismologically defined 4–12-kmdeep magma source (Moran, 1994).

The broad P-T regime reflected in the variety of phenocrysts also could be an artifact of repeated injection of 

different crystal-laden, dacite magma batches under more 

restricted P-T conditions in the uppermost zone of the magma 

reservoir. Amphibole petrology and several other lines of evidence suggest that much of the blending of the 2004–6 dacite 

mush is achieved between ~8- and 4-km depth. The average 

amphibole in 2004–6 magma is interpreted to have crystallized 

at 850°C and 220 MPa (fig. 25). Coincidentally, those same 

source conditions were interpreted for the May 18, 1980, pumice (Rutherford and others, 1985). Considering that the range 

of amphibole compositions of 2004–6 dacite is comparable 

to that of amphibole in 1980–86 lava, we suggest a consistent 

staging area at and above ~8-km depth for both eruptions. 

Thermobarometric calculations indicate that cyclic variations within the range of the predominant 2004–6 amphibole 

compositions could result from heating and cooling within this 

upper part of a magma reservoir (fig. 25). Compositional zoning in the outer parts of plagioclase in the 2004–6 dacite are 

likewise consistent with injections of hotter, porphyritic dacite 

magma into crystal-laden dacite mush (Streck and others, this 

volume, chap. 34).

Isotopic studies of plagioclase in 2004–6 dacite (Cooper 

and Donnelly, this volume, chap. 36; Kent and others, this 

volume, chap. 35) suggest that the majority of phenocrysts are 

“zero-age” and cognate to the host dacite. Of the remaining 

phenocrysts, a small proportion are Tertiary xenocrysts, but 

20–40 percent of analyzed plagioclase crystals are considered 

antecrysts, derived from prior generations of Mount St. Helens 

dacite (see Bacon and Lowenstern, 2005, for further discussion of the term, antecryst). Our data indicate that amphibole 

phenocrysts have a similarly diverse heritage. Although 

2004–6 dacite antecrysts may be an inevitable product of longterm mixing, it is also likely that some of them, as well as 

xenocrysts, are assimilated from variably aged roof rocks and 

surrounding crustal materials by intrusive processes.

The disequilibrium reflected both in the textural variety 

of amphiboles and in the wide variation of their outermost rim 

compositions is a clear indication of a dynamic, preeruptive, 

magma-mixing environment for the 2004–6 dacite. Uniform 

reaction rims on all subhedral, rounded, resorbed, and fragmented amphibole phenocrysts of all compositional affinities indicate that this crystal assemblage was being broken, 

abraded, and dissolved in the magma immediately before and 

during early stages of ascent. An ascent path that accommodates these features involves differential flow rates through 

regions of the lowermost conduit and uppermost magma 

reservoir. In this scenario, supported by experimental observations of Browne and Gardner (2006), some magma must 

stagnate long enough at 90–60 MPa (~3.5–2.5-km depth) for 

amphiboles to be resorbed, but other magma does not. This 

situation seems plausible for a magma ascending through 

deep conduit roots that must extend into a mushy reservoir 

cupola. The crystal-laden magma is viscous enough to fragment and entrain additional amphibole phenocrysts during 

ascent through crystal mushes in the upper reservoir and lower 

conduit at ~5–3.5 km. Small reaction rims on the well-mixed 

amphibole assemblage result from rapid decompression of 

host-dacite magma during ascent through a small conduit 

(≤10-m radius) extending upward from ~3.5 km to within ~1 

km of the crater floor. From ~1 km to the surface, the final 
32. Chemistry, Mineralogy, and Petrology of Amphibole in Mount St. Helens 2004–2006 Dacite 751

preeruptive condition of amphibole is frozen in as the magma 

quenches into a nearly solid plug of porphyritic lava. During 

and after this magma-to-lava transition, a relatively sluggish 

plug ascent is accommodated by an enlarged (≥60-m radius) 

uppermost conduit, in which the fresh dacite lava is sheared 

and gouged during forceful expulsion as lava spines.

Acknowledgments

We thank Jeff Linscott of JL Aviation, whose adept piloting skills enabled heli-dredge collection of lava samples, and 

Bobbie Myers of the U.S. Geological Survey’s Cascades Volcano Observatory (USGS–CVO), who helped to ensure the 

safety of our dome-sampling forays. Assistance with sample 

processing and archiving by Winston Stokes (USGS–CVO) 

and dedicated volunteers in the CVO Petrology Laboratory 

(Siobhan McConnell, Trystan Herriott, Taryn Lopez, and Alison Eckberg) is gratefully acknowledged. Isabelle Brownfield 

(USGS, Denver) provided many of the SEM images used 

in the course of this investigation. Thoughtful reviews of an 

early version of this chapter by Roz Helz (USGS, emeritus) 

and Mac Rutherford (Brown University) resulted in significant improvements.

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754 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Analytical Methods

Microbeam analyses and imagery used in this study were accomplished at three different laboratories. Most data were 

obtained using the USGS Denver JEOL 8900 electron microprobe and a JEOL 5800-LV scanning electron microscope (SEM). 

Nine samples were analyzed and imaged by M. Rowe at Oregon State University on a Cameca SX-100 electron microprobe. 

At both laboratories, amphibole grains were analyzed using 20-nA beam current and a 15 keV accelerating voltage. For all but 

the smallest of grains, which required a fully focused beam, a 10-μm beam diameter was used for crystal interiors and near-rim 

analyses. Count times for major elements varied between 30 and 10 s on peak (15 to 5 s background count times). For F and Cl, 

count times on peak were increased to 60 and 40 s, respectively. USNM natural silicate mineral standards were used for calibration. Kakanui hornblende (Jarosewich and others, 1980) was repeatedly analyzed to monitor accuracy and precision during and 

among runs.

Additional analyses of SH304-2A (early dacite) and SH304-2G (glassy lithic fragment) amphibole crystals were done by 

C. Mandeville at the American Museum of Natural History (AMNH) using a Cameca SX-100 electron microprobe utilizing a 

15-keV accelerating voltage, 10-nA beam current for major elements, 15-μm beam diameter, 30 s count time on peak and 15 s 

on background. F, Cl, and Cr analyses of amphiboles by electron microprobe at AMNH were done during the same analytical 

session with 15-keV accelerating voltage, 40-nA beam current, and 140 s count time on peak and 70 s on background. Precision 

and accuracy of major-element analyses were checked by repeated analyses of Kakanui hornblende. Precision and accuracy of 

F, Cl, and Cr analyses at AMNH were monitored by repeated analyses of Macusani peraluminous rhyolite glass (1.3 percent F, 

450 ppm Cl; Pichavant and others, 1987), NMNH 164905 chromium-augite (0.85 percent Cr; Jarosewich and others, 1980), and 

NIST SRM 610 glass (415±49.5 ppm Cr; Pearce and others, 1997).

All amphibole analyses used in this investigation fit the mineralogical criteria established for clinoamphiboles by Leake 

and others (1997). See text for discussion of the stoichiometric calculation used to assess cation site occupancy and proper 

nomenclature.

Trace-element concentrations of amphiboles were analyzed by laser ablation inductively coupled plasma mass spectrometry 

(LA ICP-MS) in the W.M. Keck Collaboratory for Plasma Spectrometry, Oregon State University, and collected from the same 

phenocryst areas that were analyzed and imaged by the Cameca SX-100 electron microprobe. Amphiboles were ablated with a 

NuWave 213 nm Nd:YAG laser using a 40-μm stationary spot and 20-Hz pulse rate. Ablation time and data acquisition for each 

analysis was 30 s, with 45 s of washout time before and after ablation. Measured counts were normalized to 43Ca. Trace-element 

concentrations were then calculated relative to the BCR-2G glass standard. BHVO-2G glass standard was ablated under identical conditions as the amphiboles as a secondary standard. Precision, estimated by repeated analysis of BHVO-2G, is within 10 

percent of reported values for all trace elements (Kent and others, 2004).

Appendix 2. Major- and Trace-Element Compositions of Amphibole Phenocrysts 

in Mount St. Helens 2004–2006 and 1980–1986 Dacite Lava

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

The Mount St. Helens amphibole phenocryst chemistry described in this chapter is tabulated in six worksheets of a 

Microsoft Excel file. The six worksheets are organized as follows:

Worksheet 1. Electron microprobe analyses of amphibole phenocrysts in Mount St. Helens 2004–6 dacite.

Worksheet 2. Electron microprobe analyses of amphibole phenocrysts in Mount St. Helens 2004 glassy lithic fragments.

Worksheet 3. Electron microprobe analyses of amphibole phenocrysts in Mount St. Helens 2004–6 lithic inclusions.

Worksheet 4. Electron microprobe analyses of amphibole phenocryst cores in Mount St. Helens 1986 gabbro xenoliths.

Worksheet 5. Electron microprobe analyses of amphibole phenocryst cores in Mount St. Helens 1980–86 dacite.

Worksheet 6. LA-ICPMS rare-earth-element analyses of amphibole phenocrysts in Mount St. Helens 2004–6 dacite.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 33

1

 Department of Earth Sciences, University of Bristol, Wills Memorial 

Building, Bristol BS8 1RJ, United Kingdom

2

 Department of Geological Sciences, 1272 University of Oregon, Eugene, 

OR 97403

3

 Department of Earth Sciences, Wills Memorial Building, Bristol BS8 1RJ, 

UK; now at Earth & Planetary Sciences, McGill University, 3450 University 

St., Montreal, Quebec, Canada H3A 2A7

Evolving Magma Storage Conditions Beneath Mount St. 

Helens Inferred from Chemical Variations in Melt Inclusions 

from the 1980–1986 and Current (2004–2006) Eruptions

By Jon Blundy1, Katharine V. Cashman2, and Kim Berlo3

Abstract

Major element, trace element, and volatile concentrations in 187 glassy melt inclusions and 25 groundmass glasses 

from the 1980–86 eruption of Mount St. Helens are presented, 

together with 103 analyses of touching Fe-Ti oxide pairs 

from the same samples. These data are used to evaluate the 

temporal evolution of the magmatic plumbing system beneath 

the volcano during 1980–86 and so provide a framework in 

which to interpret analyses of melt inclusions from the current 

(2004–2006) eruption.

Major and trace element concentrations of all melt inclusions lie at the high-SiO2 end of the data array defined by eruptive products of late Quaternary age from Mount St. Helens. 

For several major and trace elements, the glasses define a trend 

that is oblique to the whole-rock trend, indicating that different 

mineral assemblages were responsible for the two trends. The 

whole-rock trend can be ascribed to differentiation of hydrous 

basaltic parents in a deep-seated magma reservoir, probably at 

depths great enough to stabilize garnet. In contrast, the glass 

trends were generated by closed-system crystallization of the 

phenocryst and microlite mineral assemblages at low pressures.

The dissolved H2O content of the melt inclusions from 

1980–86, as measured by ion microprobe, ranges from 0 to 6.7 

wt. percent, with the highest values obtained from the plinian 

phase of May 18, 1980. Water contents decrease with increasing SiO2, consistent with decompression-driven crystallization. 

Preliminary data for dissolved CO2 in melt inclusions from 

the May 18 plinian phase and from August 7, 1980, indicate 

that XH2

O in the vapor phase was approximately constant 

at 0.80, irrespective of H2O content, suggestive of closedsystem degassing with a high bubble fraction or gas streaming 

through the subvolcanic system. Temperature and fO2 estimates 

for touching Fe-Ti oxides show evidence for heating during 

crystallization owing to release of latent heat. Consequently, 

magmas with the highest microlite crystallinities record the 

highest temperatures. Magmas also become progressively 

reduced during ascent and degassing, probably as a result of 

redox equilibria between exsolving S-bearing gases and magmas. The lowest temperature oxides have fO2

≈ NNO, similar 

to high-temperature fumarole gases from the volcano. The 

temperature and fO2 of the magma tapped by the plinian phase 

of May 18, 1980, are 870–875°C and NNO+0.8, respectively.

The dissolved volatile contents of the melt inclusions 

have been used to calculate sealing pressures; that is, the 

pressure at which chemical exchange between inclusion and 

matrix melt ceased. These are greatest for the May 18 plinian 

magma (120 to 320 MPa); lower pressures are recorded by 

samples of the preplinian cryptodome and by all post-May 18 

magmas. Magma crystallinity, calculated from melt-inclusion 

Rb contents, is negatively correlated with sealing pressure, 

consistent with decompression crystallization. Elevated 

contents of Li in melt inclusions from the cryptodome and 

post-May 18 samples are consistent with transfer of Li in a 

magmatic vapor phase from deeper parts of the magma system to magma stored at shallower levels. The Li enrichment 

attains its maximum extent at ~150 MPa, which is ascribed to 

separation of a single vapor phase into H2O-rich gas and dense 

Li-rich brine at the top of the magma column.

There are striking correlations between melt-inclusion 

chemistry and monitoring data for the 1980–86 eruption. Dissolved SO2 contents of melt inclusions from any given event, 

multiplied by the mass of magma erupted during that event, 
756 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

correlate with the measured flux of SO2 at the surface, suggesting that magma degassing and melt-inclusion sealing are 

closely related in time and space.

Textural and chemical evidence indicates that melt inclusions became effectively sealed (physically or kinetically) 

shortly before eruption. Thus by converting pressure to depth 

using a density model and edifice-loading algorithm for the 

volcano, changing depths of magma extraction with time can 

be tracked and compared to the seismic record. The plinian 

eruption of May 18, 1980, involved magma stored 5–11 km 

below sea level; this is inferred to be the subvolcanic magma 

chamber. The preceding eruptions, including the May 18, 

1980, blast, involved magma withdrawal from the cryptodome 

and conduit down to 5 km below sea level. Subsequent 1980 

eruptions tapped magma down to depths of ≤10 km below 

sea level. Tapping of magma stored deeper than 2 km below 

sea level stopped abruptly at the end of 1980, coincident 

with the onset of extensive shallow seismicity and a change 

from explosive to effusive eruption style from 1981 to 1986. 

Overall, the 1980–86 eruption is consistent with the evisceration of a thin, vertically extensive body of magma extending 

from 5 to at least 11 km below sea level and connected to the 

surface by a thin conduit. In the absence of sustained high 

magma-supply rates from depth, decompression crystallization of magma ascending through the system leads eventually 

to plugging of the conduit.

The current eruption of Mount St. Helens shares some 

similarities with the 1981–86 dome-building phase of the 

previous eruption, in that there is extensive shallow seismicity 

and extrusion of highly crystalline material in the form of a 

sequence of flows and spines. Melt inclusions from the current eruption have low H2O contents, consistent with magma 

extraction from shallow depths. Highly enriched Li in melt 

inclusions suggests that vapor transport of Li is a characteristic 

feature of Mount St. Helens. Melt inclusions from the current 

eruption have subtly different trace-element chemistry from 

all but one of the 1980–86 melt inclusions, with steeper rareearth-element (REE) patterns and low U, Th, and high-fieldstrength elements (HFSE), indicating addition of a new melt 

component to the magma system. It is anticipated that increasing involvement of the new melt component will be evident as 

the current eruption proceeds.

Introduction

Quenched melt inclusions in phenocrysts from volcanic 

rocks can provide information on preeruptive conditions within 

the subvolcanic magma body. Concentrations of H2O and other 

volatile species can be used to infer preeruptive storage depths, 

whereas concentrations of major and trace elements can be 

used to elucidate the arrival of new magma batches and, in 

the case of highly incompatible elements, the crystallinity of 

the magma at the time of inclusion entrapment. Of particular 

value are plagioclase-hosted melt inclusions, because of their 

ubiquity in calc-alkaline magmatic rocks and because of their 

ability to record a wide range of magmatic conditions through 

partial reequilibration with the matrix melt during magma 

ascent and crystallization (Blundy and Cashman, 2005). By 

allying melt inclusion data to determinations of temperature 

and oxygen fugacity (fO2) from coexisting iron-titanium oxides, 

it is possible to provide a detailed image of evolving subvolcanic magmatic conditions, which can in turn be linked to 

monitoring data such as volatile flux and seismicity.

We have previously published data on H2O and major 

elements in melt inclusions from the 1980–86 eruption of 

Mount St. Helens. Here we augment the published dataset 

with additional data, including previously unpublished traceelement data. The total dataset for the 1980–86 eruption now 

comprises 212 glasses, including 172 melt inclusions hosted in 

plagioclase, 8 in amphibole, 4 in orthopyroxene, 3 in clinopyroxene, and 25 groundmass glasses. Major elements, H2O, and 

light trace elements (Li-Ti) have been determined for all of 

these glasses; heavy trace elements (Ti-U) have been determined for 74 of them. A preliminary study of dissolved CO2

was carried out on two samples from the May 18 and August 

7, 1980, eruptive episodes. In order to compare the conditions 

of magma storage during the 1980–86 eruption with those of 

the current (2004–2006) eruption, we have analyzed 11 melt 

inclusions in three samples from the current eruption for the 

same suite of elements. We also present new data on touching 

Fe-Ti oxide pairs for 14 samples spanning the entire 1980–86 

eruption, for comparison with data from the new eruption (Pallister and others, this volume, chap. 30).

Materials and Methods

A full list and brief description of the 32 samples analyzed, together with their origin and any previous publications that describe them, are presented in table 1. Most of the 

samples were prepared as grain mounts of plagioclase and 

mafic minerals; in some cases thin sections were also used.

Melt inclusions are widespread in plagioclase phenocrysts 

from all samples studied. Most inclusions are glassy without 

evidence of daughter crystals. About 20 percent of all inclusions 

analyzed contain small vapor bubbles exposed at the surface 

of the thin section or grain mount. Thin rims or embayments 

of plagioclase around the walls of most plagioclase-hosted 

inclusions testify to some crystallization after the inclusion was 

first formed (Blundy and others, 2006). Melt inclusions are 

less common in mafic phenocrysts (amphibole, clinopyroxene, 

orthopyroxene). These inclusions typically lack clear evidence 

of host-crystal precipitation on their walls. Some of the 1980–

86 samples with relatively low (or zero) microlite abundance 

also have matrix glass pools large enough for analysis. It was 

not possible to find any large matrix glass pools in samples 

from the current eruption.

After initially examining each sample for melt inclusions 

using a scanning electron microscope (SEM), a subset of the 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 757

Sample No. Eruption date Days since 

Mar. 17, 1980 Sample type (and location)

SH10 Apr. 12, 1980 27 Dense juvenile clast

USNM115379-34 May 18, 1980 62 Cryptodome gray dacite erupted during lateral blast 

SH80D May 18, 1980 62 Cryptodome gray dacite erupted during lateral blast.

Collected from Pumice Plain

SHKB24 May 18, 1980 62 Blast deposit, nonvesicular margin of cryptodome

(UTM 565536E, 5119874N)

C85-310 May 18, 1980 62 Pale gray (microlite-bearing) pumice erupted during

early stage of plinian eruption

MSH006 May 18, 1980 62 Plinian pumice

KC518PFB May 18, 1980 62 Microlite-free pumice, pyroclastic flow. Multiple sample splits

(KCHB, KCPL, PLZ, MAY)

May25 May 25, 1980 69 Fallout pumice

KC612PF June 12, 1980 87 Pumice from pyroclastic flow

KC722U July 22, 1980 127 Pumice

KC807B Aug. 7, 1980 143 Pumice from pyroclastic flow

SHKB23 Aug. 7, 1980 143 Denser pumice in the levee of October 1980 pyroclastic flow 

(UTM 562602E, 5117871N)

USNM115418-60 Oct. 16, 1980 213 Dome fragment

USNM115418-60-2 Oct. 16, 1980 213 Dome fragment

USNM115418-42 Oct. 16, 1980 213 Pumice

USNM115418-61 Oct. 16, 1980 213 Dome

USNM115427-1 Dec. 27, 1980 285 Pumice

USNM115427-4 Dec. 27, 1980 285 Dome

USNM115465 June 18, 1981 458 Dome

KC681 June 18, 1981 458 Dome

USNM115773-18 Mar. 19, 1982 732 Dome

USNM115773-3 Mar. 19, 1982 732 Pumice

SH127 May 14, 1982 788 Dome

SH131 Aug. 18, 1982 884 Dome

SH135 Feb. 7, 1983 1,057 Spine

SH156 June 17, 1984 1,553 Dome (collected in June, probably erupted March 1984)

SH201 May 24, 1985 1,894 Spine

SHKB20 May 8, 1986 2,243 Vesicular dome rock from top of dome

(UTM 562619E, 5116559N)

SHKB21 Oct. 21, 1986 2,409 Light-colored sample from top of dome

(UTM 562619E, 5116559N)

SH304-2A Oct. 18, 2004 8,981 Spine

SH305-1 Nov. 20, 2004 9,014 Spine

SH315-4 Apr. 1, 2005 9,146 Spine

Table 1. Inventory of samples studied.

[Samples prefixed SH (and May 25 sample) are from the Cascades Volcano Observatory collection except SH80D (collected by D. Pyle). Samples prefixed 

USNM are from Smithsonian Institution (see Melson, 1983). Samples prefixed KC are from collection of K. Cashman. Samples prefixed SHKB were collected 

by the authors in September 2003; UTM eastings and northings referable to zone 10, datum WGS84. Sample MSH006 was provided by S. Carey.]

inclusions was selected for analysis of H2O and trace elements 

by ion microprobe. The same inclusions were then analyzed 

by electron microprobe analysis (EMPA). The analyses were 

performed in this order because of the known damage that 

results from EMPA, especially for volatile elements (Humphreys and others, 2006). In choosing melt inclusions for 

analysis we used backscattered electron intensity to select melt 

inclusions with a range of compositions. Because backscatter 

intensity correlates strongly with dissolved H2O, our analyses 

bracket the full range of observed H2O in each sample.

Ion microprobe analyses were carried out on Au-coated 

polished mounts using a CAMECA IMS-4f instrument at 

the University of Edinburgh with a primary beam of O−

 ions 

and detection of positive secondary ions. Typical operating 

conditions were 10 kV (nominal) primary beam and 2–6-nA 

current at the sample surface, corresponding to an 8–15-μm 

sputtered area. To prevent sample charging, a small raster 

(typically ≤10 μm diameter) was applied when analyzing 1

H 

and light trace elements (up to 47Ti). Secondary ions were 

extracted at 4.5 kV with an offset of 75±20 V to reduce transmission of molecular ions. To minimize magnet hysteresis 

we analyzed isotopes in two separate batches. The first batch 

included the light element isotopes: 1H, 7Li, 9Be, 11B, 30Si, 45Sc, and 47Ti. Interference on 45Sc by 29Si16O was monitored 
758 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

using 42Ca and 44Ca and then subtracted by peak-stripping. 

The second batch of isotopes, measured on a subset of melt 

inclusions, included the heavier isotopes: 30Si, 47Ti, 85Rb, 88Sr, 89Y, 90Zr, 93Nb, 133Cs, 138Ba, 139La, 140Ce, 141Pr, 143Nd, 149Sm, 157Gd, 159Tb, 161Dy, 165Ho, 171Yb, 178Hf, 181Ta, 232Th, and 238U. 

Background was monitored at mass 130.5 and found to be 

consistently <0.01 counts per second.

Interferences of light rare-earth-element (REE) oxide ions 

on heavy REE, Ta, and Hf were removed by peak-stripping 

using the oxide/ion ratios of Hinton (1990). The efficacy of the 

peak-stripping procedure was monitored by analyzing two isotopes of Gd (156 and 157) to check for consistency. The high 

Ba content of all glasses and the large interference of BaO on 

both Eu isotopes (BaO/Ba ≈ 0.05) means that Eu cannot be 

precisely determined. The largest oxide corrections are those 

involving Ce (CeO/Ce ≈ 0.21), which imparts an uncertainty 

of ±20 percent (relative) on peak-stripped Gd count rates. 

Count times varied from isotope to isotope, according to 

abundance in the glass, but were always sufficient to generate 

a minimum of 100 counts over the analysis period. For both 

batches of isotopes, 30Si was used as an internal standard, and 

values were corrected for their SiO2 content using the subsequent analysis of each inclusion by EMPA. Trace element 

calibration was carried out using NIST SRM610 multielement 

glass. Analyses of natural-glass secondary standards reveal 

that accuracy is within ±15 percent relative for all elements. 

This indicates that differences in ion yield between SRM610 

and natural silica-rich glasses are small. There is evidence that 

some of these small ion yield differences are systematic, but 

we have not corrected for this effect, which would make only a 

small difference to the data presented.

Water was measured using 1H and a working curve of 1

H/30Si versus H2O based on analysis of 5–12 hydrous andesite, dacite, and rhyolite glasses of known H2O content (0.09–

5.8 wt. percent). A working curve was generated on each day 

of analysis. There are small variations in 1H ion yield between 

different sessions, but the working curves are consistently 

linear with correlation coefficients in excess of 0.99. We 

used 1

H/30Si for calibration, rather than 1H/30Si×SiO2, because 

of the systematic variation in 1H ion yield with matrix SiO2

content (Blundy and Cashman, 2005). Our working curves 

can be used for glasses ranging in composition from basalt to 

rhyolite, although all of the glasses analyzed in this study are 

rhyolitic. In a typical analysis of 15 cycles across the mass 

range, only the final 10 cycles were averaged owing to the 

presence of a small amount of signal instability at the onset of 

the analysis. This routine corresponds to a presputter period 

of a few minutes for 1H analysis. No such effect is observed 

for other light isotopes, for which all 15 cycles were included 

in the averaging.

Electron microprobe analysis for major elements plus 

total S (expressed as equivalent SO2), F, and Cl, was carried 

out on a CAMECA SX100 five-spectrometer WDS instrument, using a range of minerals, oxides, and metals for 

calibration. Operating conditions (2-nA beam current, 15-μm 

diameter spot) were those shown by Humphreys and others 

(2006) to minimize the loss of alkalis (especially Na) during 

analysis. This analytical protocol reproduces the measured 

H2O content (as estimated from the analytical total) of hydrous 

glass standards to within 0.4 weight percent average absolute 

deviation. A few of the major element analyses, obtained early 

in the study, were analyzed on a JEOL-733 four-spectrometer 

instrument using slightly higher beam currents (Blundy and 

Cashman, 2005). These analyses did show some Na loss, 

as measured by ion microprobe analysis of 23Na. For these 

analyses we have used the ion microprobe value of Na2O in 

preference to the EMPA value.

After performing both ion microprobe analysis and 

EMPA, the data were screened for quality according to the 

following criteria: the analytical total, including H2O, must 

lie between 98.5 and 100.5 percent; Ti contents measured by 

EMPA and ion microprobe (heavy and light element routine) 

for a single inclusion must be within 15 percent relative; and 

the ion microprobe spot, as examined by SEM, must not have 

any contact with the host mineral. Analyses that failed one or 

more of these tests were discarded. Representative analyses 

are presented in tables 2 and 3, and the full dataset of accepted 

analyses can be found in appendix 1 (included in the digital 

version of this paper).

Touching Fe-Ti oxide pairs were analyzed by EMPA in 

thin sections of selected 1980–86 samples using a Cameca 

SX100 with 20-nA beam current and a focused spot. We used 

only touching pairs because of the known rapid reequilibration 

of Fe-Ti oxides to changes in temperature and fO2

 (for example, 

Venezky and Rutherford, 1999). By selecting only oxides in 

direct contact it is possible to get the closest approximation 

to equilibrium compositions. This approach is used in preference to the conventional practice of averaging large numbers 

of separate ilmenite and magnetite analyses. We screened each 

magnetite-ilmenite pair for Mg-Mn exchange equilibrium 

using the method of Bacon and Hirschmann (1988) and then 

calculated T–fO2using the recalculation procedure of Spencer 

and Lindsley (1981) and the thermometer of Andersen and 

Lindsley (1988) (table 4; see digital appendix 2 for geochemical analyses of oxide minerals).

Major Element Systematics

Melt inclusions from the 1980–86 eruption are rhyolites 

with 68–79 percent SiO2 (on an anhydrous basis). Matrix 

glasses span the same range. Melt inclusions from the current 

eruption extend the range in SiO2 to 80 percent. As previously 

shown by Blundy and Cashman (2001), the high SiO2 of the 

glasses requires crystallization at low pressures because of the 

increase in SiO2 solubility with decreasing pressure. It is therefore likely that the major element variation in melt inclusions 

records crystallization of magma within the magma chamber 

and conduit.

In figure 1 we compare the major-element chemistry of 

melt inclusions and matrix glasses to whole-rock data from the 
Sample No. Analyzed point Host Na2O Al2O3 SiO2 MgO K2O CaO TiO2 FeO MnO P2O5 Cl F SO2 H2O 

1980–86 eruption sh10 s4-1 P 4.07 11.27 75.12 0.83 2.74 0.66 0.52 3.16 0.04 0.12 0.13 0.10 0.00 0.9

USNM115379-34 11A* O 5.86 14.65 71.15 0.35 2.45 2.21 0.29 2.27 0.26 -- -- -- -- 1.8USNM115379-34 2A* P 6.18 13.63 69.48 0.54 2.25 1.80 0.23 2.00 0.13 0.31 -- -- -- 3.7USNM115379-34 3B P 4.56 12.27 72.57 0.70 2.70 1.07 0.38 2.65 0.05 0.10 0.14 0.04 0.05 2.USNM115379-34 6A* P 5.47 11.78 73.58 0.81 2.76 0.91 0.34 3.15 0.03 -- -- -- -- 2.2SH80D 2A P 4.34 11.17 73.66 0.77 3.45 0.81 0.40 2.52 0.10 0.15 0.15 0.01 0.02 1.9SH80D 8A P 4.42 11.15 75.26 0.30 3.41 0.68 0.52 1.69 0.06 0.11 0.16 0.10 0.00 1.2C85-310 pl3-3 P 4.30 11.41 75.49 0.53 2.61 0.70 0.38 2.32 0.03 0.03 0.16 0.03 0.01 1.9MSH006 gm1* gm 5.93 14.20 71.96 0.49 2.17 2.14 0.32 2.13 -- -- -- -- -- 1.5MSH006 9A* P 5.75 13.74 65.91 0.71 2.15 2.27 0.33 2.03 -- 0.76 -- -- -- 5.9MSH006 10A* P 5.49 13.31 64.91 0.71 1.99 1.95 0.41 2.78 0.02 0.33 -- -- -- 6.3KC518PFB 518b-4-1 P 3.92 14.50 67.94 0.47 1.97 2.54 0.34 1.75 0.01 0.09 0.14 0.06 0.01 5.KC518PFB(KCHB) 7A A 5.54 14.07 69.01 0.36 1.85 2.25 0.37 2.19 0.02 0.00 0.11 0.00 0.15 3.KC518PFB(KCHB) 11A* A 5.65 15.03 65.81 0.37 2.44 2.35 0.51 2.13 0.00 -- -- -- -- 4.9KC518PFB(KCHB) GM3* gm 6.28 14.87 71.10 0.51 1.66 2.62 0.35 2.06 0.16 -- -- -- -- 1.KC518PFB(KCPL) 13A* P 5.76 12.71 69.58 0.26 2.10 1.59 -- 1.01 0.00 0.18 -- -- -- 4.7KC518PFB(MAY) 6-2 P 4.89 17.04 63.99 0.35 1.58 4.15 0.24 1.49 0.00 0.30 0.10 0.07 0.00 6.4KC518PFB(MAY) 02-1 P 4.40 13.88 65.35 0.48 2.01 2.00 0.38 2.17 0.05 0.10 0.14 0.00 0.02 6.7KC518PFB(PLZ) gm3* gm 5.96 15.26 69.44 0.45 1.99 2.29 0.36 2.27 0.00 -- -- -- -- 2.2KC518PFB(PLZ) 33A2* P 6.36 12.84 68.48 0.63 1.89 1.74 0.37 2.18 0.06 0.16 -- -- -- 4.6KC518PFB(PLZ) plz-51-1 P 3.82 14.56 67.65 0.64 1.65 2.36 0.36 2.12 0.09 0.03 0.10 0.36 0.02 4.9May25 PL6-1 P 4.82 13.16 69.97 0.48 2.23 1.52 0.38 2.11 0.02 0.05 0.14 0.12 0.01 4.9KC612PF 14-2 P 5.57 12.61 70.58 0.51 2.45 1.27 0.35 2.37 0.00 0.27 0.14 0.05 0.00 3.9KC612PF 14-GM tube 5.17 11.90 73.07 0.53 2.54 1.08 0.56 2.19 0.06 0.22 0.12 0.02 0.02 2.4KC722U 1-GM* gm 4.62 12.04 77.02 0.20 2.85 0.90 0.45 1.58 -- -- -- -- -- 0.2KC722U 12-1 P 5.08 12.61 71.74 0.28 2.33 1.22 0.37 1.85 0.04 0.12 0.17 0.08 0.09 4.2KC807B 807b-9-1 O 4.75 12.82 73.89 0.18 2.57 1.29 0.38 2.22 0.06 0.00 0.15 0.10 0.02 1.0KC807B 807b-9-2 P 3.42 12.92 70.94 0.67 2.28 1.43 0.39 2.52 0.02 0.08 0.16 0.12 0.01 4.2KC807B 15-1 P 5.04 12.98 72.13 0.55 2.26 1.42 0.37 2.15 0.00 0.07 0.12 0.10 0.05 3.9KC807B 11-1 P 5.71 12.99 71.79 0.30 2.23 1.32 0.19 1.29 0.01 0.00 0.11 0.02 0.08 4.KC807B 17GM tube 4.52 11.65 75.88 0.17 2.85 0.72 0.56 1.65 0.05 0.05 0.18 0.11 0.00 1.3USNM115418-42 c1-gm gm 4.56 11.03 77.69 0.11 2.90 0.53 0.36 1.59 0.19 0.11 0.21 0.00 0.01 0.9

Table 2. Selected major-element and H2O analyses, in weight percent, of subset of inclusion and groundmass glasses from Mount St. Helens. [Entire dataset is in appendix 1; host of inclusion—P, plagioclase; A, amphibole; O, orthopyroxene; C, clinopyroxene; gm, groundmass; tube, tube connecting inclusion and groelectron microprobe except H2O and some Na2O (* follows analyzed point) by ion microprobe. Dashes, element not analyzed. sd, 1 standard deviation of H2O measurement pro

ing statistics and uncertainties on H2O working curve. As described in text, pH2O is H2O saturation pressure in MPa at 900°C from Newman and Lowenstern (2002).]
USNM115418-42 b5-1 P 5.17 12.27 72.99 0.52 2.32 1.13 0.31 1.81 0.02 0.01 0.13 0.12 0.01 2.7USNM115418-60 16A P 4.51 11.57 72.08 0.63 2.76 0.94 0.52 2.40 0.07 0.09 0.15 0.11 0.01 6.3USNM115418-60 17C P 4.41 10.27 71.52 0.67 2.50 0.85 0.40 3.45 0.10 0.08 0.14 0.09 0.37 2.5USNM115418-60-2 pl1-1 A 4.23 14.11 72.86 0.21 3.07 1.17 0.31 2.10 0.02 0.09 0.28 0.04 0.00 1.8USNM115427-1 a7-1 P 4.89 12.00 74.44 0.39 3.20 0.67 0.40 1.70 0.04 0.12 0.19 0.14 0.00 2.0USNM115773-3 pl10-1 P 5.42 12.15 75.01 0.55 2.63 0.87 0.27 2.06 0.09 0.13 0.14 0.21 0.02 0.6USNM115773-3 pl4-1 P 4.53 11.53 74.64 0.41 2.62 0.77 0.41 1.99 0.06 0.10 0.11 0.18 0.00 2.7USNM115773-3 pl8-1 P 5.05 11.72 73.20 0.48 2.93 1.23 0.81 2.70 0.06 0.23 0.26 0.20 0.00 1.3SH127 sh127-5-1 P 5.94 10.68 74.21 0.39 3.33 1.00 0.52 2.01 0.06 0.07 0.19 0.21 0.00 1.0SH131 sh131-1-1 O 3.28 12.80 75.62 0.42 2.74 0.89 0.33 2.08 0.05 0.12 0.19 0.11 0.02 0.3SH131 sh131-3-gm gm 3.91 11.35 78.44 0.08 3.28 0.77 0.35 1.33 0.01 0.09 0.10 0.25 0.02 0.0SH131 sh131-5-1 P 5.26 11.11 75.32 0.13 3.36 0.62 0.40 1.54 0.07 0.00 0.17 0.48 0.00 0.8SH135 sh135-1-1 P 3.58 11.10 75.81 0.09 5.32 0.46 0.48 1.32 0.05 0.09 0.13 0.22 0.03 0.1SH156 sh156-1-1 P 3.68 10.46 76.58 0.21 3.71 0.26 0.43 1.62 0.03 0.07 0.23 0.00 0.00 1.8SH156 sh156-6-gm gm 4.29 11.10 78.34 0.09 3.57 0.41 0.30 1.25 0.01 0.07 0.11 0.03 0.00 0.1SH156 sh156-7-1 P 4.52 10.88 77.19 0.16 3.42 0.61 0.32 1.24 0.01 0.03 0.17 0.24 0.01 1.0SH156 sh156-8-2 A 6.18 15.23 72.44 0.12 3.48 0.77 0.30 1.60 0.03 0.09 0.21 0.07 0.00 0.1SHKB21 shkb21-2-1 C 4.71 11.02 75.61 0.08 5.12 0.40 0.33 1.40 0.04 0.10 0.24 0.11 0.00 0.1Current eruption SH304-2A sh304-1-1 A 2.83 13.84 72.70 0.84 5.65 0.82 0.29 2.24 0.09 0.08 0.10 0.21 0.00 0.1SH304-2A sh304-2-1 O 6.63 19.05 69.70 0.27 0.68 3.65 0.09 0.89 0.01 0.09 0.03 0.00 0.01 0.0SH304-2A sh304-4-2 P 4.68 10.32 77.34 0.07 2.62 0.15 0.05 0.73 0.06 0.05 0.04 0.00 0.01 2.9SH305-1 a1-1 A 5.27 13.56 72.08 0.08 6.04 0.06 0.44 1.78 0.16 0.07 0.19 0.00 0.04 0.1SH305-1 a3-1 A 3.75 12.07 76.56 0.08 5.47 0.38 0.27 1.27 0.00 0.01 0.12 0.00 0.02 0.2SH315-4 a11-1 O 2.73 12.46 76.48 0.46 4.64 0.15 0.22 2.27 0.15 0.05 0.08 0.33 0.00 0.0SH315-4 d4-1 P 3.91 13.44 73.94 0.25 4.27 1.79 0.19 1.17 0.05 0.07 0.12 0.15 0.04 0.0

Table 2. Selected major-element and H2O analyses, in weight percent, of subset of inclusion and groundmass glasses from Mount St. Helens. —Co[Entire dataset is in appendix 1; host of inclusion—P, plagioclase; A, amphibole; O, orthopyroxene; C, clinopyroxene; gm, groundmass; tube, tube connecting inclusion and groelectron microprobe except H2O and some Na2O (* follows analyzed point) by ion microprobe. Dashes, element not analyzed. sd, 1 standard deviation of H2O measurement pro

ing statistics and uncertainties on H2O working curve. As described in text, pH2O is H2O saturation pressure in MPa at 900°C from Newman and Lowenstern (2002).]

Sample No. Analyzed point Host Na2O Al2O3 SiO2 MgO K2O CaO TiO2 FeO MnO P2O5 Cl F SO2 H2
Sample Analyzed point Host Li B Sc Rb Sr Y Zr Nb Cs Ba La Ce Pr Nd Sm Gd Dy Ho 1980–86 eruption USNM115379-34 11A O 57 -- 2.0 44 200 7.3 183 6.2 2.1 383 13.6 27.4 3.2 14.4 3.0 -- -- -- 

USNM115379-34 3B P 48 -- 5.6 49 68 11.2 206 6.6 2.8 447 14.9 32.3 3.3 18.1 3.7 4.2 3.5 0.5USNM115379-34 6A P 47 -- 6.5 51 61 11.9 226 9.5 2.6 460 16.2 34.5 3.9 17.3 3.4 3.7 3.1 0.5SH80D 2A P 33 -- 7.1 69 57 11.6 224 6.9 -- 361 15.0 28.7 3.7 -- -- -- -- -- SH80D 8A P 32 -- 7.5 58 38 12.0 234 8.3 -- 378 15.7 33.0 4.0 -- -- -- -- -- MSH006 gm1 gm 34 -- 5.4 40 180 9.5 179 6.6 1.8 360 13.3 26.8 3.5 12.6 2.6 2.0 2.6 0.4MSH006 9A P 32 -- 3.8 36 212 7.6 154 5.4 1.5 317 11.9 22.9 2.9 12.1 2.4 2.1 2.1 0.3MSH006 10A P 38 -- 4.0 34 192 6.9 136 5.2 -- 276 11.3 21.8 2.2 -- -- -- -- -- KC518PFB(KCHB) 11A A 32 -- 3.2 47 267 7.0 320 28.8 1.8 439 20.3 35.8 3.7 12.6 3.4 2.1 1.9 0.4KC518PFB(KCHB) GM3 gm 35 -- 4.5 36 224 8.5 180 5.8 1.5 341 12.4 26.3 3.2 13.4 2.7 -- -- KC518PFB(KCPL) 13A P 32 -- 1.9 45 190 2.9 87 4.3 3.0 421 11.6 20.7 2.1 8.6 1.5 1.1 -- -- KC518PFB(PLZ) 33A2 P 37 -- 6.5 50 195 11.0 240 8.6 2.7 454 17.9 36.2 4.1 17.4 3.2 3.3 2.5 0.5KC612PF 14-2 P 47 25 4.8 45 88 9.5 176 5.9 2.5 424 14.1 26.5 3.1 12.3 3.0 3.3 -- 0.5KC612PF 14-GM tube 42 25 5.6 52 64 9.5 205 6.8 2.8 385 12.3 28.2 3.3 15.1 3.5 3.6 2.8 0.5KC722U 1-GM gm 37 31 6.3 49 150 12.0 233 8.5 2.6 452 15.1 33.6 4.0 16.8 3.2 4.2 3.5 0.7KC722U 12-1 P 58 21 3.2 40 129 7.7 168 6.5 2.3 479 14.3 27.8 3.1 13.7 2.8 2.1 2.3 0.4KC807B 807b-9-1 O 37 23 5.8 45 86 11.1 230 8.9 2.4 426 14.4 31.2 3.8 14.7 3.1 2.6 2.9 0.6KC807B 807b-9-2 P 61 21 5.7 44 118 9.5 205 8.3 2.0 459 15.6 33.7 3.8 16.6 3.1 2.8 2.9 0.5KC807B 15-1 P 98 25 5.7 50 117 8.5 191 6.9 3.3 425 13.4 27.1 3.3 13.0 2.5 2.7 2.9 -- 

KC807B 11-1 P 89 25 3.6 47 144 7.8 155 6.2 2.5 423 13.1 26.3 3.4 12.9 3.1 2.2 2.3 0.5

USNM115418-42 b5-1 P 60 23 5.6 40 89 8.8 186 5.8 2.3 385 13.3 26.7 3.0 12.2 2.3 3.5 2.3 0.4USNM115427-1 a7-1 P 57 21 9.1 42 38 4.1 239 11.6 1.7 492 4.3 8.8 1.0 3.8 0.6 0.7 1.1 0.2USNM115773-3 pl10-1 P 34 25 3.2 41 56 5.2 214 5.6 2.1 354 13.0 25.0 2.6 10.2 1.6 2.5 1.4 0.3USNM115773-3 pl4-1 P 65 19 4.3 43 57 7.6 185 6.2 1.9 346 13.6 27.9 3.0 12.8 2.0 2.8 2.0 0.4SH127 sh127-5-1 P 118 29 7.9 61 36 14.6 264 11.0 2.8 457 17.3 37.3 4.1 18.7 3.9 2.7 3.6 0.7SH131 sh131-1-1 O 23 37 4.6 49 42 13.5 213 7.6 1.9 479 15.9 32.0 3.8 16.5 3.3 2.4 2.9 0.7SH131 sh131-3-gm gm 38 37 5.3 50 30 13.6 286 9.6 2.4 417 14.3 32.3 3.3 15.5 3.3 2.8 3.4 0.6SH131 sh131-5-1 P 71 33 5.8 53 24 13.5 268 9.1 2.3 408 15.8 35.0 4.2 15.1 3.1 -- 3.4 0.6SH156 sh156-1-1 P 276 28 5.9 64 22 12.0 235 8.9 3.2 498 18.5 37.1 4.2 17.8 3.2 2.6 3.0 0.6SH156 sh156-7-1 P 63 33 4.6 58 26 11.1 275 8.7 3.4 416 13.7 29.2 3.3 14.1 2.9 -- 3.0 0.6SHKB21 shkb21-2-1 C 122 44 3.0 66 15 14.4 379 8.3 3.2 260 18.1 38.8 4.6 18.7 3.4 2.9 3.6 0.7Current eruption SH304-2A sh304-1-1 A 60 24 1.5 172 66 8.7 150 6.4 2.4 512 12.8 27.1 3.2 12.3 2.3 2.4 1.8 0.3SH304-2A sh304-4-2 P 169 28 3.0 53 156 4.6 42 4.9 3.2 370 9.6 19.6 2.2 8.7 1.9 2.0 1.2 0.2SH305-1 a3-1 A 159 27 11.8 109 22 11.2 172 5.3 4.0 356 14.0 28.8 3.5 15.2 3.2 3.8 3.3 0.6SH315-4 d4-1 P 28 34 3.4 187 7 7.0 148 5.6 4.2 116 12.1 26.4 2.8 12.1 2.1 -- 1.5 -- 

Table 3. Selected trace-element analyses, in parts per million, of subset of inclusion and groundmass glasses from Mount St. Helens. [Entire dataset is in appendix 1; host of inclusion—P, plagioclase; A, amphibole; O, orthopyroxene; C, clinopyroxene; gm, groundmass; tube, tube connecting inclusion and groion microprobe. Dashes, element not analyzed.]
762 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample Texture Pair T, °C log fO2

1980–86 eruption

SH80D P SH80Dpr3 891 -11.79

C85-310 P C85_310 pr3 942 -10.54

C85-310 I C85_310 pr6 inc opx 870 -11.23

C85-310 G C85_310 pr7tiny 904 -11.55

KC518PFB P 518pfapr1 874 -11.51

KC518PFB P 518pfapr4 899 -11.04

May-25 P may25pair2 908 -10.94

May-25 P may25pr1 958 -10.50

May-25 P may25pr2 914 -11.44

May-25 P may25pr8 869 -12.06

KC612PF P june12pr3 877 -11.71

KC612PF P june12p6 901 -11.29

KC722U P 722upr4 895 -11.37

KC722U P 722upr6 878 -11.80

SHKB23 I SHKB23pr1 incPl 859 -12.32

SHKB23 G SHKB23pr2late gm 872 -11.99

SHKB23 P SHKB23pr3pheno 949 -10.93

KC681 P 681 pair 2 902 -11.61

KC681 P 681 pair 4 751 -14.65

KC681 P 681 pair 5 950 -10.73

SH131 G SH131gm 862 -12.39

SH131 P SH131pr8 883 -11.21

SH131 G SH131pr9gm 911 -11.33

SH135 P SH135_10pr6 479 -29.05

SH156 E SH156pr3ex 871 -12.07

SH156 P SH156pr4 864 -11.90

SH156 G SH156pr8tiny 927 -10.63

SH156 G SH156pr9tiny 864 -12.19

SH156 P SH156pr11 899 -11.40

SH201 P sh201pr2 833 -13.41

SH201 P sh201pr3 857 -11.76

SHKB20 P SHKB20pr6 848 -12.35

SHKB20 P SHKB20pr6bpheno 893 -11.83

Current eruption

SH315-4 P a5pair7 850 -12.19

SH315-4 P a5pair2rpt 868 -11.95

SH315-4 P a5pair6gmsrpt 1,022 -9.96

Table 4. Calculated temperature and oxygen fugacity (fO2) for 

selected touching pairs of magnetite and ilmenite in Mount St. 

Helens samples from 1980–86 and current (2004–2006) eruptions.

[All data plotted in figure 12. Each oxide pair is identified in the column 

labeled “pair,” for cross-referencing to appendix 2 (in digital versions of this 

work in CD and on Web), which contains the entire dataset and full oxide 

analyses. Texture: P, phenocryst; G, groundmass; I, inclusion in silicate phenocryst; E, exsolution lamellae and host.]

1980–86 and current eruptions, as well as to data for magmas 

erupted at Mount St. Helens throughout late Quaternary time, 

taken from a compilation of 94 published analyses (Cashman 

and Taggart, 1983; Criswell, 1987; Fruchter and others, 1980; 

Gardner and others, 1995; Halliday and others, 1983; Hooper 

and others, 1980; Irving and others, 1980; Leeman and others, 

1990; Melson, 1983; Pallister and others, 1992; Rutherford and 

Devine, 1988; Sarna-Wojcicki and others, 1981; Scheidegger 

and others, 1982; Smith and Leeman, 1987, 1993; Smithsonian Institution, 1980) and eight additional unpublished 

analyses. These data show a number of key systematics. 

Whole-rock compositions of 1980–86 magmas and those from 

the current eruption are silicic andesite or dacite with similar 

major-element compositions, which lie within the overall 

range for pre-1980 late Quaternary magmas from Mount St. 

Helens. In terms of both compatible (CaO, Al2O3, MgO, FeO, 

TiO2) and incompatible (K2O, Na2O) oxides, the glasses all lie 

at and beyond the high SiO2 end of the whole-rock data.

The relatively smooth chemical trends through the 

whole-rock data are consistent with derivation of the dacites 

by fractional crystallization or partial melting of basaltic rocks 

at depth (Smith and Leeman, 1987) or a combination of these 

processes. We cannot rule out mixing of magmas of different SiO2 contents to generate some of the chemical diversity 

in the whole rocks, especially those with >62 percent SiO2. 

However, marked inflections in the trends for some major and 

trace elements require a change in fractionating or residual 

mineralogy during differentiation. For example, the inflection 

in Al2O3 (fig. 1A) and TiO2 (fig. 1D) at ~60 percent SiO2 is 

interpreted as the arrival of plagioclase as a controlling phase, 

and the inflection in the FeO versus MgO plot at ~1.5 percent 

MgO (fig. 1C) marks the arrival of magnetite. The continuous 

linear drop in CaO with increasing SiO2 (fig. 1B) is a result of 

the involvement of amphibole and clinopyroxene in addition to 

plagioclase, such that the bulk CaO content of the crystalline 

residue remains approximately constant.

The location of the glass analyses at the high-SiO2 extrapolation of the whole-rock trends suggests that they represent 

the continued crystallization of magma to low pressures. The 

fact that the dacites in which the glasses are found have such 

limited compositional range, whereas the glasses themselves 

cover a wide compositional range, even within a single 

sample, indicates that low-pressure crystallization involved 

very little change in bulk composition. In other words, the 

glasses record a process of closed-system crystallization in 

which some crystals (for example, plagioclase) are overgrown 

by rims of different composition, whereas others (for example, 

mafic minerals, oxides) are able to partially or fully reequilibrate. There is no evidence for significant physical removal 

of crystals from the melt, because the compositional vectors 

defined by the glasses are oblique to those defined by their 

host rocks. This is consistent with the high viscosity of silicic 

melts, which would preclude efficient crystal separation on 

appropriate time scales.

The variation in K2O with SiO2 in glasses (fig. 1E) curves 

strongly upward at high SiO2, consistent with the saturation 

of a silica phase (tridymite or quartz), as observed in the most 

evolved glasses of both the 1980–86 (Blundy and Cashman, 

2001) and current eruptions (Pallister and others, this volume, 

chap. 30). The elevated K2O of glasses from the current eruption distinguishes them from those of the 1980–86 eruption, 

although a single glass from the current eruption plots at 

anomalously low K2O (0.67 percent). The overall trend suggests that most of the glasses of the current eruption reached 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 763

8

12

16

20

24

45 50 55 60 65 70 75 80 85

Al2O3, IN WEIGHT PERCENT

CaO, IN WEIGHT PERCENT

K2O, IN WEIGHT PERCENT

Na2O, IN WEIGHT PERCENT

TOTAL FeO, IN WEIGHT PERCENT

TiO2, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

MgO, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

0

2

4

6

8

10

45 50 55 60 65 70 75 80 85

0

2

4

6

8

10

12

0 1 2 3 4 5 6 7 8

0

0.5

1

1.5

2

2.5

45 50 55 60 65 70 75 80 85

0

1

2

3

4

5

6

7

45 50 55 60 65 70 75 80 85

EXPLANATION

Pre-1980 (bulk)

1980–86 (bulk)

Current (bulk)

1980–86 (MI-plag)

1980–86 (MI-hbl)

1980–86 (MI-opx)

1980-86 (MI-cpx)

1980–86 (gm)

Current (MI-plag)

Current (MI-hbl)

Current (MI-opx)

0

1

2

3

4

5

6

7

8

45 50 55 60 65 70 75 80 85

A B

C D

E F

Figure 1. Major-element variation diagrams (normalized to 100 percent anhydrous) for Mount St. Helens whole rocks of 

late Quaternary age (sources given in text) and melt inclusion (MI) and groundmass (gm) glasses from 1980–86 and current 

(2004–2006) eruptions (data from table 2). All glass data by electron microprobe analysis. Explanation indicates host mineral 

for melt inclusions (MI): plag, plagioclase; hbl, hornblende; opx, orthopyroxene; cpx, clinopyroxene. A, Al2O3–SiO2. B, CaO–

SiO2. C, FeOtot–MgO. D, TiO2–SiO2. E, K2O–SiO2. F, Na2O–SiO2.
764 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

the silica-saturation surface at low pressure, whereas silica 

saturation was relatively rare in 1980–86 glasses. Silica saturation requires higher degrees of crystallization, cooler temperatures, or an initially more SiO2-rich melt.

The Na2O-SiO2 variation (fig. 1F) is unusual in that it 

shows a marked inflection between the whole-rock data, which 

show increasing Na2O with SiO2, and the glasses, which show 

the opposite. This cannot be a result of analytical error, as the 

technique used for glass analyses effectively eliminates Na loss 

(Humphreys and others, 2006). The decrease in Na2O with SiO2

in the glasses cannot be generated by fractionation of an Na-rich 

crystal phase, because the most sodic phase in any of these 

rocks is groundmass plagioclase (An35) with ~7 percent Na2O. 

Instead, we suggest that Na is preferentially partitioned into 

the exsolving vapor phase during volatile-saturated crystallization, causing it to decrease in the melt. This is consistent with 

the elevated Na content of fumarole gases from the volcano 

(Symonds and Reed, 1993). The slight increase in Na2O with 

SiO2 in the whole-rock data suggests that, in contrast to the 

glasses, differentiation did not involve a free-vapor phase. The 

fact that the glass trend is markedly oblique to the whole-rock 

trend indicates that the onset of low-pressure crystallization 

and the onset of volatile saturation were almost coincident, as 

is typical of volatile-saturated decompression crystallization 

(Blundy and Cashman, 2001, 2005; Annen and others, 2006).

Trace Element Systematics

In figure 2 we compare the trace-element concentrations 

in glasses with the whole-rock data for magma erupted before 

1980, from 1980 to 1986, and during the current eruption. The 

trace-element behavior can be divided into compatible elements (for example, Sr, fig. 2A), which decrease with increasing SiO2; incompatible elements (Cs, Rb, Th, figs. 2B–D), 

which increase; and those elements that show inflections (Ba, 

Zr, Nb, Y, figs. 2E–H). The compatible behavior of Sr at ≥65 

percent SiO2 is consistent with plagioclase becoming the controlling phase. The curvature of the Sr-SiO2 trend (fig. 2A) is 

reminiscent of the Al2O3-SiO2 trend (fig. 1A) and testifies to an 

0

100

200

300

400

500

600

700

800

45 50 55 60 65 70 75 80 85

Sr, IN PARTS PER MILLION

Cs, IN PARTS PER MILLION

Rb, IN PARTS PER MILLION

Th, IN PARTS PER MILLION

0

1

2

3

4

5

45 50 55 60 65 70 75 80 85

0

50

100

150

200

45 50 55 60 65 70 75 80 85

0

1

2

3

4

5

6

7

8

45 50 55 60 65 70 75 80 85

A B

C D

SiO2, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

Figure 2. Trace element variation versus SiO2 (normalized to 100-percent anhydrous) for Mount St. Helens whole rocks of 

late Quaternary age (sources as in fig. 1) and melt inclusion and groundmass glasses from 1980–86 and current (2004–2006) 

eruptions (data from tables 2 and 3). A, Strontium. B, Cesium. C, Rubidium. D, Thorium. E, Barium. F, Zirconium. G, Niobium. H, 

Yttrium. Symbols as in figure 1.
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 765

increasingly important role for plagioclase as differentiation 

proceeds. This is allied to the fact that the plagioclase-melt 

partition coefficient for Sr (DSr) increases as plagioclase An 

content decreases (Blundy and Wood, 1991).

The increase of Cs, Rb, and Th with increasing SiO2

reflects their high incompatibility in all phenocryst phases at 

Mount St. Helens. Trends for Cs and Th, although showing 

greater scatter, are similar for the 1980–86 and current eruptions, showing a steady increase with increasing SiO2 in both 

whole rocks and glasses. In contrast, the behavior of Rb at 

high SiO2 is enigmatic. Three glasses for the current eruption 

are displaced to very high Rb, as previously noted for K2O. 

It is hard to ascribe this to silica saturation alone, because 

there is no such displacement in Cs, Ba, or Th. Decoupling of 

Rb (and K) from other trace elements could be generated by 

involvement of a crystal phase with high partition coefficients 

for Rb and K, such as mica. However, as micas have DBa>DRb

(Icenhower and London, 1995), the effect should be greater for 

Ba than for Rb, which is the opposite of what is observed. We 

suggest instead that high initial Rb is a distinctive feature of a 

small number of melt inclusions from the current eruption.

The slight inflection in the Ba-SiO2 trend (fig. 2E) at 

>74 percent SiO2 is consistent with the increase in DBa for 

plagioclase with decreasing An content (Blundy and Wood, 

1991). According to the model of Blundy and Wood, at 900°C 

Ba becomes compatible in plagioclase at An27, a composition 

slightly more sodic than the most An-poor microlite observed 

at Mount St. Helens (An33). However, as this is within the 

uncertainty of the Blundy and Wood model, we suggest that 

crystallization of sodic plagioclase microlites is the most 

likely cause of the inflection to lower Ba at high SiO2. Early 

stages of fractionation involved Ca-rich plagioclase, and so Ba 

remained incompatible.

The inflections in the behavior of Zr, Nb, and Y (figs. 

2F–H) cannot be attributed to microlite crystallization. These 

elements behave incompatibly during differentiation of the 

glasses (that is, concentrations increase with SiO2), but compatibly (decrease with SiO2) in the differentiation trend from 

0

5

10

15

20

25

30

35

45 50 55 60 65 70 75 80 85

0

5

10

15

20

25

30

45 50 55 60 65 70 75 80 85

G H

Nb, IN PARTS PER MILLION

SiO2, IN WEIGHT PERCENT

Y, IN PARTS PER MILLION

SiO2, IN WEIGHT PERCENT

Ba, IN PARTS PER MILLION

Zr, IN PARTS PER MILLION

0

100

200

300

400

500

600

700

800

45 50 55 60 65 70 75 80 85

0

100

200

300

400

500

45 50 55 60 65 70 75 80 85

E F

SiO2, IN WEIGHT PERCENT SiO2, IN WEIGHT PERCENT

Figure 2.—Continued.
766 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

basalt to dacite. This behavior was previously noted by Smith 

and Leeman (1987), who attributed the low concentration of 

many incompatible trace elements in dacite to their generation by melting of basalt in the lower crust. During this process, garnet was stable in the residue and residual amphibole 

occurred at much higher modal proportion than it occurs as a 

phenocryst phase in the dacite. We agree with this interpretation in a general sense, although we note that crystallization 

of hydrous basalt, rather than melting of basalt, can also produce residues with garnet and modally abundant amphibole. 

Annen and others (2006) have argued on the basis of thermal 

models that deep crustal crystallization of hydrous basalt is 

the most effective method of producing andesite and dacite. 

We suggest that high-pressure crystallization accounts for the 

trends of decreasing Y, Zr, and Nb with increasing SiO2 in 

the bulk rocks. Garnet and/or amphibole crystallization best 

explains the behavior of Y and possibly Zr, whereas crystallization of rutile or ilmenite can account for the behavior of Nb. 

We note that the Zr concentrations in all Mount St. Helens 

rocks are too low for zircon saturation at ≥850°C (Watson 

and Harrison, 1983), and it is unlikely that this mineral plays 

a role in controlling the behavior of Zr in the whole rocks. It 

is possible, however, that the three melt inclusions with Zr 

concentrations <100 ppm in figure 2F attained zircon saturation during cooling, although this would require a temperature ≤720°C (Watson and Harrison, 1983), which is below 

the H2O-saturated haplogranite solidus at pressures ≤50 MPa 

(Johannes and Holtz, 1996).

Detailed quantitative modeling of the whole-rock and 

melt-inclusion differentiation trends is beyond the scope of 

this paper. However, we conclude that the whole-rock chemical variations seen in figures 1 and 2 are consistent with a 

high-pressure, vapor-undersaturated differentiation episode to 

generate the silicic andesite and dacite and that this is followed 

by low-pressure, vapor-saturated crystallization to generate 

the compositional variation in the glasses. Mixing of more 

and less differentiated magmas at depth (Annen and others, 

2006) or entrainment of crystalline residues from depth may 

be responsible for some of the linearity observed in the wholerock data, especially in rocks with >62 percent SiO2. However, 

the overall whole-rock trend from basalt to dacite cannot be 

attributed to magma mixing alone.

For the 1980–86 eruption, melt inclusions hosted by 

plagioclase, amphibole, orthopyroxene, and clinopyroxene lie 

on the same overall trends (figs. 1, 2). In keeping with Rutherford and Devine (1988), we observe that amphibole-hosted 

inclusions are confined to the SiO2-poor end of the array, but 

they do not show a consistent offset relative to plagioclase- or 

pyroxene-hosted inclusions. This observation indicates that the 

major and trace-element chemistry of a melt inclusion is not a 

product of simple closed-system postentrapment crystallization or dissolution of the host mineral. If it were, then elements 

that are compatible in only one host phase (for example, Sr in 

plagioclase) would show different behavior from one host mineral to another; but figure 2A indicates that this is not the case. 

Similarly, we see no consistent offset in MgO and FeO contents 

of those inclusions hosted in mafic minerals compared to those 

in plagioclase (fig. 1C; Blundy and Cashman, 2005).

Rare-Earth Elements (REE)

Rare-earth elements (except Eu) have been measured in 

melt inclusions and groundmass glasses from the 1980–86 

eruption and from melt inclusions in three samples of the 

current eruption. Selected data are plotted on chondritenormalized (Sun and McDonough, 1989) variation diagrams 

in figures 3B–D, together with whole-rock data (fig. 3A) from 

the literature and from Pallister and others (this volume, chap. 

30). As noted by Pallister and others, bulk rock REE patterns of samples from both the current and 1980–86 eruptions 

are similar. The only difference is a lower overall content of 

heavy REE in the current eruption. Neither eruption produced 

rocks with any detectable Eu anomaly (fig. 3A), despite the 

abundance of plagioclase as a major crystallizing phase at low 

pressure. This is consistent with the suggestion above that 

much of the chemical variation in the basalt to dacite magmas 

at Mount St. Helens was generated in the lower crust (~30 

km depth), where plagioclase was a minor phase (Berlo and 

others, 2007). Elevated fO2

 in these magmas would also serve 

to minimize any Eu anomaly due to the low Eu2+/Eu3+ ratio. 

Plagioclase appears to have become a major crystallizing 

phase only above 62 percent SiO2. We cannot rule out some 

plagioclase fractionation followed by plagioclase addition at a 

later stage to eliminate any Eu anomaly, although it would be 

surprising if this process precisely eliminated the Eu anomaly 

in all samples analyzed.

The melt inclusion and groundmass glasses have slightly 

elevated REE concentrations relative to the bulk rocks owing 

to crystallization. This crystallization demonstrably involved 

plagioclase, and we would expect the glasses to show negative Eu anomalies if this element could be measured by ion 

microprobe. Plagioclase has higher partition coefficients for 

the light REE relative to the heavy REE; hence the increase 

in heavy REE concentration with fractionation is greater than 

that of the light REE. This is clearly seen in the groundmass 

glasses (fig. 3B). A characteristic of almost all inclusion and 

groundmass glasses is a flattening out of the REE patterns 

between Sm and Yb, leading to a relatively low chondritenormalized ratio (Sm/YbN) of ~2. This behavior is characteristic of silicic magmas that have equilibrated with amphibole, 

which has elevated, but near-constant, partition coefficients for 

Sm to Yb (Sisson, 1994). The Sm/YbN of most glasses and the 

whole rocks is broadly similar, suggesting that at least some 

of the amphibole fractionation occurred in the source region 

of the dacites, in accord with inferences from trace-element 

systematics. Melt inclusions in amphibole and orthopyroxene 

from the current eruption (fig. 3C) are broadly similar in REE 

chemistry to those of 1980–86, with Sm/YbN ≈ 2.

Most plagioclase-hosted melt inclusions have similar 

REE patterns to melt inclusions in mafic phenocrysts and 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 767

1

101

102

La Ce Nd Sm Eu Gd Tb Ho Tm Yb Lu

Whole rocks

July 22, 1980

May 18, 1980

May 18, 1980

June 12, 1980

Oct 1986

SH304-2B

SH306

SH315-1

SH318-1

SH323-1

La Ce Pr Nd Sm Gd Tb Dy Ho Yb

Groundmass

glass

006-GM1

PLZ-GM1

JUL-1-GM

773-18-GM1

SH127-3-gm

SH131-3-gm

JUN-14-GM

1

101

102

1

101

102

1

101

102

La Ce Pr Nd Sm Gd Tb Dy Ho Yb

Melt inclusions in

mafic minerals

807b-9-1(opx)

SH131-1-1(opx)

427-1-A6-1(cpx)

KCHB-11A(amph)

KCHB-11B(amph)

SH304-1-1(amph)

SH305-1-A3-1(amph)

CHONDRITE (C1)-NORMALIZED

CONCENTRATION (LOG SCALE)

CHONDRITE (C1)-NORMALIZED

CONCENTRATION (LOG SCALE)

CHONDRITE (C1)-NORMALIZED

CONCENTRATION (LOG SCALE)

CHONDRITE (C1)-NORMALIZED

CONCENTRATION (LOG SCALE)

La Ce Pr Nd Sm Gd Tb Dy Ho Yb

Melt inclusions

in plagioclase

006-9A

KCPL-6A

KCPL-11A

JUL-14-1

SH131-4-1

SH156-1-1

SH304-4-2

427-1-A7-1

KCPL-12A

A B

C D

Figure 3. Chondrite-normalized REE patterns for selected glasses (data from table 3) and whole rocks from Mount St. 

Helens. Samples from current (2004–2006) eruption shown in red. A, Whole rocks (Halliday and others, 1983; Smith and 

Leeman, 1987; Pallister and others, this volume, chap. 30). B, Groundmass glasses. C, Melt inclusions in mafic minerals, 

identified in parentheses in legend. D, Melt inclusions in plagioclase. Low value for Sm in KCPL-12A is probably an 

analytical artifact. Note that horizontal axis differs slightly between panels A and B–D.

groundmass glasses, with Sm/YbN ≈ 2 (fig. 3D). However, of 

the 172 plagioclase-hosted inclusions from 1980–86, we have 

identified two with anomalous patterns. The first (KCPL12A), from the plinian phase of May 18, 1980, has similar 

overall light REE concentrations but a much higher Sm/YbN

of 4. The second (427-1-A7-1), from December 27, 1980, 

has very low REE concentrations and a distinctive spoonshaped pattern, strongly suggestive of amphibole fractionation from the melt inclusion after entrapment. The steep 

pattern from the May 18 sample cannot have been generated 

by any postentrapment process and appears to represent a 

chemically distinctive, but rare, batch of melt within the 

system, trapped in plagioclase before the melt could be fully 

homogenized with the rest of the melt in the system. Two 

plagioclase-hosted melt inclusions from SH304-2A, an early 

erupted sample from the current eruption (see Pallister and 

others, this volume, chap. 30), also show significantly steeper 

REE patterns (for example, SH304-4-2 on fig. 3D), similar to 

the plagioclase-hosted inclusion of May 18, 1980, described 

above. This similarity is borne out by other distinctive chemical features, including significantly lower U, Th, Y, TiO2, Zr, 

and Hf and higher Sr for its SiO2 concentration (table 3).

Given the limited number of melt inclusions analyzed 

from the current eruption compared to 1980–86, the fact 

that two plagioclase-hosted melt inclusions show anomalous 

chemistry suggests that this component was considerably more 

common in the current eruption than in 1980–86. The preservation of melt inclusions with distinctive chemistry suggests 

a magma system that is continually replenished with melts 

from depth. These melts can be trapped sufficiently rapidly 

that extensive chemical interaction with the dominant matrix 

melt cannot occur. Humphreys and others (2008) drew similar 

conclusions regarding “exotic” melt inclusions from Shiveluch 

Volcano, Kamchatka.

There are several chemical lines of evidence to support the involvement of a new type of magma in the current 

eruption. When other trace elements are plotted against Sm/

YbN, three anomalous melt inclusions lie at the high Sm/YbN
768 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

extreme of the trend defined by the majority of melt inclusions 

and whole rocks (fig. 4). The anomalous melts are defined 

by slightly lower U, Th, and high-field-strength elements 

(HFSE). Significantly, the whole-rock trace element chemistry of the samples from the current eruption is also displaced 

towards this high Sm/YbN component (fig. 4), consistent with 

its greater prevalence in the current eruption than in 1980–86 

(Pallister and others, this volume, chap. 30). Conversely, the 

major-element chemistry of this component is not sufficiently 

different from that of other melt components in the system to 

significantly modify the major element chemistry of magmas 

from the current eruption. We suggest that input of melts with 

this chemistry into the magmatic system before the current 

eruption has played an important but subtle role in modifying 

its bulk trace-element chemistry. The only whole-rock analyses from Mount St. Helens that lie close to this postulated high 

Sm/YbN, low-HFSE component (fig. 4) are from the Ape Canyon eruptive stage (300–35 ka) and the Smith Creek eruptive 

period (3.3–3.9 ka) (Halliday and others, 1983; Clynne and 

others, this volume, chap. 28).

Water

Using a subset of the data presented here for the various 

1980 eruptive phases, Blundy and Cashman (2005) showed 

that the variation in H2O with SiO2 is consistent with vaporsaturated crystallization in response to decompression. Those 

data, augmented by new data for May 18 and July 22, 1980, 

and June 18, 1981, are presented in figure 5A. The data 

describe a trend of decreasing H2O with increasing SiO2. The 

maximum H2O content is 6.7 percent, in a plagioclase-hosted 

melt inclusion from the plinian phase of May 18, 1980; the 

minimum is at the ion microprobe detection limit (~0.04 

weight percent). The lowest H2O (and highest SiO2) values 

occur in highly crystalline samples erupted during the preplin0

5

10

15

20

25

30

35

0 1 2 3 4 5

0

50

100

150

200

250

300

350

400

0 1 2 3 4 5

0

0.5

1

1.5

2

2.5

3

0 1 2 3 4 5

0

1

2

3

4

5

6

7

0 1 2 3 4 5

A B

C D

Nb, IN PARTS PER MILLION U, IN PARTS PER MILLION

Zr, IN PARTS PER MILLION Th, IN PARTS PER MILLION

Sm/YbN, NORMALIZED TO CHONDRITIC VALUES Sm/YbN, NORMALIZED TO CHONDRITIC VALUES

Sm/YbN, NORMALIZED TO CHONDRITIC VALUES Sm/YbN, NORMALIZED TO CHONDRITIC VALUES

Pre-1980 (bulk)

EXPLANATION

1980–86 (bulk)

2004–5 (bulk)

Plinian (MI-plag)

1980–86 (MI-plag)

1980–86 (MI-opx)

1980–86 (MI-cpx)

1980–86 (gm)

2004–5 (MI-plag)

2004–5 (MI-hbl)

Figure 4. Variation in selected trace-element concentrations as function of chondrite-normalized Sm/YbN for Mount St. 

Helens whole rocks of late Quaternary age (sources as in fig. 1) and melt inclusion and groundmass glasses from 1980–86 and 

current (2004–2006) eruptions (data from table 3). A, Niobium. B, Zirconium. C, Uranium. D, Thorium.
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 769

0

1

2

3

4

5

6

7

66 68 70 72 74 76 78 80

1980–81

Apr 10 (plag)

Crypto (plag)

Crypto (opx)

May 18 (plag)

May 18 (hbl)

May 25 (plag)

Jun 12 (plag)

EXPLANATION

All 1980 except June 1981 sample

Jul 22 (plag)

Aug 7 (plag)

Aug 7 (opx)

Oct 16 (plag)

Oct 16 (hbl)

Dec 27 (plag)

Dec 27 (cpx)

Jun 1981 (plag)

May 18 (gm)

1980 (gm)

Ruptured

inclusions

0

1

2

3

4

5

6

7

66 68 70 72 74 76 78 80

1982–86

Cryptodome

May 18, 1980 (plinian)

1980–81

Mar 1982

May 1982

Aug 1982

Feb 1983

Jun 1984

May 1985

Oct 1986

1982 (gm)

0

1

2

3

4

5

6

7

66 68 70 72 74 76 78 80 82

Current

Cryptodome

May 18, 1980 (plinian)

1980–81

SH304-2A

SH305-1

SH315-4

H2O, IN WEIGHT PERCENT H2O, IN WEIGHT PERCENT H2O, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT

SiO2, IN WEIGHT PERCENT

A

B

C

EXPLANATION

EXPLANATION

ian phase (cryptodome) of May 18, 1980, and during the subsequent dome-forming eruptions. This variation in melt-inclusion composition can be attributed to rapid extraction of the 

plinian magma from relatively deep in the subvolcanic plumbing system, whereas the cryptodome and post-May 18 domes 

were derived from magma that either ascended more slowly 

from the same depths as the plinian magma or had stalled at 

shallower levels (lower pressure) before extrusion (Blundy 

and Cashman, 2005). All subplinian eruptions of 1980 contain 

melt inclusions that show some affinity with those of the May 

18 plinian eruption in having elevated H2O and low SiO2.

Superimposed upon the trend of decreasing H2O with 

increasing SiO2 is a vertical trend of decreasing H2O at nearly 

constant SiO2 (shaded oval in fig. 5A). This trend is seen in 

some plagioclase-hosted melt inclusions from the plinian 

eruption of May 18 and the June 12 and October 16 domes (all 

1980). The trend is also evident in many amphibole-hosted 

inclusions and a single orthopyroxene-hosted inclusion. The 

vertical trend terminates in the groundmass glasses of the 

plinian pumices, which contain 0.5 to 2.2 percent H2O; the 

pumice analyses represent the H2O content of the melt at or 

close to the point of fragmentation during eruption. As magma 

ascent during the plinian eruption was extremely rapid, there 

was insufficient time for full chemical exchange between 

phenocrysts and matrix; thus the groundmass glasses had sufficient time to lose H2O syneruptively but not enough time to 

modify their SiO2 (Blundy and Cashman, 2005). We consider 

that some plagioclase, amphibole, and orthopyroxene crystals 

fractured during ascent, allowing H2O to escape without concomitant crystallization, such that they are displaced towards 

the groundmass glasses in figure 5A. This process primarily 

affected explosively erupted samples. In subsequent discussions we will refer, informally, to those inclusions lying within 

the shaded region of figure 5A as “ruptured inclusions.”

Melt inclusions from the 1982–86 phase of the eruption 

(fig. 5B) lie at the low-H2O, high-SiO2 end of the 1980–81 

trend. Only the March 19, 1982, sample contains melt inclusions with >2 percent H2O. The 1982–86 phases of the eruption 

therefore appear to derive largely from slowly ascending magma 

stored preeruptively at shallow levels, with the possible exception of March 19, 1982, where there is seismic evidence for 

involvement of new deeper magma (Malone and others, 1983).

Melt inclusions from the current eruption (fig. 5C) show 

contrasting behavior to those of 1980–86. All but two of the 

Figure 5. Variation in dissolved H2O versus SiO2 (normalized 

to 100 percent anhydrous) for melt inclusion and groundmass 

glasses, distinguished on basis of host mineral and eruption age. 

A, 1980–81. Gray oval denotes field of groundmass glasses and 

ruptured melt inclusions, thought to have lost H2O syneruptively 

(Blundy and Cashman, 2005). B, 1982–86. C, Current eruption 

(2004–2006). The 1980–81 variation is reprised in panels B and C

for comparative purposes.
770 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

inclusions have negligible H2O contents, close to ion microprobe detection limits. The variable SiO2 contents of these 

inclusions strongly suggest that they fall into the ruptured 

inclusion category. The striking exceptions are three inclusions 

from SH304-2A with elevated H2O (~3 percent) and SiO2 (~81 

percent). These inclusions lie quite remote from the 1980–86 

trend, suggesting that different processes have shaped their 

chemistry. The same inclusions have the anomalous REE patterns seen in figure 3C. We noted previously that there is evidence for significant silica-phase fractionation in glasses from 

the current eruption (for example, fig. 1E). The elevated H2O 

content of the inclusions from SH304-2A suggests that they 

encountered the silica-saturation surface at higher pressures 

than those of the 1980–86 eruption. Higher pressure saturation 

with silica can occur if the magma is initially more silica rich, 

cooler, or both (Blundy and Cashman, 2001, fig. 2).

Partial Pressure of H2O and Other Volatile 

Species

Changes in inclusion-sealing conditions can be conveniently expressed in terms of the partial pressure of H2O (pH2

O) 

(Blundy and Cashman, 2005). The pH2

O can be calculated 

from the measured H2O content and well-known solubilitypressure relations. We have used the software VolatileCalc 

of Newman and Lowenstern (2002) for rhyolitic melts at a 

nominal temperature of 900°C. Pressures calculated in this 

way are reported in table 2. The pH2

O varies from 248 MPa (for 

the maximum H2O content of 6.7 percent) to zero. The pH2

O

values calculated for pumice erupted during the May 18, 1980, 

plinian event are consistently higher than those of subsequent 

explosive and effusive eruptive episodes.

Partial pressure of H2O does not equate directly to total 

pressure (Ptot) because of the presence of other volatiles, notably CO2, halogens, and sulfur species. We can use the measured concentrations of F, Cl, and S (as SO2) in glasses (table 

2) to estimate partial pressures of other volatile species.

The chlorine content of glasses in the 1980–86 eruptive 

products ranges from 0.05 to 0.30 weight percent (fig. 6A) and 

shows a slight increase with decreasing pH2

O. This variation is 

consistent with the known negative pressure dependence of 

Cl solubility in silicate melts (Metrich and Rutherford, 1992). 

Enrichment of melt in Cl at low pressure can result either 

from crystallization or from interaction between Cl-rich vapor 

liberated from magma at higher pressures and magma stored at 

lower pressure, or a combination of both processes. The higher 

values measured are close to the measured solubility of Cl in 

hydrous rhyolite melts in equilibrium with a NaCl-KCl-H2O 

vapor (Shinohara and others, 1989). Metrich and Rutherford 

(1992) report slightly higher values, probably attributable to 

subtle differences in rhyolite starting composition. The fact 

that the measured solubilities lie at the upper end of the measured Cl contents strongly suggests that the 1980–86 Mount 

St. Helens melts were at or close to saturation with Cl-bearing 

vapor. However, this does not imply high pCl2

 (or pHCl2) in 

the vapor phase. For example, the Cl-saturated 60–120-MPa 

experiments of Shinohara and others (1989) involve fluids 

with Cl molalities as low as 1, which equates to a mole fraction of Cl in an H2O-rich vapor of less than 2 percent. Thus the 

contribution of pCl2 (or pHCl) to Ptot is negligible. We note that at 

pressures of ≤120 MPa the experiments of Shinohara and others (1989) show evidence for exsolution of a dense brine phase 

from the vapor. Thus it is possible that the Mount St. Helens 

melt inclusions from 1980–86 trapped at pH2

O <120 MPa were 

also brine saturated, which has implications for the chemical 

signature of fumarole gases released at the surface. Chlorine 

contents of two glasses from the current eruption plot at the 

lower end of the 1980–86 range (≤0.06 percent; fig. 6A).

The fluorine content of the glasses ranges from the 

EMPA detection limit of ~0.1 weight percent up to 0.5 weight 

percent (fig. 6B). No F was detected in the two melt inclusions 

from the current eruption with elevated H2O contents. There is 

considerable scatter in the data, largely a consequence of low 

count rates for F at the analytical conditions. Overall, however, 

F contents are similar to Cl in any given melt inclusion and 

show a similar slight tendency for increase in F with decreasing pH2

O. Fluorine contents are below the experimentally 

determined fluorite solubility of Dolejš and Baker (2006) and 

Price and others (1999).

Concentrations of SO2 (fig. 6C) are typically less than 

0.1 percent, with only two melt inclusions having significantly 

higher concentrations. Surprisingly, the highest SO2 concentrations are observed at lower pressures, and mostly in post-May 

18 samples. The highest values of SO2 lie at pH2

O of 50–150 

MPa, suggesting that S may be concentrated in the upper 

reaches of the subvolcanic plumbing system. The solubility 

and speciation of S in silicate melts depend on pressure, temperature, fO2

, fS2, and melt composition (see, for example, Carroll and Rutherford, 1985; Luhr, 1990; Clemente and others, 

2004; Scaillet and Pichavant, 2005). For the May 18 plinian 

eruption of Mount St. Helens, Whitney (1984) calculated fH2

S 

= 3.3 MPa and fSO2 = 2.0 MPa, indicating that the contribution 

of sulfur species to the overall fluid pressure is small. Sulfur 

dioxide contents of melt inclusions from the current eruption 

are similar to the lower values from 1980–86.

Concentrations of CO2 in melt inclusions were not measured routinely as part of this study. Even small amounts of 

dissolved CO2 equate to significant partial pressures, however, 

because of its low solubility in rhyolitic melts. Thus knowing CO2 concentrations is a prerequisite for converting pH2

O

to Ptot and thence to depth. Many of the characteristics of the 

1980–86 eruption are consistent with volatile-saturated dacite 

magma (Rutherford and others, 1985), but the coexisting 

vapor need not be pure H2O. Recently we have measured CO2

concentrations in bubble-free melt inclusions from the May 

18 plinian eruption and the August 7, 1980, eruption using 

ion microprobe analysis of 12C calibrated against a working 

curve consisting of rhyolite and andesite glasses of known 

CO2 content. The data and analytical method will be described 

in detail elsewhere. However, we note here that the analyzed 

melt inclusions do contain detectable CO2. The most H2O-rich 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 771

0

50

100

150

200

250

0 0.1 0.2 0.3 0.4

0

50

100

150

200

250

0 0.1 0.2 0.3 0.4 0.5

A B

PARTIAL PRESSURE H2O (pH2O),

IN MEGAPASCALS

Cl, IN WEIGHT PERCENT

PARTIAL PRESSURE H2O (pH2O),

IN MEGAPASCALS

F, IN WEIGHT PERCENT

1980 cryptodome

May 18, 1980 (plinian)

Post-May 18, 1980

2004–2005 dome

Experimental data Melt inclusions

EXPLANATION

Metrich and Rutherford,

 1992 

Shinohara and others,

 1989 Cryptodome

May 18, 1980 (plinian)

post-May 18, 1980

EXPLANATION

800°C (Dolejš and Baker,

 2006) 

850°C (Price and others,

 1999)

Experimental data Melt inclusions

2004–2005 dome

0

50

100

150

200

250

0 0.1 0.2 0.3 0.4 0.5

1980 cryptodome

May 18, 1980 (plinian)

Post-May 18, 1980

2004–2005 dome

EXPLANATION

Experimental data

Melt inclusions

Po sol. (900°C, FMQ)

Anhyd. sol. (900°C, MnO)

Po-Anhyd (900°C, NNO)

C

PARTIAL PRESSURE H2O (p

H2O), IN MEGAPASCALS

SO2, IN WEIGHT PERCENT

Figure 6. Concentration of volatile elements in 

melt inclusions as a function of pH2

O, as calculated 

from H2O contents using VolatileCalc (Newman and 

Lowenstern, 2002) at 900°C. Data are subdivided on 

basis of eruption age. Also plotted are experimental 

solubility data for comparison. A, Chlorine. 

Experimentally determined Cl solubility in system 

rhyolite-vapor±brine from Metrich and Rutherford 

(1992) at 830–850°C and Shinohara (1989) at 810°C. 

B, Fluorine. Experimentally determined fluorite 

solubility from Dolejš and Baker (2006) and Price 

and others (1999). C, Sulfur, expressed as SO2. 

Experimentally determined pyrrhotite (Po) and 

anhydrite (Anhyd) solubility at FMQ, NNO, and MnO 

buffers (Luhr, 1990).

(6.1 percent H2O) melt inclusion of those analyzed for CO2, 

from May 18, contains 400 ppm CO2. This equates to Ptot of 

281 MPa, using the model of Newman and Lowenstern (2002) 

at 900°C, compared to 216 MPa if only the H2O content is 

considered. Thus the presence of CO2 contributes 65 MPa to 

the calculated Ptot. Samples with lower H2O have lower CO2. 

For example, at 4 percent H2O, the maximum CO2 content is 

160 ppm, which equates to Ptot of 140 MPa, compared to 113 

MPa for the CO2-free case. The August 7, 1980, melt inclusions lie at the low H2O extrapolation of the May 18 samples 

and extend down to 1.6 percent H2O and 60 ppm CO2, where 

the incorporation of CO2 increases Ptot from 23 to 31 MPa. We 

do not yet have any CO2 measurements for melt inclusions in 

samples of the current eruption.

These preliminary results confirm the conclusions of 

Rutherford and others (1985) and Rutherford and Devine 

(1988) that the fluid phase in equilibrium with the May 18, 

1980, magma was a mixed H2O-CO2-SO2-H2S fluid. On the 

basis of experimental determination of phase relations and 

composition, they proposed a preeruptive equilibration pressure 

(Ptot) for the May 18 plinian magma of ~220 MPa with XH2

O in 

the fluid of 0.67. Our most H2O- and CO2-rich melt-inclusion 

analysis from this eruption corresponds to Ptot=281 MPa and 

XH2

O≈0.8. As pressure decreases, our preliminary data show 

that XH2

O remains approximately constant, suggestive of nearly 

closed-system degassing with high bubble fractions (Newman 

and Lowenstern, 2002) or gas streaming through the magma 

system, buffering XH2O (Rust and others, 2004).
772 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Without CO2 data for all of the analyzed melt inclusions, 

it is impossible to calculate Ptot for each one. However, we 

can use the preliminary results for CO2 to bracket the correction required to convert pH2

O to Ptot. We will assume that 

XH2

O = 0.8 for all melt inclusions and use the model of Newman and Lowenstern (2002) to derive a relation between pH2O

(calculated CO2 free) and Ptot (calculated at measured H2O and 

XH2

O = 0.8). This simple procedure gives the following 

empirical correction for the presence of CO2:

Ptot = 1.287 × pH2

O . (1)

The maximum Ptot for any melt inclusion in table 2, 

calculated in this way, is 248×1.287 = 319 MPa. In the following discussion we will use both pH2

O and Ptot (calculated from 

equation 1) as the ordinate axes for plotting.

Lithium

One of the most striking geochemical features of Mount 

St. Helens is the considerable variability in lithium (Li) 

content of melt inclusions and plagioclase phenocrysts (Berlo 

and others, 2004; Kent and others, 2007). For the 1980–86 

eruption, Li in melt inclusions ranges from 20 to 100 ppm 

(Berlo and others, 2004). New data for the 1982–86 phase of 

that eruption extends the range to 270 ppm (table 3). There 

is a corresponding elevation of Li contents in plagioclase 

phenocrysts from the same samples that show elevated meltinclusion Li (Berlo and others, 2004; Kent and others, 2007; 

Kent and others, this volume, chap. 35). For the 1980 samples 

(fig. 7A), Li reaches its maximum level at Ptot≈140 MPa. The 

high Li concentrations are confined to the cryptodome and 

post-May 18 eruptive phases. At lower Ptot, Li contents fall 

systematically to ~25 ppm, the same concentration observed in 

quenched matrix glasses from these samples. Melt inclusions 

from the plinian phase of May 18 have near-constant 30 ppm 

Li over a wide range of Ptot. Quenched matrix glasses from this 

phase of the eruption also have 25–30 ppm Li.

Berlo and others (2004) ascribed the behavior of Li in the 

post-May 18, 1980, to 1981 phases of the eruption to transfer 

of Li from deeper parts of the magma system to magma stored 

at shallower level. Enrichment of Li in cryptodome samples 

suggests that a similar pattern of Li transfer preceded the May 

18 eruption. A vapor phase was considered the most likely 

transport agent. The fact that Li enrichment reaches a maximum at a value of Ptot that marks the upper limit of the range 

Figure 7. Lithium variation as a function of pH2O and Ptot in meltinclusion and groundmass glasses. Values for Ptot calculated from 

pH2O using equation 1. A, 1980–81. B, 1982–88. C, Current (2004–

2006) eruption. The 1980–81 variation is reprised in panels B and C 

for comparative purposes. Horizontal scale differs in each panel.
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 773

defined by melt inclusions from the May 18 plinian samples 

that have not suffered syneruptive H2O loss (fig. 7A) suggests 

that Li enrichment takes place at the top of the main magma 

reservoir. May 18 melt inclusions with > 4 percent H2O (and 

two melt inclusions from June 12 and October 16, 1980) lack 

significant Li enrichment. Berlo and others (2004) argued that 

the longer a magma batch is stored at low pressure (Ptot ≤150 

MPa) before eruption, the greater the extent of Li enrichment 

observed. As magma continues to ascend above the level of 

Li enrichment, further degassing occurs, stripping the highly 

vesicular groundmass melt of its elevated Li but preserving 

high Li in melt inclusions and plagioclase.

New data for the 1982–86 phase of eruption (fig. 7B) 

paint a slightly different picture. A single melt inclusion from 

the March 1982 eruption has elevated H2O (pH2

O=60 MPa) and 

a Li content of 60 ppm, similar to melt inclusions in samples 

produced by the explosive-effusive activity of 1980. All subsequent eruptions, however, show Li enrichment at significantly 

lower pressures; the maximum Li concentration (276 ppm) 

occurs at Ptot=40 MPa. These data suggest that the locus of 

maximum Li enrichment shifted to lower pressures after 1981. 

The greater extent of Li enrichment in these inclusions is consistent with a longer period of shallow magma storage between 

eruptions. Groundmass Li from these samples has much lower 

Li, about 35–40 ppm.

The current eruption also shows elevated Li, as high as 

210 ppm (table 3; Kent and others, 2007). The high levels of 

Li enrichment in the current eruption probably reflect the long 

dormancy between 1986 and 2004. Most of the melt inclusions 

from the current eruption have experienced both syneruptive 

H2O loss (fig. 5C) and extensive shallow crystallization, such 

that the pH2

O recorded by the most Li-rich inclusions may be 

meaningless. However, three melt inclusions from SH304-2A 

have high Li and high H2O (pH2

O=80 MPa; fig. 7C). They 

therefore plot near the pressure level of maximum Li enrichment of the 1980 samples (pH2

O=110). A corresponding elevation of Li in plagioclase is seen in the same samples (Kent and 

others, 2007). These observations suggest that the locus of Li 

enrichment for the current eruption may have shifted back to 

higher pressures, although the data are too sparse to determine 

whether this is the same depth as that during the 1980–81 eruptive phase.

Kent and others (2007) propose that the mechanism of 

Li enrichment at Mount St. Helens involves upwards streaming of volatiles derived from deeper parts of the subvolcanic 

system, combined with condensation of a magmatic brine and 

loss of a low-salinity vapor phase at shallow levels. Lithium 

will be concentrated in the brine relative to the vapor, thereby 

providing a means of enriching the shallow-stored magmas in 

Li. This interpretation is consistent with the observed variation in Cl, which increases in concentration with decreasing 

Ptot (fig. 6A). Fractional degassing alone could not produce 

this enrichment. Concentrations of Cl in melt inclusions from 

the 1980–86 eruption are consistent, within error, with the 

experimentally measured solubility of Cl in brine-saturated 

rhyolite at pressures below ~160 MPa (Shinohara and others, 

1989). The discrete levels at which Li enrichment is observed 

at Mount St. Helens would then correspond not only to depths 

at which gas streaming was most intense (for example, at the 

top of the magma chamber), but also to depths at which phase 

separation occurs. The latter depth would be controlled by the 

NaCl content of the vapor (for example, Heinrich and others, 

1999). The proposed mechanism is analogous to that often 

invoked for the origin of hydrothermal ore deposits (Shinohara, 1994; Heinrich and others, 1999; Webster, 2004).

Crystallinity

We can use trace-element data from the melt inclusions to 

investigate the proposal that Li enrichment is associated with 

crystal-rich magmas at the top of the main magma reservoir. 

Incompatible elements, such as Cs and Rb, show increases 

in concentration with SiO2 (figs. 2B, 2C) that are most easily 

related to increases in crystallinity. As the 1980–86 eruption 

produced silicic andesite and dacite of near-constant bulk 

composition, it is possible to use the average bulk-rock trace 

element content to calculate the crystallinity of the magma at 

the time of trapping for each melt inclusion and then see how 

this varies with pressure (for example, Blundy and others, 

2006). For this exercise we have used Rb, because more melt 

inclusions have been analyzed for this element than for Cs and 

because Rb has a lower bulk partition coefficient than K. The 

average whole-rock Rb content of all magmas analyzed for the 

1980–86 eruption (n=5) is 31±3 ppm. We have used this value 

to calculate the crystallinity at the time of trapping for each 

melt inclusion assuming that bulk DRb = 0. The exact value of 

whole-rock Rb used is not crucial to these calculations. They 

provide a relative means of assessing crystallinity, provided 

that there are no significant secular variations in Rb from 

magma batch to magma batch. The typical 1σ uncertainty 

on each calculated crystallinity, based on propagation of the 

uncertainty in bulk Rb, is ±0.09.

Calculated crystallinities are plotted against pressure in 

figure 8A for the 1980–81 eruptive phases and in figure 8B for 

1982–86. The data are scattered, indicating either that there 

were variations in the initial Rb content of the magma or that 

the magma did not follow a single decompression crystallization trajectory. This is not surprising given the likely complex geometry of the subvolcanic system, which will lead to 

spatial gradients in crystallinity that depend, for example, on 

the storage period and proximity of a particular plagioclase 

grain to conduit or chamber walls. Nonetheless, the 1980–81 

melt inclusions show two crude trends of increasing crystallinity, shown by arrows in figure 8A. The first is recorded by 

the plagioclase-hosted melt inclusions from the plinian phase 

of May 18 and shows an overall increase from 0.10 at Ptot = 

300 MPa to around 0.35 at 160 MPa. The latter value is in 

good agreement with the modal abundance of phenocrysts 

in samples of the microlite-free white pumice from May 18 

(on a vesicle-free weight basis; Kuntz and others, 1981). This 

suggests that a significant proportion of the phenocrysts grew 
774 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

during decompression from Ptot ≥300 to 160 MPa (Blundy and 

Cashman, 2005). Some plagioclase-hosted melt inclusions from 

the cryptodome and post-May 18 eruptions of 1980 plot at the 

lower pressure end of this trend at 130–210 MPa (fig. 8A).

The second trend is defined chiefly by microlite-rich 

cryptodome and post-May 18 samples and extends from a 

crystallinity of ~0.35 at Ptot ≈ 80 MPa to 0.55 at the lowest 

pressures. This increase in crystallinity is consistent with the 

growth of microlites in these samples, augmented possibly by 

overgrowth rims on phenocrysts (Cashman, 1992). The highest crystallinities are recorded by microlite-rich cryptodome 

samples. Between the two trends of increasing crystallinity, 

from Ptot of ~180 to ~100 MPa, there is no marked change in 

crystallinity. Interestingly, this pressure interval brackets the 

Ptot at which the maximum Li enrichment occurs in the same 

samples (fig. 7A).

Melt inclusions from 1982–86 (fig. 8B) overlap the second 

trend described above (Ptot≤40 MPa), with broadly similar maximum crystallinities at low Ptot. The only exception is one melt 

inclusion from March 1982, which records Ptot (78 MPa) and 

crystallinity (0.30) similar to the lowest values from the May 18 

plinian pumice. Interestingly, the March 19, 1982, eruption not 

only was preceded by deep earthquakes but also produced the 

only post-1980 explosive eruption (Malone and others, 1983; 

Weaver and others, 1983). These observations are consistent 

with recharge (and subsequent degassing) of the deeper parts of 

the system at that time (Weaver and others, 1983).

The crystallinity recorded by melt inclusions from the 

current eruption (fig. 8C) is less well constrained, owing both 

to paucity of data and the fact that all but one of the inclusions 

analyzed for heavy trace elements appear to have ruptured (fig. 

5C). We have again used Rb for the crystallinity calculations, 

but with a slightly different bulk Rb content based on an average of dacites from the current eruption (33±3 ppm; Pallister 

and others, this volume, chap. 30). The one unruptured melt 

inclusion (from SH304-2A; fig. 8C) shows similar crystallinity 

and pressure (pH2

O ≈ 70 MPa) to melt inclusions from postMay 18, 1980, and March 1982 samples (fig. 8B). Evidence 

was presented above that melt inclusions of the current eruption encountered the silica-saturation surface at higher pressures than the 1980–86 magma. The fact that the crystallinity 

is comparable to those of 1980–86 inclusions, many of which 

lack evidence for silica saturation, suggests that the reason for 

higher pressure silica saturation in the current magmas is a 

result of their higher initial SiO2 contents rather than cooling. Ruptured melt inclusions from the current eruption imply 

very high magma crystallinities (≥0.7) at the time of trapping, 

consistent with textural evidence (Cashman and others, this 

volume, chap. 19). The loss of H2O from these inclusions, 

however, precludes any constraint on the pressure at which this 

high crystallinity was reached.

Oxide Thermobarometry

In order to assess the relative importance of decompression and cooling in driving crystallization at Mount St. Helens, 

Figure 8. Crystallinity calculated from Rb concentration of 

melt-inclusion and groundmass glasses as a function of pH2

O and 

Ptot . Symbols as in figure 7. A, 1980–81. Arrows denote the two 

episodes of crystallization described in text. B, 1982–86. C, Current 

eruption (2004–2006).
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 775

we have analyzed touching pairs of oxides from the 1980–86 

and current eruptions (table 4). A more comprehensive discussion of the oxide data from the current eruption is presented by 

Pallister and others (this volume, chap. 30).

The majority of phenocryst temperatures from 1980–86 

span a range from 840 to 960°C; fO2 ranges from 0 to 0.8 

log units above the NNO buffer (fig. 9A). A single oxidized 

pair from August 1982, showing exsolution lamellae, lies at 

NNO+1.8. Four phenocryst pairs define lower temperatures: 

one pair each from June 1981 and the cryptodome, and two 

pairs from a spine sample erupted in 1985. These lower temperature oxides straddle the NNO buffer and overlap the range 

of fO2 of high-temperature fumarole gases from Mount St. 

Helens (Gerlach and Casadevall, 1986; fig. 9A). Some oxides 

from August 1982 and June 1984 show exsolution textures 

typical of slow cooling, as often observed in plutonic rocks. In 

detail (fig. 9B), individual eruptions show discrete fO2 variations, with a tendency for fO2

 to decrease (relative to NNO) 

progressively from the May 18 plinian eruption to subsequent 

subplinian and dome-forming eruptions. Thus, the highest 

fO2 phenocrysts (at NNO+0.8) come from the microlite-free, 

H2O-rich plinian phase of May 18, 1980, whereas subsequent 

eruptions of microlite-bearing, H2O-poor magmas define 

subparallel trends but displaced to slightly lower fO2 (fig. 9B). 

The most reduced samples lie at NNO and come from August 

7, 1980, and June 1981. Microlite oxides from 1980–86 (fig. 

9C) cover essentially the same spreads in temperature and fO2 

as the phenocrysts from post-May 18 plinian samples; that is, 

displaced to slightly lower fO2 relative to NNO.

The reduced nature of groundmass oxides relative to 

phenocrysts (fig. 9C) and the progressive reduction of the 

phenocrysts as magmas become progressively less H2O rich 

and more crystalline (fig. 9B) suggests that magma reduction occurs concomitantly with degassing and crystallization. 

This is consistent with the relatively reduced nature of the 

high-temperature fumarole gases (Gerlach and Casadevall, 

1986) relative to the phenocryst oxides from the May 18 

plinian phase (fig. 9A). The lower temperature (≤850°C) 

and fO2 of oxide pairs from several samples appear to have 

equilibrated with the fumarole gases at very low pressures. 

Reduction of the magma during degassing suggests that 

equilibria between exsolved gases and magma are important 

in constraining the redox state of the melt. For the case of 

sulfur-free, H2O- and Fe-bearing melt, Candela (1986) and 

Burgisser and Scaillet (2007) show that degassing results in 

oxidation of the melt. Our contrary observation that degas-

-18

-16

-14

-12

-10

Cryptodome

EXPLANATION

May 18, 1980

May 25, 1980

June 12, 1980

July 22, 1980

August 7, 1980

June 1981

August 1982

June 1984

May 1985

May 1986

SH315-1

fumarole

650 700 750 800 850 900 950 1,000

NNO

TEMPERATURE, IN DEGREES CELSIUS

TEMPERATURE, IN DEGREES CELSIUS

TEMPERATURE, IN DEGREES CELSIUS

-12.5

-12.0

-11.5

-11.0

-10.5

840 860 880 900 920 940 960

NNO

-13.0

-12.5

-12.0

-11.5

-11.0

-10.5

-10.0

Cryptodome

August 7, 1980

August 1982

June 1984

Current

Plinian phenocrysts

850 900 950 1,000

NNO

A

B

C

OXYGEN FUGACITY (fO2), IN LOG10 UNITS OXYGEN FUGACITY (fO2), IN LOG10 UNITS OXYGEN FUGACITY (fO2), IN LOG10 UNITS

Phenocrysts and

 fumaroles

Phenocrysts

Microlites

EXPLANATION

Figure 9. Temperature versus fO2 variations for selected 1980–86 

and current samples calculated from touching Fe-Ti oxide 

pairs. Data from table 4. A, Phenocrysts and fumarole gases 

(Gerlach and Casadevall, 1986). B, Detail of phenocrysts, showing 

temporal evolution of fO2

; lines show linear fits to oxides from 

selected eruptions. Explanation same as in A. C, Groundmass 

microlites, with May 18, 1980, plinian-deposit oxides from panel 

A for comparison. Nickel-nickel oxide (NNO) buffer (at 200 MPa) 

shown for reference.
776 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

sing leads to reduction of the melt is consistent with the presence of reduced sulfur (S2-) in the melt, and the degassing 

equilibrium (Whitney, 1984; Candela, 1986; Burgisser and 

Scaillet, 2007):

6Fe2O3(melt) + 3S2-(melt) = 12FeO(melt) + 3SO2(vap). 

The magnitude of the observed reduction (≤0.7 log units 

in fO2) is consistent with the thermodynamic calculations of 

Burgisser and Scaillet (2007) for degassing and ascent of 

sulfur-rich, relatively oxidized (NNO+1) silicic magma.

The process of reduction can also account for the 

observed increase in melt sulfur contents at lower pressures 

(fig. 6C). For example, the experimental data of Clemente and 

others (2004) show that a modest decrease in fO2 from NNO+1 

to NNO increases sulfur solubility in rhyolite melt by a factor 

of three.

Following Blundy and others (2006), we consider the 

phenocryst temperature spread in figure 9B to reflect real 

temperature variations within the magma body. The preservation of the spread indicates that there was insufficient time for 

all oxides to fully equilibrate to changing temperatures before 

eruption. Possible causes of the spread include injection of 

new magma from depth into the system, loss of heat to wall 

rocks during magma storage, and latent heat release during 

decompression crystallization. Blundy and others (2006) argue 

that the latter process, which amounts to ~2.5°C for each 1 

percent crystallized, dominates during isenthalpic magma 

ascent. This explains why microlite-bearing samples with 

the highest crystallinities tend to have the highest maximum 

temperatures. Interestingly, the highest temperatures, and 

largest range of temperatures, are found in a May 25, 1980, 

pumice sample (fig. 9B) that shows evidence of extensive, 

very shallow crystallization. Extensive crystallization, coupled 

with a short repose interval (one week), would limit the extent 

to which the shallow magma could thermally reequilibrate 

after the rapid crystallization pulse. The profound effect of 

late-stage heating indicates that the most reliable estimate of 

temperature in the deep magma body that fueled the May 18 

plinian event lies at the low temperature end of the spread, that 

is, ~870°C, in good agreement with the estimate (880±10°C) 

of Venezky and Rutherford (1999) but somewhat lower than 

earlier estimates by Melson and Hopson (1981) and Rutherford and others (1985).

Temperature estimates from touching phenocrysts in a 

single sample (SH315-4) of the current eruption are displaced to slightly lower temperatures (850–870°C) relative to 

the main 1980–86 trend as observed by Pallister and others 

(this volume, chap. 30). A curious feature of SH315-4 is the 

unusually high temperature (1,020°C) recorded by a single 

touching microlite pair (fig. 9C). Pallister and others (this 

volume, chap. 30) also present sparse high Fe-Ti-oxide temperatures (≤1,019°C) from other samples of the current eruption. These data raise the possibility of some recharge of the 

magma system by hotter magma at depth. Alternatively, they 

could result from significant latent-heat release as ascending 

magmas undergo rapid and considerable crystallization (Pallister and others, this volume, chap. 30).

Temporal Variations

In this section we track temporal trends in several parameters at Mount St. Helens since March 17, 1980, the approximate 

date on which the first seismic activity was observed (Endo and 

others, 1990), in order to investigate evolving magma-storage 

conditions beneath the volcano. The most useful parameters to 

examine in time series are: (1) the volatile elements H2O, SO2, 

and Cl (fig. 10), which monitor changes in degassing; (2) the 

major elements K2O, MgO, and TiO2 (fig. 11), which monitor 

variations in crystallinity and the arrival of new, less-evolved 

magma batches; and (3) temperature and fO2 (fig. 12), which 

monitor cooling (or heating) and redox conditions.

There is an overall drop in the maximum H2O content of 

melt inclusions with time, from a maximum of nearly 7 percent during the plinian phase of May 18, 1980, to almost zero 

in 1986 (fig. 10A). Maximum H2O content of groundmass 

glass decreases from 2.2 to 0 percent over the same period. In 

general, explosive eruptions are characterized by melt inclusions with high H2O contents, whereas effusive eruptions are 

characterized by low-H2O inclusions. This strongly suggests 

a relation between magma storage conditions and eruptive 

style. In detail, the decrease in maximum H2O content with 

time is nonlinear and appears to describe a crude sawtooth 

cyclic pattern in which abrupt peaks in H2O are followed by a 

steady fall. Although this may in part be an artifact of sample 

size, our attempts to analyze the most H2O-rich melt inclusions from each sample using SEM screening, as described 

above, suggest that this variability is real. Peaks are observed 

on May 18, 1980 (plinian eruption), March 19, 1982, and 

June 17, 1984. It is unclear whether the single H2O-rich melt 

inclusion from October 16, 1980, represents a peak or simply 

an inherited crystal from an earlier eruption. The three peaks 

delineate cycles in eruptive behavior. The first cycle started 

with the plinian eruption on May 18, after which a steady 

decrease in magma supply rate led to increasingly short 

explosive eruptions that finally changed to discrete effusive 

events (Scandone and Malone, 1985). The second cycle initiated with an explosive eruption on March 19, 1982, preceded 

and accompanied by deep earthquakes, after an unusually long 

repose interval of 5 months. This cycle then evolved to a year 

(February 1983–February 1984) of continuous slow magma 

effusion and endogenous dome growth. The third cycle began 

with renewal of discrete extrusive dome growth events in 

March 1984 (although the first sample we have of this cycle 

was collected in June 1984) and continued until the end of 

effusive activity in 1986. This cyclicity suggests that changes 

in eruptive activity in March 1982 and March 1984 may 

reflect new inputs of magma. There are too few melt-inclusion 

data from the current eruption to discern any trend.

Chlorine shows the opposite behavior with time to H2O, 

with an overall increase in Cl from May 18, 1980, to 1986 (fig. 

10B). This is consistent with the increase in solubility of Cl 

with decreasing pH2

O (for example, fig. 6A). As is the case for 

H2O, the change in Cl is nonlinear with time and appears to 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 777

MI in plag

EXPLANATION

MI in amph

MI in opx

MI in cpx

groundmass

DAYS SINCE MARCH 17, 1980

DAYS SINCE MARCH 17, 1980

DAYS SINCE MARCH 17, 1980

April 10, 1980May 18, 1980June 12, 1980Aug. 7, 1980Oct. 16, 1980Dec. 27, 1980June 18, 1981Mar. 19, 1982Feb. 7, 1983June 17, 1984Oct. 21, 1986Oct. 18, 2004

A

B

C

0

1

2

3

4

5

6

7

100 1,000 10,000

0

0.1

0.2

0.3

100 1,000 10,000

0

0.04

0.08

0.12

0.16

100 1,000 10,000

SO2, IN WEIGHT PERCENT H2O, IN WEIGHT PERCENT Cl, IN WEIGHT PERCENT

0

1

2

3

4

5

6

7

0

0.2

0.4

0.6

0.8

1.0

1.2

0

0.2

0.4

0.6

0.8

1.0

A

B

C

DAYS SINCE MARCH 17, 1980

DAYS SINCE MARCH 17, 1980

DAYS SINCE MARCH 17, 1980

100 1,000 10,000

100 1,000 10,000

100 1,000 10,000

TiO2, IN WEIGHT PERCENT K2O, IN WEIGHT PERCENT MgO, IN WEIGHT PERCENT

April 10, 1980May 18, 1980June 12, 1980Aug. 7, 1980Oct. 16, 1980Dec. 27, 1980June 18, 1981Mar. 19, 1982Feb. 7, 1983June 17, 1984Oct. 21, 1986Oct. 18, 2004

Figure 10. Temporal variation of volatile components in melt 

inclusion and groundmass glasses for 1980–86 and current 

(2004–2006) eruptions, plotted as a function of days since 

March 17, 1980 (log scale). Melt inclusions (MI) distinguished on 

basis of host mineral: plag, plagioclase; amph, amphibole; opx, 

orthopyroxene; cpx, clinopyroxene. Selected eruptive episodes 

are marked on the upper abscissa in panel A for reference. A, 

Water. B, Chlorine. C, Sulfur dioxide.

Figure 11. Temporal variation of major-element oxides 

(normalized to 100-percent anhydrous) in melt inclusion 

(MI) and groundmass glasses for the 1980–86 and current 

(2004–2006) eruptions, plotted as a function of days since 

March 17, 1980 (log scale). Symbols as in figure 10. A, K2O. B, 

MgO. C, TiO2. All elements analyzed by electron microprobe.
778 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 12. Temporal variation of (A) 

temperature and (B) log10fO2

 (relative 

to NNO) in touching oxide pairs 

for 1980–86 and current eruptions, 

plotted as a function of days since 

March 17, 1980 (log scale). Fe-Ti oxide 

pairs are distinguished on a textural 

basis; “inclusions” denotes two 

oxides included in a single silicate 

phenocryst. Data from table 4. Small 

red dots denote temperature and fO2

in high-temperature fumarole gases 

(Gerlach and Casadevall, 1986). 

Horizontal lines show best estimate of 

conditions in magma body erupted in 

plinian phase of May 18, 1980, based 

on touching phenocrysts. Also shown 

for comparison are current eruption 

data of Pallister and others (this 

volume, chap. 30).

show distinct peaks (for example, October 1980) and troughs 

(for example, June 1981). The increase in Cl with time is 

particularly marked during the 1980 events, with the highest values reached in October. The temporal trend for Cl in 

groundmass glass broadly mirrors that of the melt inclusions. 

Chlorine contents of melt inclusions from the current eruption 

are variable but generally displaced to values lower than those 

of the 1980–86 eruption.

Total sulfur (expressed as SO2) drops abruptly from May 

18 through December 1980 to 1986, with consistently low 

values thereafter (fig. 10C), except for a slight elevation in 

March 1982. As with H2O and Cl, the June 1981 melt inclusions are distinguished by low SO2 contents. The highest value 

recorded (0.37 weight percent) comes from a melt inclusion in 

an October 16, 1980, sample (table 2; not plotted). This unusually high value, which greatly exceeds the known solubility of 

sulfur in rhyolite melts (Clemente and others, 2004), may result 

from sample contamination by tiny sulfide grains, although 

none were evident in backscattered electron images. Like H2O, 

SO2 exhibits a secondary peak in 1984 with the resumption of 

episodic dome-building eruptions. Groundmass glasses also 

show diminishing sulfur with time. The SO2 contents of melt 

inclusions from the current eruption are comparable to those of 

post-1980 melt inclusions.

Another monitor, K2O, which like Rb serves as a proxy 

for crystallinity (Cashman, 1992; Blundy and Cashman, 2005; 

Cashman and McConnell, 2005), shows a steady increase 

with time, from ≤2 percent on May 18, 1980, to ≥5 percent in 

February 1983 and October 1986 (fig. 11A). Minimum K2O 

values increase during 1980–81, before decreasing slightly in 

March 1982 pumice. Minimum K2O then increases again, with 

maximum values achieved in May 1985 (see also Cashman 

1992; Geschwind and Rutherford, 1995). The maximum K2O 

content during 1982–86 is highly variable, reaching a maximum in a February 7, 1983, sample from a highly crystalline 

lava spine. Overall these temporal trends are consistent with 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 779

increasing microlite crystallization with time in response to 

decreasing H2O. Interestingly, the temporal trends for K2O 

and H2O (fig. 10A) are not closely correlated, a reflection of 

the varying rates of degassing and crystallization, also seen as 

scatter in the pressure-crystallinity plot (fig. 8). Groundmass 

glasses have similar K2O contents to melt inclusions. The current eruption shows highly variable K2O, spanning more than 

the total variation from 1980–86. A single melt inclusion in 

orthopyroxene from October 18, 2004, shows unusually low 

K2O, whereas several melt inclusions preserve K2O values that 

exceed anything seen in 1980–86. The wide variability in K2O 

for the current eruption is consistent with extrusion of magma 

of highly variable crystallinity (Cashman and others, this 

volume, chap. 19), although we cannot rule out involvement of 

melts with initially very different K2O contents.

The MgO contents of melt inclusions (fig. 11B) are 

inversely correlated with K2O but show considerably greater 

variability for any one eruptive phase. From May 18, 1980, 

to June 1981, the maximum MgO content drops off sharply 

from <1 percent to ~0.25 percent, consistent with increasing 

amounts of microlite crystallization. There is then a marked 

increase in maximum MgO during the period March 1982 

to February 1983, more evidence suggestive of the involvement of slightly less crystalline magma during this eruption. 

After 1983 the MgO content falls to low levels similar to 

the minimum MgO contents of the earlier effusive phases. 

Groundmass glasses lie consistently at the low end of the meltinclusion data. As with K2O, MgO in melt inclusions from the 

current eruption is extremely variable, spanning almost the 

entire range observed in 1980–86 (fig. 11B).

The behavior of TiO2 (fig. 11C) differs from that of MgO. 

During 1980, minimum TiO2 contents increase slightly with 

time. After 1980 the TiO2 baseline remains almost constant 

through 1986, but the maximum TiO2 peaks between 1982 

and 1983. Groundmass glasses lie consistently at the low end 

of the melt-inclusion range. Melt inclusions from the current 

eruption are also variable in TiO2 but generally displaced to 

lower levels than 1980–86.

Iron-titanium oxide temperatures (fig. 12A) show a span 

of values throughout the 1980–86 eruption, with a slight 

(~50°C) overall decrease in maximum temperature with time, 

suggestive of some secular cooling within the subvolcanic 

magma reservoir. Oxide inclusions in phenocrysts from the 

May 18, 1980, plinian eruption give temperatures of 875°C. 

Phenocryst and groundmass oxide pairs from subsequent 

eruptions in 1980 are displaced to higher temperatures, which 

Blundy and others (2006) attribute to the latent-heat release 

during crystallization of phenocryst rims and microlites. 

Several phenocryst oxide pairs from post-May 18 samples 

show significantly lower temperatures that decrease with 

time. This trend matches that of high-temperature fumarole 

gases from Mount St. Helens collected in 1980–81 (Gerlach 

and Casadevall, 1986; fig. 12A), suggesting that these lowtemperature oxides equilibrated with fumarole gases as they 

streamed though the shallow magma-storage system. Over the 

same period fO2

 (relative to the NNO buffer; fig. 12B) shows 

fluctuations of ~1 log unit, decreasing during 1980, showing a marked increase in March 1982, and then decreasing 

again to 1986. For the 1980–81 period, the fO2

 values of the 

most reduced oxide pairs match those of the high-temperature 

fumarole gases (Gerlach and Casadevall, 1986; fig. 12B), also 

suggestive of equilibration of selected oxide pairs with vapors 

exsolved from underlying magma.

Iron-titanium oxide temperatures for the current eruption (fig. 12A) span a considerable range (730–1,020°C). The 

lowest temperatures are close to those of the coolest oxides 

recorded in 1986. The trend to temperatures ≥960°C may be 

caused by the addition of significantly hotter new magma or 

latent-heat release during crystallization (Pallister and others, 

this volume, chap. 30). In either case it is likely that the lowest 

oxide temperatures correspond to magmas held over from the 

end of the 1986 eruptive phase and that the heating event was 

sufficiently recent that no wholesale resetting of phenocryst 

oxide temperatures was possible. Oxygen fugacity estimates 

for the current eruption (fig. 12B) also cover almost the entire 

1980–86 range. This may result from extensive redox reactions involving fumarole gases or from addition of more 

oxidized magmas into the system.

Relation Between Petrological and 

Monitoring Data

In principle there should be a relation between the subterranean magmatic record, as preserved in melt inclusions, and 

the monitoring record, in the form of seismic events, ground 

deformation, volatile flux, and so forth. A relation has already 

been demonstrated for Fe-Ti oxide temperature-fO2 measurements and fumarole gases (fig. 12). In this section we will 

explore two further correlations, between the SO2 content 

of melt inclusions and the SO2 flux from the volcano, and 

between the pressures recorded by melt inclusions and the 

depths of subvolcanic earthquakes.

In figure 13A we compare the SO2 content of melt inclusions with the measured SO2 flux (McGee and Casadevall, 

1994). Sulfur dioxide degasses both during and between eruptive events, with some of the highest fluxes unassociated with 

the appearance of magma at the surface. A striking correlation 

is apparent between the two datasets. The high SO2 content 

of melt inclusions during May to August 1980 corresponds 

to high SO2 fluxes over the same period. The melt-inclusion 

data show a slight decrease in SO2 from October 1980 to June 

1981, before increasing again slightly through March 1982 

to June 1984. It is difficult to correlate this increase with the 

slight increase in SO2 flux over this period because of the 

sparseness of data. The clear temporal correlation between 

SO2 in melt inclusions and the SO2 flux suggests that magma 

degassing and the sealing off of melt inclusions are closely 

associated in time. Indeed it is possible that gas loss (which 

drives the SO2 flux) and sealing of melt inclusions (which 

traps SO2) are related processes. Significantly, SO2 degassing 
780 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

between eruptions produces spikes in SO2 flux which are not 

matched in the melt-inclusion data.

For those eruptive phases in which the total erupted 

volume and duration are known, it is possible to calculate a 

mean SO2 flux from the melt-inclusion data for comparison 

with the monitoring data. To perform the calculations we have 

taken the erupted volumes and eruption durations from Swanson and others (1987) and Swanson and Holcomb (1990), 

assumed a magmatic density of 2,500 kg/m3

, and used the 

maximum measured SO2 in melt inclusions for each eruption 

(that is, assuming negligible SO2 in the groundmass glass). 

The calculations for October 16, 1980, using the measured 

maximum of 0.37 weight percent, yielded anomalously high 

fluxes, confirming our suspicion that this is an anomalously 

high value and not representative of the eruption as a whole. 

We have therefore adopted the second highest value (0.04 percent) for this eruption. To calculate the average measured SO2

flux for each eruptive phase, we have taken the data of McGee 

and Casadevall (1994) and averaged them over the duration of 

the phase. The standard deviation is also calculated. In figure 

13B we plot the observed SO2 flux versus calculated (“petrological”) SO2 flux. The agreement between the two estimates 

is striking and within the combined errors on both sets of 

measurements. This is in contrast to calculated petrological 

fluxes for major plinian eruptions (Wallace, 2001), which 

consistently underestimate the measured fluxes by factors of 

up to 10. This suggests that plinian eruptions are driven in part 

by gas derived from deeply stored volatile-rich magma that 

is not erupted (Wallace, 2001). This interpretation is consistent with the SO2 flux between dome-forming eruptions: the 

good correlation between the calculated and observed eruptive fluxes during eruptions in figure 13B suggests that SO2

expelled between eruptions derives from a deeper source 

rather than from the magma that subsequently appears at the 

surface. Certainly there is evidence that syneruptive degassing 

processes differ between large plinian eruptions and smaller 

dome-forming eruptions.

The relation between the pressures recorded by melt 

inclusions and the seismic record for 1980–86 is equally 

significant. For each plagioclase-hosted melt inclusion we 

have calculated pH2

O and Ptot, using the methodology described 

above. We have not used melt inclusions hosted in mafic 

minerals, because of their tendency to lose H2O syneruptively. 

In addition, we have filtered the dataset for plagioclasehosted melt inclusions that appear to have ruptured (that is,

they plot within the gray ellipse in fig. 5A). The resultant 

dataset is plotted in figure 14A. There is a clear decrease in 

maximum pressure with time; a small amount of the decrease 

in pressure immediately after May 18, 1980, results from 

decompression of the entire system following removal of the 

upper 400–1,000 m of the edifice during the sector collapse 

that preceded the lateral blast (Moore and Albee, 1981). In 

detail, we note that each eruption contains melt inclusions 

with a pressure range (for example, Cashman and McConnell, 

2005), which diminishes with time. Magmas erupted during 

1980 cover a wide range in Ptot, from 320 to 0 MPa. There is 

a clear distinction between melt inclusions in the blast dacite, 

which tapped magma from the cryptodome and conduit over 

a pressure range of 0–160 MPa, and the main plinian phase, 

which tapped magma with melt inclusions stored from 120 

to 320 MPa. The minimum melt-inclusion pressure for the 

plinian magmas is slightly greater than the minimum pressure 

at which amphibole is stable at ~875°C, consistent with the 

Figure 13. Comparison of dissolved SO2 in melt inclusions and 

measured SO2 flux (in metric kilotons per day) for post-May 18, 

1980, episodes of 1980–86 eruption (McGee and Casadevall, 

1994). A, Temporal evolution of melt inclusion SO2 (left hand axis, 

red dots) and SO2 flux (right hand axis, black line). Key eruptive 

events shown on top axis for clarity. B, Observed SO2 flux versus 

calculated SO2 flux, based on duration of each eruptive phase, 

erupted volume, and maximum SO2 content in melt inclusions. 

Bars show 1 standard deviation error of observed flux. Error on 

calculated flux is ~10–20 percent relative. Full details of calculation 

procedure are given in text. Black line denotes a 1:1 correlation.
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 781

0

80

160

240

320

100 1,000

0

100

200

300

400

-3

0

3

6

9

12

15

100 1,000

DAYS SINCE MARCH 17, 1980

DEPTH BELOW SEA LEVEL,

IN KILOMETERS

DAYS SINCE MARCH 17, 1980

PARTIAL PRESSURE H2O (pH2O),

IN MEGAPASCALS

TOTAL PRESSURE (Ptot),

IN MEGAPASCALS

A

B

April 10, 1980May 18, 1980June 12, 1980Aug. 7, 1980Oct. 16, 1980Dec. 27, 1980June 18, 1981Mar. 19, 1982Feb. 7, 1983June 17, 1984Oct. 21, 1986

April 10, 1980May 18, 1980June 12, 1980Aug. 7, 1980Oct. 16, 1980Dec. 27, 1980June 18, 1981Mar. 19, 1982Feb. 7, 1983June 17, 1984Oct. 21, 1986

Figure 14. Evolution of melt-inclusion sealing pressures and 

depths with time for unruptured melt inclusions from 1980–86. 

A, Variation in pH2

O and Ptot as calculated from H2O contents and 

equation 1, respectively. May 18, 1980, cryptodome samples, 

erupted during blast phase and early plinian phase (Criswell, 

1987) are shown as open circles; all other eruptions, including 

main plinian phase of May 18, 1980, shown as dots. B, Comparison 

of melt-inclusion sealing depths (red dots), as calculated from 

equation 2a for eruptions of May 18, 1980, and earlier and 

from equation 2b for subsequent eruptions, with earthquake 

hypocenters (gray dots) from Pacific Northwest Seismic Network 

catalog, corrected by 1.1 km to account for mean seismic-station 

altitude. Note development of a shallow seismic lid before May 

18, 1980, and after December 27, 1980. Magmas erupted during the 

interval between these dates show consistently higher maximum 

extraction depths than those erupted before or after.

lack of amphibole breakdown rims in these samples (Rutherford and Hill, 1993). Eruptions after October 1980 produced 

magma with melt inclusions that record Ptot <50 MPa, with 

the exception of a single melt inclusion from March 1982 at 

80 MPa. All of these magmas contain amphibole with reaction rims (Rutherford and Hill, 1993). The clear impression 

from figure 14A is that the 1980 eruptions of Mount St. Helens 

sampled magma (as recorded by melt inclusions) stored over 

a wide pressure range, whereas the subsequent eruptions predominantly involved magma stored preeruptively at shallow 

levels. This change occurs abruptly between October 16 and 

December 27, 1980, and is accompanied by an abrupt change 

in eruptive activity from transitional explosive-effusive to 

predominantly effusive.

Earthquake depths for the period 1980–86 are available 

from the Pacific Northwest Seismic Network catalog. These 

data adopt as their datum the mean seismic station altitude at 

Mount St. Helens, which is ~1.1 km above sea level. In order 

to compare earthquake depths to Ptot, we require a relation 

between these two variables. We have integrated the depthdensity model for rocks beneath the volcano from Williams 

and others (1987), coupled with the volcano edifice-loading 

algorithm of Pinel and Jaupart (2000), to calculate lithostatic 

pressure as a function of depth. For the period before and 

including May 18, 1980, we assume a conical edifice 2.95 km 

high standing 1.5 km above the surrounding landscape and 

having a basal radius of 3.2 km. The edifice density is taken 

as 2,170 kg/m3 (as in Moran, 1994). The calculated pressures 

are then parameterized in terms of depth, z, (below sea level) 

using a polynomial:

 z (km) = 0.029654 Ptot+ 0.22704Ptot

0.5−2.95 , (2a)

where Ptot is the calculated total pressure, in MPa, for each 

melt inclusion using the relation between Ptot and pH2

O in equation 1.

Following the removal of the volcanic summit on May 

18, 1980, there is a small pressure readjustment in the rocks 

beneath the volcano. We have modeled this change using a 

truncated conical edifice with its top 2.55 km above sea level 

and the method of Pinel and Jaupart (2005) to derive a second 

expression for calculating depths in samples erupted after May 

18, 1980, as follows:

z (km) = 0.03074 Ptot+ 0.18334Ptot

0.5−2.55 . (2b)

Note that the pressure drop is small (≤6 MPa) and confined 

to the upper reaches of the subvolcanic system (above 500 m 

above sea level).

In applying equations 2a and 2b we are making a number of assumptions. First, we assume that the pressure in the 

magma reservoir is lithostatic, without significant magmatic 

overpressure through volatile buildup, for example, or significant regional deviatoric stresses. As magmatic overpressures 

are likely to be ≤20 MPa (Massol and Jaupart, 1999), we 

consider this effect to be small. Second, implicit in our use 

of equation 1 is the assumption that XH2

O is approximately 

constant at 0.8 throughout the magma column and that the 

partial pressures of other volatile species (such as SO2, H2S, 

HCl, HF) are very low.

We have calculated z for all unruptured plagioclasehosted melt inclusions, for comparison to the earthquake 

depths (fig. 14B), after subtracting 1.1 km from the latter 

to account for the mean station altitude. There is a striking overall consistency between the calculated ranges of 
782 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

melt-inclusion depths and the range of earthquake depths 

associated with any given phase of the eruption. This agreement is particularly strong for May 18, 1980, for which the 

maximum melt inclusion and earthquake depth ranges agree 

to within 1–2 km. Here we note that the deep earthquakes 

followed the explosive eruption and have been interpreted 

to reflect collapse in response to magma withdrawal (Scandone and others, 2006). Following May 18, 1980, there is a 

dramatic dropoff in earthquake event frequency throughout 

1980. Melt-inclusion depths over the same period extend 

down to ~10 km below sea level but are consistently shallower than those in the plinian deposit. Here again, deep 

seismic activity occurred in response to, rather than before, 

eruptive activity, an interpretation that is consistent with the 

observed correlation between earthquake and melt-inclusion 

depths. This pattern changed in March 1982, when deep 

earthquakes preceded the explosive eruption by about a 

week (Scandone and others, 2006). Here we see a mismatch 

between the earthquake depth and maximum recorded meltinclusion depth that suggests that the week of intrusion was 

sufficient for degassing and equilibration at intermediate 

storage levels before eruption.

After 1980 there is an increase in the number of shallow 

earthquakes, reflecting a gradual closing of the conduit system 

that remained open through the summer of 1980 (Scandone 

and others, 2006). During this period no melt inclusions record 

depths greater than 1.5 km below sea level. Apparently the 

development of extensive shallow-level seismicity in December 

1980 marks the formation of a barrier to the eruption and/or 

preservation of deeper melt inclusions. This may be a consequence of slower magma-ascent rates during this period, 

such that higher pressure melt inclusions are not preserved, or 

the preeruptive stalling of each magma batch at depths corresponding to the region of shallow seismicity. This scenario is 

consistent with effective sealing off of the subvolcanic system 

after December 27, 1980 (for example, Scandone and Malone, 

1985). Shallow seismicity persisted throughout the noneruptive period of 1986–2004, gradually deepening with time and 

leading eventually to the formation of a distinct “seismic lid” at 

depths of ~1 km below sea level after 1992 (Moran and others, 

this volume, chap. 2). Seismic activity associated with the current eruption has been consistently shallow (within the seismic 

lid). Melt inclusions from the current activity tell the same 

story; most record very low pressures, whereas the most H2Orich originate approximately at the base of the pre-October 

2004 seismic lid.

Discussion

The melt-inclusion data are consistent with a vapor-undersaturated silicic magma generated in the deep crust, ascending 

to the point of vapor saturation, and undergoing decompression 

crystallization (Blundy and Cashman, 2001, 2005; Annen and 

others, 2006). Melt inclusions from the 1980–86 samples record 

a vertical variation in crystallinity with pressure, such that the 

lowest pressure samples have the highest crystallinity and the 

lowest bulk magmatic H2O content. The plinian phase of May 

18, 1980, tapped the deeper, less crystalline part of the magma 

reservoir―from a relatively crystal-rich cap at Ptot=120 to a 

crystal-poor, deeper part at Ptot=320 MPa. The few plagioclaseand amphibole-hosted melt inclusions that record lower pressures lost H2O syneruptively (fig. 5A). Preplinian and postplinian 

eruptions tapped magma stored at somewhat shallower levels, 

which had undergone further decompression crystallization. The 

deepest, most crystal poor of these magmas came from Ptot of 

150–220 MPa during 1980, that is, overlapping the shallowest 

plinian magmas, and from much smaller Ptot during 1982–86. 

This picture of the subvolcanic magma system is consistent with 

seismic (Lees, 1992; Moran 1994), petrologic (Pallister and others, 1992; Rutherford and Hill, 1993), and geodetic (Mastin and 

others, this volume, chap. 22) data that suggest a thin, vertically 

extensive magma chamber less than 500 m in diameter at a depth 

of ~5 km below sea level, connected to the surface by a narrow 

conduit (≤40 m in radius) that may be only intermittently filled 

with magma (Scandone and others 2006). As magma ascends 

through this system, it loses H2O and crystallizes, leading to a 

vertical gradient in H2O content and crystallinity.

Within the magma chamber, crystallinity increases 

from ~10 to 30 weight percent. It is unclear whether 

10-percent crystallinity represents the crystal content of the 

magma as it enters the deep system or whether the magma is 

essentially aphyric at that point. A possibility that we favor 

is that the 10 percent initial crystal load was entrained from 

the deep-seated magma generation zone, as proposed by 

Annen and others (2006). Alternatively, rapid crystallization 

of initially aphyric magma at the point of vapor saturation 

may have prevented the preservation of any melt inclusions 

with compositions matching the bulk Mount St. Helens 

magmas. Magma from the deeper reservoir periodically fills 

the shallow plumbing system, undergoing further crystallization en route. This crystallization involves both rim 

overgrowths on phenocrysts and microlite nucleation. Total 

crystallinity reaches a maximum of ~60 weight percent at 

the shallowest levels.

Between eruptions, magmatic gases streaming out of 

the deeper part of the system interact with the shallow-stored 

magma, enriching the magma in Li at the point of phase 

separation to a brine. The same process of intereruption gas 

streaming enriches shallow-stored magmas in Cl and F (figs. 

6A, B) and 210Pb (Berlo and others, 2006), enriches plagioclase 

phenocrysts in Li (Berlo and others, 2004), generates pulses in 

the SO2 flux (fig. 13A), and may also buffer XH2O in the vapor 

phase. The observed 210Pb excesses in the same magmas that 

show Li enrichments (Berlo and others, 2004, 2006) suggest 

a subvolcanic magma system wherein a large deep reservoir 

episodically supplies 222Rn-bearing gas to a much smaller shallow reservoir.

The significance of melt-inclusion pressures (and corresponding depths) hinges on the mechanisms by which 

inclusions form, evolve, and become sealed. Melt inclusions 

trapped in plagioclase, amphibole, and pyroxenes show 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 783

similar chemical characteristics (figs. 1, 2), indicating that 

the compositions of the melt inclusions are not controlled by 

host-inclusion exchange. There is, nonetheless, clear evidence 

that melt-inclusion compositions have been modified since 

they first became incorporated. The fact that melt inclusions 

and groundmass glasses have overlapping compositions for all 

nonvolatile components suggests that the dominant process of 

melt-inclusion modification is exchange with the groundmass 

melt, rather than exclusively by reaction with the host mineral. 

We suggest that exchange of components between inclusion 

and groundmass melts, via thin channels or capillaries connecting inclusion and matrix, persists for some time after the 

inclusion is first trapped (Anderson, 1991; Stewart and Pearce, 

2004; Blundy and Cashman, 2005; Humphreys and others, 

2008). Thus, melt-inclusion chemistry follows groundmassliquid lines of descent, despite the absence of daughter minerals in the inclusion, up to the point that the inclusion becomes 

sealed off (occluded) from the matrix.

Occlusion may be a physical process, whereby a new 

overgrowth of plagioclase blocks the channel connecting 

inclusion and groundmass, or it may be a kinetic process, 

limited by the time available for chemical exchange. For a 

typical channel length of 100 μm (Blundy and others, 2006) 

and magma temperature of ~900°C, the diffusivity data of 

Baker (1991) indicate that diffusive exchange of Si (the slowest diffusing species) will occur on time scales of 7 to 46 

hours for melt H2O contents of 3 and 6 percent, respectively. 

Thus, even without physical occlusion, melt inclusions are 

unlikely to equilibrate chemically with the groundmass in the 

last day or so before eruption. Whether physically or kinetically occluded, there comes a point after which each melt 

inclusion undergoes no further chemical modification, except 

for syneruptive loss of H2O and other volatiles from ruptured 

inclusions. We refer to this as the “melt-inclusion sealing 

point,” which corresponds to a depth and time coordinate in 

the overall magma ascent trajectory. There may be a considerable time gap between the moment when the inclusion first 

exists as a physical entity within the crystal, the “melt-inclusion formation point,” and the sealing point. During this time 

gap the melt-inclusion chemistry evolves in harmony with the 

groundmass melt.

There are several lines of evidence from this study to 

suggest that the melt-inclusion sealing point corresponds 

closely in time to the onset of the event in which the host 

crystal is erupted. For example, the correlation between SO2

contents of melt inclusions and the SO2 flux associated with 

a particular eruptive event (fig. 13B) suggests sealing off 

shortly before eruption. If the inclusion were sealed several 

weeks or more before the eruption, the petrological sulfur 

flux and the measured sulfur flux would be out of phase. 

The correspondence between inclusion-sealing pressures and 

eruptive style (fig. 14A) suggests that rapid upwards movement of the magma may be sufficient to seal off the inclusion. Thus, the explosive eruptions of 1980 preserve melt 

inclusions sealed over a wide pressure range, because there 

was insufficient time to allow the inclusions to modify their 

chemistry during the final stages of ascent. Conversely, after 

1980, when eruptions were less vigorous and predominantly 

effusive, sealing depths are consistently shallow, reflecting 

accumulation of magma beneath the growing seismic lid 

between eruptions as well as slower ascent and more time to 

chemically exchange with the groundmass.

There is also evidence to suggest that the melt-inclusion 

sealing point corresponds closely in space to the region of the 

subvolcanic system where the host crystal was stored before 

eruption. The close temporal correspondence between sealing 

depths and earthquake depths (fig. 14B) suggests that there is a 

relation between magma withdrawal and earthquakes. During the period May 18 to December 27, 1980, sealing depths 

show a wide range, down to 11 km below sea level. During 

this time period, earthquakes occurring in response to evacuation of magma from the system cover the same depth range. 

At the same time, limited precursory seismic activity and high 

intereruptive gas fluxes show that the conduit system was 

relatively open (Scandone and others, 2006). After December 

27, shallow earthquakes increase dramatically in abundance to 

form a seismic lid, and sealing depths become almost exclusively shallow. The only post-December 1980 melt inclusion 

with a sealing depth below sea level is found in the explosive 

event of March 19, 1982, which punctuated nearly 5 years of 

predominantly effusive activity. This event was preceded by 

a deep earthquake swarm, apparently testifying to renewed 

magma supply from depth and tapping of a small part of this 

magma during the eruption. The dominant control over the 

transition in eruptive behavior after December 1980 is likely 

to have been dwindling magma supply from depth (Scandone 

and Malone, 1985), such that the addition of new magma to 

the base of the system was unable to maintain a fully connected conduit or to keep pace with the effects of decompression crystallization. The current (2004–2006) eruption appears 

to be a continuation of this process, with magma ascent now 

sufficiently slow that a plug of largely crystallized magma is 

being extruded as a spine from depths at which there are abundant shallow earthquakes (Iverson and others, 2006).

We propose that, for any given eruptive event, the melt 

inclusion sealing depths are a good indication of both the 

depth range over which magma was stored preeruptively 

and the rate at which it was extracted. Before May 18, 1980, 

magma resided from just below the volcanic summit to at 

least 11 km below sea level. The maximum Li enrichment 

observed at ~5 km below sea level probably marks a constriction at the top of the magma column. All magma at shallower 

levels was probably stored in a plexus of dikes that formed the 

conduit or within the growing cryptodome. Magma stored at 

very shallow depths in the nascent cryptodome was erupted 

as early as April 10, 1980 (Cashman and Hoblitt, 2004). The 

lateral blast of May 18 eviscerated the cryptodome and magma 

residing in the conduit down to ~5 km below sea level (see, 

for example, Scandone and others, 2006). The May 18 plinian 

event erupted a large volume of magma stored within the subvolcanic chamber from 5 to 11 km below sea level. Residual 

magma left at shallow levels was probably the source of much 
784 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

of the May 25 eruption. The subvolcanic system was then 

efficiently recharged, leading once again to a system in which 

a relatively deep magmatic reservoir supplied gas to shallow 

stored magma, giving rise to Li (and 210Pb) enrichments (Berlo 

and others, 2004, 2006). The post-May 18 transitional eruptions tapped magma stored at depths ≤10 km below sea level, 

probably as a result of evacuation of the vestigial conduit 

system developed after the plinian event (see also Cashman 

and McConnell, 2005). This was the last “deep” magma to be 

erupted from Mount St. Helens.

After December 27, 1980, all melt inclusions have sealing 

depths at or above sea level, a single exception being a melt 

inclusion from the March 19, 1982, explosive event. Coincident with the change in inclusion sealing depths, there is the 

development of extensive shallow seismicity and the change to 

dominantly effusive behavior. The depth at which maximum Li 

enrichment occurs is also displaced to sea level or shallower, 

indicating either a change in the composition of gas released 

from depth, such that the brine condensation pressure is 

reduced, or a change in magma storage depth. In either case it 

is clear that after October 1980 the subvolcanic magma system 

contained predominantly crystalline magma that was unable 

to sustain explosive activity. An earthquake swarm in March 

1982 may testify to limited recharge of the deeper system (≤ 9 

km below sea level) by new magma. However, very little of this 

deeper magma was erupted directly, most being stored instead 

at shallow level (<2 km below sea level) before eruption.

There are similarities between the post-December 1980 

preeruptive magma-storage conditions and those of the current 

eruption, where once again exclusively shallow-stored magma 

(≤2 km below sea level) is extruded. However, continuous 

slow magma extrusion during the current eruption results in a 

conduit plugged by largely crystalline magma, leading to successively erupted holocrystalline spines lubricated along their 

margins by well-developed fault gouge (Iverson and others, 

2006; Cashman and others, this volume, chap. 19; Iverson, this 

volume, chap. 21). We would expect activity to revert to a more 

explosive eruptive style only after the conduit has been completely cleared to liberate magma stored beneath. This would 

require a substantial increase in the magma supply rate from 

depth or destruction of the shallow volcanic plumbing system.

The decay of 222Rn to 210Pb and diffusion of Li in plagioclase provide important time constraints on the process 

of degassing and the storage and ascent of magma at shallow 

level. Berlo and others (2006) argue that the enrichment in 

210Pb occurs because significant decay of 222Rn occurs during 

gas transport from depth. Given the 3.8-day half-life of 222Rn, 

the 210Pb enrichment constrains gas transport time to be at least 

of this order. The rapid diffusion of Li in plagioclase (Giletti 

and Shanahan, 1997) indicates that diffusive enrichment (and 

loss) of Li on the scale of millimeters (a typical plagioclase 

phenocryst diameter) occurs on time scales of hours to days 

at 900°C. Thus magma must be stored at shallow levels for at 

least hours to days in order to acquire the observed Li enrichment. The magma must then ascend on a time scale that is long 

enough to allow Li to diffuse out of the melt and into the vapor 

phase but too short to allow the Li to escape from the plagioclase phenocrysts. This behavior is consistent with repose 

periods of a week or more between those eruptive events showing Li-enriched plagioclase (Berlo and others, 2004), and with 

ascent times of several days or less for post-18 May eruptions 

of Mount St. Helens (Rutherford and Hill, 1993).

The current eruption also shows very high Li concentrations in melt inclusions and plagioclase (Kent and others, 2007). Although the data are insufficient to discern the 

depth of Li enrichment, it is clear that substantial time is 

required to generate such high Li levels via the gas-streaming 

mechanism. Certainly it appears that the process of enrichment is similar in the current eruption to that of 1980–86 and 

this represents a characteristic feature of magma ascent and 

degassing at Mount St. Helens.

A key question about the current eruption is whether, 

petrologically, it represents a continuation of the 1980–86 

eruption, following a brief period of dormancy, or whether it 

heralds the arrival of new magma at depth. There are several 

lines of evidence presented elsewhere in this volume that support the involvement of some new magma, not present at the 

time of the 1980–86 eruption. First, there is seismic evidence 

for recharge of the deeper system as early as 1987–92 (Moran, 

1994). Focal mechanisms for these events are consistent with 

repressurization of the system at depths of 5–9 km below 

sea level, in the depth range of the 1980–86 magma chamber. Additional deep (8–10 km below sea level) earthquake 

swarms occurred in 1996 and 1998–99, suggesting continued 

recharge (Musumeci and others, 2002). Second, whole-rock 

geochemical data (Pallister and others, this volume, chap. 30) 

reveal that, although the bulk chemistry of the currently erupting dacite is slightly more evolved than that of 1980–86, the 

concentrations of some compatible trace elements (Ni, Cr) are 

higher, suggestive of addition of a more mafic input into the 

system. Third, Fe-Ti oxide temperatures (Pallister and others, 

this volume, chap. 30) indicate the appearance of high-temperature (>1,000°C) pairs in the groundmass of some recent 

samples, in accord with data presented in this work. No such 

high temperatures were evident in 1980–86. This is suggestive of heating by a new hotter magma at depth, although we 

cannot rule out the effects of significant latent heat release in 

these highly crystalline magmas.

The melt-inclusion data presented here also support the 

arrival of new magma at depth. The trace-element compositions of two melt inclusions from the current eruption are in 

marked contrast to all but one of the melt inclusions analyzed 

from 1980–86 in their steep REE pattern and lower U, Th, Y, 

Zr, and Hf. In terms of major elements, there are subtle differences between the current and 1980–86 eruptions. Although 

the melt inclusions of the current eruption are still rhyolitic 

in composition, the range of MgO and K2O contents in them 

almost equals the entire range seen in 1980–86, whereas TiO2

and Cl contents are uniformly lower. The simplest interpretation of these data is mixing of a new and different magma into 

a reservoir filled with the unerupted residue of the 1980–86 

eruption. At present the evidence for this mixing event is 
33. Evolving Magma Storage Conditions Beneath Mount St. Helens Inferred from Chemical Variations in Melt Inclusions 785

subtle, and it is inferred only through slight changes in traceelement chemistry consistent with small parcels of the new 

magma invading volumetrically dominant older magma stored 

at shallow levels. This interpretation is consistent with seismic 

evidence for such batches (for example, Scandone and others, 

2006). Over time, the addition of increasing quantities of this 

melt component displaces slightly the bulk chemical composition of the currently erupting magma, as evinced by figure 4. 

The melt-inclusion data do not necessarily require input of a 

more mafic magma into the system, just a more silicic magma 

with elevated Ni and Cr contents, similar, for example, to the 

dacites erupted during the Kalama period (Pallister and others, 

1992, and this volume, chap. 30). Continued monitoring of 

melt inclusions from the current eruption is required to track 

the progressive involvement of the new magma as the eruption 

taps progressively deeper levels of the subvolcanic reservoir.

Conclusions

Whole-rock chemical variations in magma erupted from 

or around Mount St. Helens during late Quaternary time 

testify to extensive chemical differentiation of basaltic magma 

beneath the volcano to produce silicic andesite and dacite, 

such as those erupted in 1980–86 and since 2004. Plagioclase 

involvement was minimal during the early stages of differentiation, probably because of elevated H2O contents in the 

parental magmas (Berlo and others, 2007). The compatible 

behavior of Y, Zr, and Nb during whole-rock differentiation 

points to the involvement of garnet, amphibole, and ilmenite 

in modal proportions at odds with the observed phenocryst 

assemblages. This suggests that differentiation occurred at 

considerable depth below Mount St. Helens, probably at the 

base of the crust (Ptot ≥1 GPa), where garnet becomes a stable 

crystallizing phase (see, for example, Müntener and others, 

2001). The curvature of many whole-rock geochemical trends 

rules out magma mixing as the primary source of chemical 

variation at Mount St. Helens. However, limited amounts of 

mixing of different magmas in the source region, and entrainment of crystal residues from depth, may account for some of 

the linearity in rocks with >62 percent SiO2, including those 

from the 1980–86 and current (2004–2006) eruptions.

Melt-inclusion and groundmass-glass compositions lie 

at the high-SiO2 end of the whole-rock arrays. The chemical 

trends exhibited by the glasses commonly are oblique to those 

of the whole rocks, indicative of a marked change in crystallizing assemblage. We interpret the glass trends to result from 

low-pressure crystallization of a dacite magma that was generated in the lower crust and ascended to the point of volatile saturation, ~11 km below sea level. Low-pressure crystallization 

involved the plagioclase-rich phenocryst assemblage, rather 

than the high-pressure amphibole-rich assemblage, which controls the whole-rock chemistry. The wide variety of matrix and 

inclusion glass compositions within host dacites of restricted 

composition indicates that low-pressure crystallization involved 

negligible physical segregation of crystals and melts.

Melt inclusions and Fe-Ti oxide data from the 1980–86 

eruption provide constraints on the evolving subvolcanic 

magma plumbing. The dissolved volatile contents of the melt 

inclusions have enabled us to determine the storage pressure of a magma batch just before its eruption. By converting 

pressure to depth, we have been able to track the changing 

magma-extraction depths with time. The plinian eruption of 

May 18, 1980, involved magma that last equilibrated 5–11 km 

below sea level. The preceding eruptions, including the May 

18, 1980, blast deposit, tapped magma from just below the 

edifice to ~5 km below sea level. Subsequent episodes in 1980 

tapped magma down to 10 km below sea level. The tapping of 

deeper magma stopped abruptly at the end of 1980, coincident with the onset of shallow preeruptive seismicity and the 

transition to purely effusive activity. All subsequent 1981–86 

eruptive episodes tapped exclusively shallow-stored magma. 

We interpret the development of shallow seismicity to mark 

the effective closure of the conduit system by stalled, highly 

crystallized magma in response to diminished magma supply. 

Therefore, 1980 corresponds to a period during which the conduit formed on May 18 remained sufficiently open that magma 

could be tapped over a wide range of depths.

The current eruption of Mount St. Helens shares similarities with the 1981–86 phase of the previous eruption, in 

that magma appears to be fed from shallow levels, beneath 

a seismic lid. However, the continuous slow effusion that 

characterizes the current activity is different from the more 

rapid episodic effusion of most of the 1980s and produces 

highly crystalline conduit material in the form of a sequence 

of collapsing spines. Several subtle petrological and chemical 

lines of evidence indicate that the deep magma driving spine 

extrusion differs from that previously present at Mount St. 

Helens. Melt-inclusion data suggest that the new magma has 

lower concentrations of REE, U, Th, and HFSE; higher Ni 

and Cr; and a steeper REE pattern than melts from 1980–86. 

To date, very little of this magma has insinuated its way into 

the choked conduit system. With time, however, we anticipate 

that the proportion of new magma in erupted products will 

increase. A careful integration of petrologic, seismologic, and 

gas-monitoring data will help to mark the progressive rise of 

any new magma through the system and, possibly, its eventual 

appearance at the surface.

Acknowledgments

This work was funded through research fellowships 

from the Natural Environment Research Council (U.K.) and 

the Royal Society to Blundy, National Science Foundation 

grants EAR-0207362 and EAR-0510437 to Cashman, and a 

University of Bristol Ph.D. studentship to Berlo. We would 

like to thank J. Craven, S. Kaseman, and R. Hinton for support on the Edinburgh ion microprobe, S. Kearns for help on 

the Bristol electron microprobe, and informative discussions 

with M. Humphreys, O. Melnik, S. Sparks, J. Pallister, and 
786 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

S. Moran. L. Mastin, R. Scandone, and S. Sparks provided 

thorough and helpful reviews. We are indebted to B. Chappell 

for painstaking compilation of the published whole-rock data 

from Mount St. Helens, S. Moran for providing seismic data, 

S. Carey for providing sample MSH006, D. Pyle for providing sample SH80D, and V. Pinel for providing details of the 

depth-pressure relation. This work is dedicated to the memory 

of Ben John (1935–2006).

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Appendix 1. Chemical Analyses of Glassy Melt Inclusions, Groundmass Glass, 

and Fe-Ti Oxide Pairs from the 1980–86 and 2004–2006 Eruptions of Mount St. 

Helens, Washington

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ] 

Appendix 1is a spreadsheet file that contains the entire dataset of major, trace-element, and volatile analyses, a selection of 

which is in tables 2 and 3 in the text. See tables for analytical details. 

Appendix 2. Chemical Analyses of Fe-Ti Oxide Pairs from the 1980–86 and 2004–

2006 Eruptions of Mount St. Helens, Washington

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume and as 

a separate file accompanying this chapter on the Web at: http://pubs.usgs.gov/pp/1750. ]

Appendix 2 is a spreadsheet file that contains the entire dataset of Fe-Ti analyses and calculated temperature and oxygen 

fugacity, a selection of which is in table 4 in the text. See table for analytical details.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 34

1

 Department of Geology, Portland State University, Portland, OR 97207

2

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

3

 U.S. Geological Survey, 345 Middlefield Road, Menlo Park, CA 94025

Plagioclase Populations and Zoning in Dacite of the 

2004–2005 Mount St. Helens Eruption: Constraints for 

Magma Origin and Dynamics

By Martin J. Streck1, Cindy A. Broderick1, Carl R. Thornber2, Michael A. Clynne3, and John S. Pallister2

Abstract

We investigated plagioclase phenocrysts in dacite of the 

2004–5 eruption of Mount St. Helens to gain insights into the 

magmatic processes of the current eruption, which is characterized by prolonged, nearly solid-state extrusion, low gas 

emission, and shallow seismicity. In addition, we investigated 

plagioclase of 1980–86 dacite.

Light and Nomarski microscopy were used to texturally 

characterize plagioclase crystals. Electron microprobe analyses measured their compositions. We systematically mapped 

and categorized all plagioclase phenocrysts in a preselected 

area according to the following criteria: (1) occurrence of 

zones of acicular orthopyroxene inclusions, (2) presence of 

dissolution surface(s), and (3) spatial association of 1 and 2. 

Phenocrysts fall into three main categories; one category contains four subcategories.

The range of anorthite (An) content in 2004–5 plagioclase is about An57–35 during the last 30–40 percent crystallization of plagioclase phenocrysts. Select microphenocrysts 

(10–50 μm) range from An30 to An42. Anorthite content is 

lowest near outermost rims of phenocrysts, but zonation patterns between interior and rim indicate variable trends that 

correlate with textural features. Crystals without dissolution 

surfaces (about 14 percent of total) show steadily decreasing An content outward to the crystal rim (outer ~80 μm). 

All other crystals are banded as a consequence of dissolution; dissolution surfaces are band boundaries. Such crystals 

display normal outward An zoning within a single band that, 

following dissolution, is then overgrown abruptly by high-An 

material of the next band. Swarms of acicular orthopyroxene 

inclusions in plagioclase are characteristic of 2004–5 dacite. 

They occur mostly inward of dissolution surfaces, where band 

composition reaches lowest An content. The relative proportions of the three crystal types are distinctly different between 

2004–5 dacite and 1980s dome dacite.

We propose that crystals with no dissolution surfaces are 

those that were supplied last to the shallow reservoir, whereas 

plagioclase with increasingly more complex zoning patterns 

(that is, the number of zoned bands bounded by dissolution 

surfaces) result from prolonged residency and evolution in 

the reservoir. We propose that banding and An zoning across 

multiple bands are primarily a response to thermally induced 

fluctuations in crystallinity of the magma in combination with 

recharge; a lesser role is ascribed to cycling crystals through 

pressure gradients. Crystals without dissolution surfaces, 

in contrast, could have grown only in response to steady(?) 

decompression. Some heating-cooling cycles probably 

postdate the final eruption in 1986. They resulted from small 

recharge events that supplied new crystals that then experienced resorption-growth cycles. We suggest that magmatic 

events shortly prior to the current eruption, recorded in the 

outermost zones of plagioclase phenocrysts, began with the 

incorporation of acicular orthopyroxene, followed by last 

resorption, and concluded with crystallization of euhedral 

rims. Finally, we propose that 2004–5 dacite is composed 

mostly of dacite magma that remained after 1986 and underwent subsequent magmatic evolution but, more importantly, 

contains a component of new dacite from deeper in the magmatic system, which may have triggered the new eruption.

Introduction

The 2004 eruption of Mount St. Helens is remarkable for 

several reasons. Nearly solid, gas-poor dacite lava has been 
792 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

extruded continuously for 28 months (at time of this writing, 

early 2007). Earthquakes are limited to the upper 3 km, with 

most located less than 1 km below the surface; no deeper seismicity has been observed (Moran and others, this volume, chap. 

2). Emissions of SO2, H2S, and CO2 are extremely low, indicating eruption of degassed magma (Gerlach and others, this volume, chap. 26). Low gas emission argues against the possibility 

of proximal mafic magma at depth, which might be called upon 

to have initiated the current eruption, and is consistent with an 

apparent lack of direct evidence for mingling with more mafic 

magma. This combination raises two important questions: (1) 

what is the driving force for the current eruption, and (2) is there 

petrologic evidence for magma recharge as an eruption trigger?

We studied plagioclase phenocrysts in 2004–5 dacite 

(fig. 1) and 1980–86 dacite dome rocks with the goal of using 

these data to investigate magma origin and reservoir dynamics, 

including evidence for recharge. We used polarized light and 

Nomarski microscopy in combination with detailed microprobe traverses to texturally and compositionally characterize 

single plagioclase crystals, focusing on areas near phenocryst 

rims. In addition, we systematically mapped and classified all 

plagioclase phenocrysts along thin-section traverses to evaluate variability in plagioclase crystal populations. On the basis 

of these datasets, we infer that the new dacite is composed 

of three components. The first is magma that remained in the 

reservoir after 1986 and whose plagioclase crystals typically 

underwent crystallization and resorption as crystals cycled 

through cooler and hotter (and possibly deeper) parts of the 

reservoir, respectively. Whether or not some of this magma was 

isolated and escaped any modification remains open to debate. 

The second component is magma that was probably supplied 

recently by recharge of the shallow Mount St. Helens magma 

chamber. This magma carried plagioclase phenocrysts that 

grew continuously and are compatible with crystallization controlled largely by decompression (Blundy and Cashman, 2001). 

In addition to these magmatic components, there is evidence 

that some plagioclase in erupting dacite magmas was derived 

by disintegration of wall-rock xenoliths constituting the third 

component. These include crystals with distinct sieve-textured 

cores probably derived from gabbroic source rocks (Heliker, 

1995) and older Mount St. Helens dacite containing plagioclase 

with unusually low An content and sporadically hosting quartz 

inclusions (M.J. Streck, unpub. data; Clynne and others, this 

volume, chap. 28). Our data are compatible with rechargedriven initiation of the current eruption. We suggest that new 

magma is dacite and is hotter than and has fewer crystals than 

the resident dacite with which it has blended.

Samples and Analytical Procedures

Samples for this study are splits of samples collected by 

the staff of the Cascades Volcano Observatory (CVO) as the 

eruption proceeded (table 1). We report collection dates for 

2004–5 samples and eruption dates for 1980–86 samples. The 

SH304-2B

SH315-4

1 mm

SH305-1

Figure 1. Thin sections of dacite from 2004–5 Mount 

St. Helens eruption showing crystal-rich nature and 

plagioclase occurrence mostly as single, equant 

phenocrysts. Note sporadic plagioclase with sieved 

interior and clear overgrowth.

full sample names are given in table 1 but are abbreviated. The 

reader is referred to Pallister and others (this volume, chap. 

30) for information about likely eruption dates for 2004–5 

samples. For most of our samples, the estimated eruption date 

precedes collection date by 1–3 weeks. Sample localities are 
2004–5 Dacite samples Normalized abundance, in %, with parenthetical 1σ deviation Dissol

Typ

(s

Sample No. Collection date EMP analyses Sheets per thin section Total crystals examined Crystals per sheet 1σ dev. Type 1 crystals Type 2 crystals Type 3 crystals Indeter- minate crystals A 

SH304-2B Nov. 4, 2004 16 17 1,939 114 17 66 (7) 24 (7) 10 (5) 12 70 

SH305-1 Jan. 3, 2005 6 16 1,118 70 8 40 (6) 29 (6) 31 (6) 20 75 SH308-3 Feb. 22, 2005 -- 18 1,653 92 12 52 (6) 28 (8) 20 (6) 11 70 SH311-1B Jan. 19, 2005 9 -- -- -- -- -- -- -- -- -- 

SH315-4 Apr. 19, 2005 8 13 1,354 104 21 57 (11) 34 (10) 9 (4) 8 69 

SH321 Aug. 19, 2005 -- 17 1,275 75 18 52 (5) 37 (7) 11 (5) 5 69 

SH323-2 Oct. 18, 2005 -- 16 1,213 76 15 65 (6) 27 (6) 8 (2) 7 82 SH324-3 Dec. 15, 2005 -- 16 1,290 81 6 66 (5) 28 (4) 6 (3) 5 77 

average 57 (10) 30 (4) 14 (9)

1980–86 Dacite samples Normalized abundance, in %, with parenthetical 1σ deviation Disso

Typ

(

Sample No. Eruption date EMP analyses Sheets per thin section Total crystals examined Crystals per sheet 1σ dev. Type 1 crystals Type 2 crystals Type 3 crystals Indeter- minate crystals A 

SH226 October 1986 -- 17 747 44 12 80 (10) 4 (3) 16 (9) 16 36 SH157 June 1984 -- 15 490 33 7 91 (4) 5 (3) 4 (3) 7 14 SH131 August 1982 -- 10 326 33 6 88 (7) 5 (5) 7 (4) 12 19 SC334-99B ~April 1981 9 14 835 60 8 89 (4) 1 (1) 10 (4) 16 50 SH52 December 1980 -- 10 181 18 6 84 (9) 9 (6) 7 (5) 11 8 

average 86 (6) 5 (3) 9 (6)

Table 1. Samples analyzed in this study and results of crystal mapping.[All plagioclase assigned to crystal type: Type 1, dissolution surfaces; Type 2, dissolution surfaces and acicular orthopyroxene (opx); Type 3, no dissolution surface and with or withcrystals further grouped by position of acicular opx relative to dissolution surface(s): A, opx abundant at outermost dissolution surface; B, opx at some surface inboard of the outermbut dissolution surface is near core of crystal; D, opx is near core but dissolution surface is near rim of crystal. EMP, electron microprobe analyses; dashes indicate no analysis. Colunumber of images from two transects across each thin section (see fig. 2). Column for indeterminate crystals shows proportion (percent) of unresolved assignments relative to total 
794 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

shown on photogeologic maps in Herriott and others (this 

volume, chap. 10).

We prepared polished thin sections that were etched in 

concentrated hydrofluoroboric acid for 25 s in preparation for 

Nomarski differential interference contrast (NDIC or Nomarski, for brevity) microscopy. This microscopy, in combination with transmitted-light microscopy, was employed to 

investigate textural aspects of plagioclase crystals. To associate textural characteristics with composition, we performed 

microprobe analyses along traverses in crystals representative 

of the range of textural features displayed. Point spacing along 

analytical traverses was 3–4 μm, concentrating on the outer 

~100 μm of the crystal. We analyzed 39 crystals from 2004–5 

samples and 9 crystals from 1981 dome lava (table 1). Analyses were done using the five-spectrometer CAMECA SX100 

electron microprobe housed at Oregon State University, which 

was mostly operated remotely from Portland State University. Some additional crystals were analyzed in single-point 

mode at the University of Stuttgart, Germany, to characterize 

very small plagioclase crystals. Natural mineral standards 

were used to calibrate the instruments. Analytical conditions 

included an accelerating voltage of 15 kV, a beam current 

of 15 nA, and a focused (~1 μm diameter) beam. Collection 

times on peak and background positions were 10 and 5 s for 

Na (counted first), Si, Al, and Ca; times were 30 and 15 s for 

K, Mg, and Fe.

Encouraged by initial efforts to categorize plagioclase 

phenocrysts into different textural types (see below), we 

designed a mapping procedure to determine proportions 

of plagioclase crystal types. For each thin section, we used 

transmitted-light images at magnification × 25 and mapped 

SH304-2B (Nov. 2004)

x 25 image (sheet)

Figure 2. Thin section showing 

mapped areas. Stripes comprise 

sequential image sheets, enlarged 

x 25, on which all plagioclase 

phenocrysts (long dimension >80 

μm) were classified.

two separate transects, comprising a total of 10–18 adjacent images, or sheets (table 1, fig. 2). In each × 25 sheet, 

we inspected every plagioclase phenocryst (≥80 μm) with 

transmitted and Nomarski microscopy and categorized crystals according to selected textural features. In this way, we 

obtained textural data on all plagioclase phenocrysts within an 

area covering roughly one-third of each thin section (table 1).

Plagioclase in 2004–2005 Mount St. 

Helens Dacite

The phenocryst assemblage of 2004–5 dacite is dominated by plagioclase with subordinate amounts of orthopyroxene, amphibole, and Fe–Ti oxides. Plagioclase consists 

dominantly (≥90 percent) of single, euhedral, clear, equant 

phenocrysts 80–800 μm across (fig. 1). Equant plagioclase 

also occurs as smaller crystals between 80 and 10 μm. We call 

these microphenocrysts even though they overlap in size with 

microlites in the surrounding interstitial groundmass. A few 

plagioclase crystals are texturally (and likely compositionally) 

distinct. Typically these are considerably larger (≥1 mm), subhedral to anhedral, and commonly display a sieved-textured 

interior. Plagioclase glomerocrysts are rare.

Textural Features

Individual plagioclase phenocrysts are characterized by 

growth features in the form of oscillatory zoning (parallel lines 

in Nomarski images) or by practically textureless 

parts (thus appearing as flat areas in Nomarski 

images) (fig. 3). Resorption textures are commonly found within crystals as dissolution surfaces 

that cut obliquely across oscillatory growth zones 

(figs. 3A–D) (note: we use the term “resorption” 

for the general process that creates a “dissolution 

surface”). The extent of resorption varies and most 

commonly is expressed as variable degrees of 

rounding of the corners of interior zones within a 

crystal. Commonly the phenocrysts have multiple 

dissolution surfaces; examples with four or more 

were observed. Sieved or pitted textures are rare 

in crystals of the dominant population, especially 

near their rims, but occur frequently in the interiors 

of the large crystals (>800 μm). Plagioclase phenocrysts may contain inclusions of orthopyroxene, 

Fe–Ti oxide, amphibole, apatite, and glass (melt). 

A characteristic feature of plagioclase of the 

2004–5 eruption is the occurrence and distribution 

of tiny orthopyroxene inclusions. They are acicular, with length-to-width ratios of ~10 or greater. 

Their length is variable but typically about 30 μm. 

Their maximum widths are on the order of 4–5 μm. 

Acicular orthopyroxene inclusions occur mostly 
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 795

in swarms and are concentrated in bands close to dissolution 

surfaces (figs. 3C, 3D), although they also occur independently of the surfaces (figs. 3E, 3F). Where associated with a 

dissolution surface, acicular orthopyroxene bands are, in most 

cases, inward of the dissolution surface, toward the center of 

the crystal (figs. 3, 4).

opx

opx

Dissolution surface

Dissolution surface

opx

opx

Dissolution surface

Dissolution surface

100 μm

100 μm

100 μm

E F

A B

C D

Compositional Features

The overall compositional range for the outer 80–100 

μm of plagioclase phenocrysts in the 2004–5 dacite ranges 

from An32–38 to An55–60 (figs. 4, 5, 6). At or near the rim (that 

is, within one traverse step of ~3 μm), compositions are 

Figure 3. Plagioclase types distinguished by combination of dissolution surfaces (arrows) and presence 

of acicular orthopyroxene inclusions (opx). Scale bars are 100 μm. Left column, transmitted light images; 

right column, corresponding Nomarski differential interference contrast images. A, B, Type 1 plagioclase, 

with two dissolution surfaces but without acicular orthopyroxene inclusions. C, D, Type 2 plagioclase, 

with acicular orthopyroxene inclusions inbound of dissolution surface. E, F, Type 3 plagioclase with no 

dissolution surface but containing, in this example, acicular orthopyroxene inclusions near rim.
796 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

An35–44, although An37–38 is most common. Analyzed microphenocrysts are 10 to 50 μm in size and range from ~An30 to 

~An42 (fig. 6). Compositions of microphenocrysts or at the 

rims of phenocrysts in 2004–5 samples have a wider range in 

An content than that reported for samples of 1980s dome dacite (Cashman, 1992), but one of our comparison samples of 

1980–86 dome dacites (SC-99-334b; table 1, fig. 6) suggests 

a similar range with minima of An38–40 and maxima of An60–65

and outermost rim compositions of An35–48. We exclude the 

large and sieve-textured plagioclase crystals from our analysis 

because they are thought to be derived from disaggregated 

plutonic or cumulate inclusions (Heliker, 1995). Contrasting 

with the above range in An content of phenocrysts are fairly 

sodic compositions of An24–32 that we found in the interior 

parts of some phenocrysts. Such phenocrysts have not been 

reported previously from Mount St. Helens dacite. They are 

compositionally distinct but texturally indistinguishable from 

typical phenocrysts unless they contain abundant mineral 

inclusions (some of which are acicular). We also exclude 

these because we believe that the unusually sodic crystal 

interiors represent recycled older crystals derived from moresilicic magma (see below).

Figure 4. Correlation of textures and composition. Solid lines on Nomarski images indicate 

location of analytical traverses, and short arrows point to acicular orthopyroxene inclusions. 

Arrows in graphs indicate location of dissolution (resorption) surfaces, and bars show 

location of acicular orthopyroxene inclusions.

30

0 10 20 30 40 50 60

0 10 20 30 40 50 60

70

0 10 20 30 40 50 60 70 80

35

40

45

50

55

60

30

35

40

45

50

55

60

30

35

40

45

50

55

DISTANCE FROM RIM, IN MICRONS

ANORTHITE CONTENT, IN MOLE PERCENT

Dissolution surface Dissolution surface

Dissolution surface

SH304-2B-F5

SH305-1-F7

SH304-2B-F8b

SH304-2B-F5

SH304-2B-F8b

SH304-2B-F5

SH305-1-F7

SH305-1-F7
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 797

Figure 5. An profiles of Type 2 crystals—those with dissolution surfaces and acicular 

orthopyroxene inclusions—except bottom row, which depicts two Type 3 crystals (no 

resorption). Lower right, one 1981 Type 1 crystal (dissolution surface but no inclusions). Arrows 

indicate location of dissolution surface and bars show location of acicular orthopyroxene 

inclusions. Resorption is typically followed by abrupt increase in An content. Sequential growth 

is rimward, toward left on each graph.

30

35

40

45

50

55

60

30

35

40

45

50

55

60

30

35

40

45

50

55

60

SH304-2B-bF7

SH305-1-AF1

SH304-2B-F8 SH305-1-AF3

SH304-2B-F11

SH304-2B-F9

SH304-2B-aF7

SH311-1B-F2 SH311-1B-F4

SH315-4-BF4

30

35

40

45

50

55

60

SH315-4-BF5 SH315-4-BF6

SH315-4-BF3

30

0 10 20 30 40 50 60 70 80

0 10 20 30 40 50 60 70 80

0 10 20 30 40 50 60 70

0 10 20 30 40 50 60 70

0 10 20 30 40 50 60 70

0 10 20 30 40 50 60 70

0 10 20 30 40 50 60 70

0 20 40 60 80 100 0 20 40 60 80 100

0 20 40 60 80 100

0 10 20 30 40 50 60

0 10 20 30 40 50 60

0 10 20 30 40 50 60 70 80

0 10 20 30 40 50 60 70 80

0 20 40 60 80

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60

30

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60

30

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50

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60

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60

30

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40

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60

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60

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55

60

SH315-4-BF7 1981: SC99-334b-F7

DISTANCE FROM RIM, IN MICRONS

ANORTHITE CONTENT, IN MOLE PERCENT

Type 3 xls Type 3 xls Type 1 xls

Near their rims, all plagioclase phenocrysts are compositionally normally zoned within bands that show oscillatory zoning or flat textures. We distinguish compositional from textural 

zoning because texturally flat areas are frequently compositionally zoned (for example, fig. 4, sample SH304-2B-F5; table 2). 

Band boundaries are dissolution surfaces. Band widths are variable and can be as narrow as ~10–15 μm. Some smaller crystals 

do not show any dissolution surface and therefore can be 

considered to consist of a single, broader band (about 200–300 

μm). Compositional changes within a single band, regardless 

of width, in most cases display An decreases from 7 to 20 mol 

percent and thus may encompass the entire range of An variability of plagioclase in 2004–5 dacite (table 2). Compositionally 

uniform bands and bands with slight reverse zoning (∆ ~An5) 

are rare. The initial overgrowth on a dissolution surface always 

has the highest An content; last growth is the lowest An content 
798 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

within a band—a feature also observed in May 18, 1980, dacite 

(Pearce and others, 1987). In fact, it is an often-observed feature in volcanic plagioclase with dissolution surfaces (Pearce, 

1994). Along a single analytical traverse, abrupt An increases 

of 5 to 25 mol percent were observed where crossing from the 

lowest An content of an inner band to the highest An content 

of the next outer band (figs. 4, 5). Thus the profiles undulate, 

depending on whether one or multiple dissolution surfaces exist. 

Rarely, overgrowth on a more interior dissolution surface can 

be of distinctly lower An content than inward of the dissolution 

surface.

Acicular orthopyroxene inclusions are consistently associated with the lower-An parts of bands. They occur where the 

plagioclase composition reaches An33–40. The composition of 

acicular orthopyroxene inclusions is similar to the composition of orthopyroxene phenocrysts (table 3), although Al and 

perhaps Ca are slightly higher. It is uncertain whether the Ca 

and Al variation of acicular inclusions compared to orthopyroxene phenocrysts is significant given the spatial difficulties 

in analyzing small crystals embedded in plagioclase.

Crystal Mapping

The mapped crystal populations are distinguished by 

presence or absence of near-rim features produced by alternating growth and resorption events and by the presence or 

absence of acicular orthopyroxene inclusions. On the basis 

of these features, we can distinguish three crystal types and 

determine their proportions (table 1):

•	 Most abundant are plagioclase crystals containing one 

or several dissolution surfaces but lacking acicular 

orthopyroxene; they are designated Type 1 crystals.

•	 Second in abundance are plagioclase crystals that show 

dissolution surfaces and contain acicular orthopyroxene inclusions; they are designated Type 2 crystals. We 

further distinguished subcategories of Type 2 crystals 

according to location of the orthopyroxene inclusions 

with respect to dissolution surfaces: (A) inboard of last 

dissolution surface, (B) only associated with an older 

dissolution surface, (C) near the rim and not associated 

with a dissolution surface (that is, the dissolution surface is near the core of the crystals and orthopyroxene 

inclusions occur near the rim), and (D) near the crystal 

core and the dissolution surface is near rim.

•	 Least abundant are crystals lacking clearly discernible dissolution surfaces; they are designated Type 

3 crystals. For the Type 3 category, we established 

subcategories based on whether acicular orthopyroxene 

inclusions are present or absent.

In some cases, categorization was difficult owing to poor 

crystal polish, fracturing, disadvantageous cutting of the crys30

0 20 40 60 80 100 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7

0 20 40 60 80 100 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7

40

50

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70

30

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50

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70

30

40

50

60

70

30

40

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DISTANCE FROM RIM, IN MICRONS

ANORTHITE, IN MOLE PERCENT

NORMALIZED DISTANCE FROM RIM

2004–5 

1981

<30

30-3333-3636-3939-42

>42

2

0

4

6

Microphenocrysts (≤50 μm)

NUMBER

An Figure 6. Anorthite content in 

plagioclase, showing all data of 

this study. Plots in left column 

show An content versus distance 

to rim; in right column, An content 

versus normalized distance to rim. 

To normalize distance to rim, we 

connected traverse end point at 

rim with center of crystals and 

took fraction that was covered by 

traverse. We projected beginning 

of traverse into “rim – center” line 

along growth zones in case where 

analysis traverse was not parallel 

to “rim – center” line. Histogram 

shows plagioclase composition of 

22 microphenocrysts (~<50 μm) in 

SH321 (table 1), similar to the two 

observed in center and along left 

side of groundmass picture of sample 

SC99-334A shown in figure 8A. All 

microphenocrysts were analyzed in 

center and rim.
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 799

SH304-2B-F5 (Type 2 crystal)

Distance to rim (μm)------- 31.5 28.0 24.5 21.0 17.5 14.0 10.5 7.0 3.5 0

SiO2 58.3 58.7 59.1 60.1 54.9 55.7 57.1 57.5 58.4 59.8

Al2

O3 26.7 26.3 26.2 25.4 28.6 27.6 27.0 26.8 26.4 25.4

FeO* 0.29 0.28 0.34 0.30 0.50 0.48 0.47 0.48 0.45 0.49

MgO 0.01 0.01 0.02 0.02 0.04 0.04 0.04 0.03 0.03 0.02

CaO 8.7 8.5 8.2 7.5 11.6 10.6 9.7 9.4 8.8 7.8

Na2

O 6.2 6.4 6.6 6.9 4.8 5.2 5.7 5.9 6.2 6.7

K2

O 0.18 0.20 0.21 0.23 0.12 0.13 0.15 0.17 0.20 0.29

Total 100.3 100.3 100.6 100.4 100.5 99.9 100.2 100.2 100.5 100.5

An 43.2 42.1 40.4 37.2 57.2 52.6 47.9 46.6 43.7 38.8

Ab 55.7 56.8 58.3 61.4 42.1 46.7 51.3 52.4 55.1 59.5

Or 1.1 1.2 1.2 1.4 0.7 0.8 0.9 1.0 1.2 1.7

SH305-1-F7 (Type 3 crystal)

Distance to rim (μm)------ 30.4 27.0 23.7 20.3 16.9 13.5 10.1 6.8 3.4 0

SiO2 55.5 55.6 55.8 56.5 56.8 57.5 58.0 58.1 58.7 58.9

Al2

O3 27.8 28.1 27.8 27.2 26.9 26.7 26.2 25.5 25.7 25.8

FeO* 0.27 0.34 0.33 0.22 0.35 0.30 0.28 0.79 0.27 0.34

MgO 0.01 0.01 0.01 0.01 0.01 0.02 0.02 0.51 0.04 0.01

CaO 9.9 9.8 9.7 9.1 8.4 8.5 7.7 7.5 7.6 7.5

Na2

O 5.5 5.5 5.7 5.8 6.1 6.4 6.7 6.5 6.7 6.9

K2

O 0.14 0.14 0.16 0.19 0.18 0.19 0.23 0.22 0.23 0.25

Total 99.2 99.5 99.5 99.0 98.6 99.7 99.2 99.2 99.2 99.7

An 49.6 49.0 48.0 45.9 43.0 42.0 38.4 38.4 38.3 37.0

Ab 49.5 50.1 51.1 53.0 55.9 56.8 60.2 60.2 60.4 61.5

Or 0.8 0.8 0.9 1.1 1.1 1.1 1.4 1.3 1.4 1.5

Table 2. Representative plagioclase compositions along profiles of two crystals shown in figure 4, starting at 

~30 μm from the rim.

[FeO*, total iron as Fe2+. Molecular components for plagioclase shown calculated: An, anorthite; Ab, albite; Or, orthoclase.]

tal leading to obliteration near the rim, and other ambiguities. 

Crystals that could not be categorized clearly are considered 

indeterminate crystals; their proportions ranged from 5 to 20 

percent (table 1). In subsequent analysis, we assumed that the 

indeterminate crystals are proportionally distributed among the 

other populations and therefore normalized the other three categories to their sum to investigate changes in their proportions.

Comparing mapping results among sheets (images at 

× 25 magnification) that make up surveyed transects yielded 

standard deviations of mapped crystal types for a single thin 

section (table 1); 1σ standard deviations are typically on the 

order of 5–10 percent of the normalized abundance, and this 

appears to be the natural variation on the level of a single thin 

section, including the error associated with recognition.

An exception to Type 1 and Type 2 crystals being first 

and second in abundance is sample SH305-2 (lava collected 

Jan. 3, 2005), in which Type 3 crystals are slightly more abundant or subequal to Type 2 crystals (fig. 7).

Variations among crystal types in samples of the 2004–5 

dacite indicate some trends after apparent natural variation 

within individual samples (based on 1σ errors) is taken into 

consideration. Notable are (1) an increase in the abundance 

of Type 3 crystals in SH305, which taper back to the level 

observed at the onset of the eruption at the expense of Type 1 

crystals, and (2) a possible slight increase in the abundance of 

Type 2 crystals from SH304 through SH321. Interestingly, the 

increase of Type 3 crystals from SH304 (November 2004) to 

SH305 (January 3, 2005) correlates with an increase in magma 
800 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Phenocryst

Acicular

Inclusion Inclusion

Core Rim Point 1 Point 2 Point 1

SiO2 54.31 53.5 53.31 54.15 53.6

TiO2 0.21 0.14 0.28 0.21 0.1

Al2

O3 0.41 0.57 2.14 1.2 1.25

FeO* 21.38 23.74 20.02 20.66 23.17

MnO 0.66 0.79 0.47 0.44 0.76

MgO 23.72 21.68 23.45 23.57 22.08

CaO 0.86 0.73 1.59 1.24 0.73

En 65.3 61.0 65.5 65.4 62.0

Fs 33.0 37.5 31.4 32.2 36.5

Wo 1.7 1.5 3.2 2.5 1.5

Mg # 66.4 61.9 67.6 67.0 62.9

Table 3. Compositions of representative orthopyroxene 

phenocryst and acicular orthopyroxene inclusions in plagioclase 

phenocrysts.

[Features described as acicular inclusions are typical; those described 

simply as inclusions are also acicular but slightly larger. FeO*, total iron as 

Fe2+. Molecular components for pyroxene shown calculated: En, enstatite; 

Fs, ferrosilite; Wo, wollastonite. Magnesium number, Mg# = (MgO×100)/

(MgO+FeO*).] ABUNDANCE OF PLAGIOCLASE CRYSTAL TYPES, IN PERCENT

1980 (SH52)

1981 (SC-334B)

1982 (SH131)1984 (SH157)1986 (SH226)

11/04 (SH304)

1/05 (SH305)2/05 (SH308)4/05 (SH315)8/05 (SH321)

10/05 (SH323)12/05 (SH324)

0

10

20

30

40

50

60

70

80

90

100

Type 1

Type 2

Type 3

1980s 2004–5

Figure 7. Proportions of different plagioclase types from 

crystal mapping. Bars showing 1σ error are based on 

variations measured among mapped sheets (25x images).

temperature as obtained from two-oxide geothermometry, 

whereas the tapering off to the initial abundance correlates with 

a drop in magma temperature in the spring of 2005 (Pallister 

and others, this volume, chap. 30). In other words, the sample 

in which crystals devoid of dissolution surfaces are most 

abundant is also the sample from which the highest temperature 

values were obtained. We will revisit this relation below.

Comparison of crystal populations of 2004–5 dacites with 

samples of 1980s dome dacite reveals a striking difference, 

even if all 2004–5 samples are averaged (fig. 7, table 1). The 

Type 2 crystals are sparse in 1980s dacite, which is dominated by Type 1 crystals. Type 3 crystals display comparable 

abundances except a temporary increase in samples collected 

in January and February 2005. In essence, the main difference 

between 2004–5 and 1980 dacites is an absence of orthopyroxene inclusions in 1980s lava.

Discussion

Acicular Orthopyroxene Inclusions as Probable 

Markers of Cooling Events

The presence of acicular orthopyroxene inclusions in plagioclase phenocrysts of the 2004–5 dacite is a striking feature 

(figs. 3, 4, 8). They are also present in plagioclase of 1980s 

dome dacite, although much rarer than in 2004–5 samples (table 

1), and have since been observed by the first author elsewhere in 

andesite from Mount Hood, Oregon, and Volcán Arenal, Costa 

Rica. Thus, acicular orthopyroxene inclusions in plagioclase are 

not unique to this eruption, but their abundance and systematics of occurrence are presently unrivaled. Texturally identical 

orthopyroxene crystals occur in interstitial glass or groundmass 

in dacite of the current eruption, as well as in 1980s dome dacite 

(fig. 8). In particular, samples composed of interstitial glass 

laced by acicular orthopyroxene microlites and of scattered plagioclase microphenocrysts (<70 μm, either dimension) such as 

sample SH304-2G of the current eruption (Pallister and others, 

this volume, chap. 30) and SC99-334A (fig. 8), a 1981 sample, 

are important because they demonstrate that conditions exist 

in which interstitial melts of Mount St. Helens dacite primarily 

nucleate orthopyroxene that grew rapidly (as suggested by being 

acicular; for example, see Lofgren, 1980).

The supersaturation required to nucleate and rapidly grow 

orthopyroxene may be caused by a drop in temperature, by 

decompression associated with degassing, or by a combination of both. Samples with a more crystalline groundmass (for 

example, SH305-1, fig. 8) contain larger acicular orthopyroxene, which indicate that growth of acicular orthopyroxene continued as groundmass crystallinity increased. Plagioclase did 

not undergo a concurrent nucleation and rapid-growth event; 

instead of acicular crystals, equant plagioclase crystals continued to grow. Plagioclase growth in interstitial melt is strongly 

governed by the rate of decompression and devolatilization 

as magma ascends to the surface (Hammer and Rutherford, 
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 801

SC99-334A (1981), groundmass

SH305-1 (Jan. 3, 2005), groundmass Opx inclusion zone in SH315-4-F7 (Apr. 19, 2005)

Opx inclusion zones in SH321 (Aug. 19, 2005)

A B

C D

30

35

40

45

50

55

60

AN

0

0.2

0.4

0.6

0.8

1.0 FeO*

FeO*, IN WEIGHT PERCENT

AN, IN MOLE PERCENT

DISTANCE FROM RIM, IN MICRONS

100 20 30 40 50 60 70 80

Figure 8. Occurrence of acicular orthopyroxene crystals in groundmass (A, C) and as inclusions in 

plagioclase (B, D). All pictures in transmitted light. A and C at x 1000 magnification, B at x 500 magnification, 

and D at x 200 magnification. Inset in D shows analysis traverse along solid line.

2002). In general, rapid nucleation rate and acicular growth of 

plagioclase are induced by faster decompression and devolatilization rates. The occurrence of acicular groundmass orthopyroxene in the absence of acicular plagioclase in SH304-2G and 

SC99-334A (both representing magma seemingly saturated in 

orthopyroxene and plagioclase) suggests that temperature may 

have played an equal or more important role than decompression and devolatilization to induce supersaturation in orthopyroxene but not in plagioclase. Only slow and steady or small 

decompression steps are compatible with equant shapes of plagioclase microphenocrysts (Hammer and Rutherford, 2002).

An alternative interpretation, that a boundary layer 

enriched in orthopyroxene components rejected by growing 

plagioclase and subsequent local saturation and crystallization, 

is unlikely. A few orthopyroxene crystals may be generated 

this way but not the swarms of inclusions observed (fig. 8). 

Constant or decreasing Fe content in profiles approaching 

bands of orthopyroxene inclusions (fig. 8D) suggests no major 

enrichment of total Fe in the boundary layer.

In conclusion, we believe that the presence of acicular 

orthopyroxene in some plagioclase phenocryst bands reflects 

an abundance of acicular orthopyroxene microlites in the 

interstitial melt at the time of embedding. Therefore, acicular 

orthopyroxene-rich zones in plagioclase record rapid orthopyroxene-nucleation events that were probably induced in large 

part by rapid cooling.

Production of Near-Rim Plagioclase—an 

Evaluation of Processes and Natural 

Constraints

Cooling, decompression, compositional and volatile 

fluxes, and kinetic effects all may control compositional and 

textural features of plagioclase—although we consider kinetic 

effects (for example, see Pearce, 1994) at best a secondary 

cause for the abrupt increase of An content by more than 2–4 

mol percent (for example, see Ginibre and others, 2002). It is 
802 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

typically a certain combination of features that narrows down 

which parameter exerted greater control.

Below we make the case that the characteristic features 

of Types 1 and 2 crystals—multiple, normally zoned bands 

separated by dissolution surfaces within the outer ~80 μm 

of crystals, combined with marked declines in An content 

within variably thick bands (as thin as 10–20 μm)—are most 

consistent with fluctuations in crystallinity as primarily 

induced by temperature, with other parameters contributing 

in secondary roles. On the other hand, normal zoning (that is, 

progressive outward decrease of An content toward rim) over 

wider (>~50 μm) distances without dissolution surfaces—as 

observed in Type 3 crystals—is less constrained and may 

record times when other parameters such as pressure outweigh temperature changes.

If the magmatic system were closed, then higher and 

lower An content would imply lower or higher crystallinity, 

respectively, at the time of crystallization. To correlate the 

degree of crystallinity with a particular An content in plagioclase, we calculated melt composition at a variety of crystallinities and varied the partition coefficient, Kd (Kd = Ca/Naplag/

Ca/Namelt) (Sisson and Grove, 1993), to calculate equilibrium 

compositions of plagioclase (fig. 9). For our model calculations, we used an average bulk composition from initial anlyses of November 4, 2004, dacite (SH304) as starting composition (final analyses in Thornber and others, 2008b). A fixed 

mineral assemblage was removed from the bulk composition 

to calculate interstitial melt composition. The assemblage 

consisted of 6 percent amphibole, 9 percent orthopyroxene, 18 

percent An53 plagioclase, 65 percent An42 plagioclase, and 2 

percent titanomagnetite and is based on mineral modes established for 2004–5 dacite (Pallister and others, this volume, 

chap. 30). Although the extraction oversimplifies any natural 

process, certain important features are illustrated. If SH304 

ever existed as pure liquid, it would initially crystallize An53

(at Kd=2) to An75 (at Kd=5) under lower to higher water-saturated conditions, respectively (Sisson and Grove, 1993) (fig. 

9). We estimate that our investigation applies to approximately 

the last 30–40 percent of plagioclase phenocryst crystallization (synonymous with whole-rock crystallinities above ~30 

percent), as suggested by lengths of microprobe traverses combined with range of phenocryst sizes studied. Thus, the highest 

commonly observed An content of 55 to 60 mol percent was 

achieved when crystallinity of SH304 dacite was already at 

or near 30 percent. This crystallinity requires a minimum Kd 

of 3, or higher at higher crystallinities, in order to generate 

An55–60 at water-saturated conditions (fig. 9), an interpretation 

supported by water-saturated experimental results obtained 

on dacite SH305, in which conditions of 870oC and 200 MPa 

produced An50 plagioclase at a crystallinity of ~30 percent 

(Rutherford and Devine, this volume, chap. 31). On the other 

hand, an An content of 37 mol percent requires crystallinity 

near 50 percent, in keeping with the observation that 2004–5 

dacite consists of ~45 percent groundmass (interstitial glass 

and/or groundmass crystals), and An37 is the typical rim and 

microphenocryst composition. As a consequence, crystals that 

have equally low or even lower An content further inward than 

at the rim (figs. 4, 5, 6) mark earlier times when crystal growth 

occurred at crystallinities between 50 and 60 percent.

A decrease of An content from 55 to 37 mol percent, as 

is observed commonly within single compositionally zoned 

bands, requires the following conditions or some combination thereof: (1) ~25 percent crystallization at constant Kd or 

(2) decreasing Kd at constant crystallinity (for example, at 40 

percent crystallinity, Kd would need to decrease from ~4.5 to 

2). A decrease in Kd at water-saturated conditions and a given 

melt composition is synonymous with a decrease in water concentration and thus decrease in Ptotal (Sisson and Grove, 1993). 

To explain crystals with several dissolution surfaces by changes 

in Kd would be problematic, as changes in Kd would need to 

be dramatic and cyclic. If Kd exerted a dominant control on 

observed An content, then a return to An50–55 after resorption 

would require a substantially higher Kd and would require the 

crystal to return to a greater depth (Hattori and Sato, 1996). 

Water-saturated experiments indicate that an increase in pressure of 60 MPa would increase An content by 5 mol percent in 

Mount St. Helens dacite magmas (Rutherford and Devine, this 

volume, chap. 31). Thus crystals with bands with An contents 

≥5 mol percent higher after resorption would require a depth 

increase of ~2 km or more, and crystals that display several 

such bands would have been cycled numerous times. In fact, to 

10

0

0 10 20 30 40 50 60

20

30

40

50

60

70

80

Kd = 2

Kd = 3

Kd = 4

Kd = 5

PLAGIOCLASE ANORTHITE

CONTENT, IN PERCENT

CRYSTALLINITY, IN PERCENT

Figure 9. Variation of An content in plagioclase versus 

degree of crystallinity based on model calculations using 

variable Kd (=(Ca/Na plag)/(Ca/Na melt)), bulk composition of 

SH304 as starting composition (table 1), and removal of a fixed 

mineral assemblage (see text for details). Model simulates 

correlation of degree of crystallinity with plagioclase An 

composition in equilibrium with corresponding interstitial 

melt. Shaded zone marks conditions covered by this study, as 

suggested by lengths of microprobe traverses combined with 

range of phenocryst sizes investigated.
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 803

explain the entire An range of An55 to An37 would require pressure to fluctuate by 150 MPa, equivalent to a depth change of 

~5.5 km (Rutherford and Devine, this volume, chap. 31). 

An additional complicating factor is that once magma 

has reached water saturation and begins to degas at shallow 

depth, it will become undersaturated as it reaches greater 

depths unless volatiles are added to keep it saturated. If it did 

not gain volatiles, descent towards higher degree of undersaturation would promote crystallization as the liquidus is raised 

and likely would crystallize plagioclase of lower An content 

(for example, Blundy and Cashman, 2001). To return multiple 

times from An37 to An55 at a constant Kd poses comparably 

little difficulty. The only requirement is that crystallinity 

needs to fluctuate by ~25 percent or less to return to an An 

content of An55; this could be achieved by local cooling and 

subsequent entrainment into more interior and hotter parts of 

the magmatic reservoir (Singer and others, 1995; Couch and 

others, 2001). In fact, Rb concentrations in melt inclusions 

of 1980s dacite suggest crystallinity variation by 20 percent 

at a constant pressure over the pressure interval from 200 to 

100 MPa (Blundy and others, 2006)—the inferred pressure 

of the upper part of the 2004–5 magmatic reservoir (Pallister 

and others, this volume, chap. 30). A temperature change of 

about 45oC would be needed to achieve this, consistent with 

experimental results showing that An55 plagioclase crystallizes at 890oC and An37 crystallizes at 845oC (Rutherford and 

Devine, this volume, chap. 31). Such temperatures are also 

within the range deduced from oxide geothermometry of 

2004–5 dacite (Pallister and others, this volume, chap. 30) and 

are similar to isobaric variations deduced from plagioclaseliquid geothermometry on 1980s dacite (Blundy and others, 

2006). Latent heat of crystallization (Blundy and others, 2006) 

may have contributed to the increase in temperature. On the 

other hand, magma heating by decompression-driven crystallization (Blundy and others, 2006) appears minimal above 50 

MPa and is most significant from about 40 MPa to the surface. 

Therefore the greatest impact of latent heat is taking place at 

pressure significantly shallower than conditions of the magma 

reservoir. Consequently, variation of magma temperatures 

observed at pressures above 50 MPa require additional explanation (see section “New Versus 1980s Residual Magma and 

Evidence for Magmatic Recharge,” below).

Returning to the discussion of our compositional data, 

alternative possibilities under conditions of an open system 

(that is, involving magma and (or) volatile flux into or out of 

the system) include: (1) an influx of higher Ca/Na magma 

and (2) switching between water-undersaturated and saturated conditions as the system ranges between lithostatic and 

hydrostatic conditions. The latter is the only instance where a 

decrease in Ptotal is not synonymous with decreasing depth as 

the system degasses under hydrostatic conditions and then is 

repressurized as connection to the surface ceases and the system returns to lithostatic conditions. Periodic influx (recharge) 

of more mafic (higher Ca/Na) magma may help to reset 

conditions for crystallization of a more An-rich plagioclase 

onto dissolution surfaces, and there is circumstantial evidence 

that this effect is occurring (see section “New Versus 1980s 

Residual Magma and Evidence for Magmatic Recharge”). 

Furthermore, this mechanism works in concert with the notion 

that abrupt near-rim decrease in An content within single 

bands is driven mostly by increased crystallinity owing to 

cooling. The other possibility—in essence reflecting shallow 

and periodic degassing events—has been advocated to occur at 

El Chichón volcano during eruption and repose cycles (Tepley 

and others, 2000). Degassing would promote crystallization, 

and resorption would then need to occur during repressurization of the system. This process may have contributed to 

features observed in plagioclase of the 2004–5 dacite at Mount 

St. Helens (see below) but cannot serve as sole explanation as 

clearly evidenced by juxtaposed plagioclase phenocrysts with 

no, single, or multiple dissolution surfaces, therefore requiring 

a range in the number of degassing events.

Crystal Histories and Populations

On the basis of our textural and compositional analysis, 

it appears that all plagioclase phenocrysts share, at best, only 

the last An decrease over the last 5–20 μm. Thus, it seems 

impossible to generate the diverse features in plagioclase 

described above during a single ascent in which essentially 

all phenocrysts grow—as is envisioned for plagioclase of the 

dome dacite of the 1980s (Blundy and Cashman, 2001). It 

appears that the only plagioclase crystals for which a singleascent history is conceivable are crystals that display monotonically decreasing An content (our Type 3) and probably 

crystals that show mostly continuously decreasing An content 

but with a dissolution surface deeper in the core of the crystal. 

These could be a product of continuous crystallization driven 

by decompression (Blundy and Cashman, 2001). In our view, 

plagioclase crystals that are poorly explained by a continuous 

ascent-driven crystallization are those with multiple, normally zoned bands separated by dissolution surfaces. Nearly 

isobaric temperature fluctuations (corresponding to degree of 

crystallinity) could account for multiple dissolution surfaces 

as discussed above. On the other hand, such zonation patterns 

(also known as sawtooth patterns) have recently been attributed to processes of magma heating by decompression-driven 

crystallization (Blundy and others, 2006). As argued above, 

such processes are strongest at depths of less than 1 km and 

therefore would occur in the conduit and would impact all 

plagioclase during their final ascent.

In the discussion of crystal populations, two important 

questions concern the significance of any recognized crystal 

population and the significance of variations from it. The 

answers to both questions depend on the processes controlling the features that define crystal types, the minimal changes 

needed to induce these features, and the minimal distances 

needed to see a response to differing environmental conditions. 

As presented above, we based our crystal types on existence of 

dissolution surfaces separating growth bands and the presence 

of acicular orthopyroxene. This is analogous to plagioclase 
804 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

crystal populations established by Pearce and others (1987) for 

dacite of May 18, 1980. The abundance of dissolution surfaces 

in plagioclase from Mount St. Helens dacite and from dacite 

elsewhere (for example, see Pearce, 1994) suggests that these 

features do not require major changes in magmatic conditions 

but are produced rather easily. Therefore, neighboring crystals 

with a variable number of dissolution surfaces likely experienced different growth histories and were subsequently juxtaposed by mixing. We propose that crystals with no dissolution 

surface are those that were supplied to the erupting dacite 

magma last, whereas increasingly more complex textures, 

especially several zoned bands bound by dissolution surfaces, 

are evidence of a prolonged residency and evolution in a shallow reservoir (for example, Tepley and others, 2000). Crystals 

with multiple bands can be explained solely by temperature 

gradients and, thus, by crystallinity gradients across the reservoir, in combination with recharge by higher Ca/Na melt (fig. 

10). Cycling through polybaric conditions, seemingly required 

to explain amphibole compositions in 2004–5 dacite (Rutherford and Devine, this volume, chap. 31), is permissible as an 

influence on the development of bands and zonation. However, 

narrow (~10–15 μm) bands with strong zonation (for example, 

∆An ≥10 mol percent), which suggest rapid, back-and-forth 

changes in growth conditions, argue against pressure as a principal control because the required cycling through a vertical 

distance of several kilometers (~4 km) is unrealistic.

New Versus 1980s Residual Magma and 

Evidence for Magmatic Recharge

One of the main questions asked about the renewed 

activity at Mount St. Helens is whether dacite lava that has 

extruded since October 2004 is magma that was stored in the 

subvolcanic reservoir since Mount St. Helens erupted in 1986 

or is freshly supplied magma. Bulk chemical analyses demonstrate that there are chemical differences between 1980s and 

2004–5 lava; for example, SiO2 is about 1.5–2.5 wt. percent 

higher in 2004–5 dacite than in most 1980s dome dacite 

(Pallister and others, this volume, chap. 30). Results from our 

crystal mapping hold clues to answering the question of new 

versus residual magma. The high ratio of Type 1 to Type 2 

crystals in 1986 versus 2004–5 lava in general and the higher 

proportion of Type 3 crystals in January and February 2005 

samples (table 1, fig. 7) suggest that lava extruded during the 

current activity is not from the same magma that was extruded 

during the final year of dome growth in 1986. Closed-system 

evolution from 1986 to today is one conceivable possibility 

to explain the observed differences in plagioclase phenocryst 

populations. Yet, any increases in Type 3 crystals in 2004–5 

lava cannot be due to modification of crystals with dissolutions surfaces, which makes the higher proportion of Type 3 

crystals early in the 2004–5 eruption in comparison to 1980s 

sample indicative of freshly supplied crystals. Furthermore, to 

explain increases in the ratio of Type 2 to Type 1 crystals by 

growth evolution since 1986, the following would be required: 

0

DISTANCE FROM RIM, IN MICRONS

180160140120100806040200

180160140120100806040200

0.2

0.4

0.6

0.2

0.3

0.4

FeO*, IN

WEIGHT PERCENT

30

35

40

45

50

55

60

~50% xls?

>60% xls?

~35% xls?

SH311-1B-F9

x-polars

K2O, IN 

WEIGHT PERCENT ANORTHITE, IN MOLE PERCENT

Figure 10. Analytical traverse of plagioclase SH311-1B-F9 

showing multiple normally zoned bands bounded by mild 

dissolution (resorption) surfaces. An peaks correlate inversely 

with K but not with Fe. We propose that highest An contents 

record magmatic conditions of low crystallinities (~35 percent 

crystals) whereas low An contents record high crystallinities 

(50–60 percent crystals). FeO*, total iron as Fe2+. 

crystallization to trap orthopyroxene followed by some resorption, in turn followed by crystallization of more plagioclase. 

From the above, it seems likely that 2004–5 dacite is composed mostly of variable proportions of magmatic components 

that have undergone shallow evolution (that is, modified, 

residual magma) and components that have been added from 

deeper parts (new magma). In other words, various magmatic 

sources provided plagioclase phenocrysts to explain multiple 

crystal histories, and phenocrysts were juxtaposed by mixing events (additionally, crustal sources supplied xenocrystic 

plagioclase, but we neglect these in the discussion here).
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 805

There is additional circumstantial evidence that the current eruption has experienced magmatic recharge by hotter 

magmas with more calcic melts (but not necessarily more calcic bulk composition). Assume that a magma batch undergoes 

closed-system crystallization to produce the same growth features in equally sized crystals. If growth were continuous, we 

would expect decreasing An content in the plagioclase crystals 

as fractional crystallization progresses and the melt fraction 

decreases (if crystal-melt equilibrium is not maintained). If 

the same system were to oscillate between higher and lower 

crystallinity during its course of crystallization, owing to 

temperature fluctuations, we would expect profiles of overall 

An decrease toward crystal rims but modulated by resorption 

and precipitation of higher-An bands (Pearce, 1994; Johannes 

and others, 1994). Partial plagioclase dissolution would raise 

the Ca/Na in the melt, and therefore initial overgrowth would 

be higher in An content than that before resorption. The An 

content of the initial overgrowth will depend on how much 

dissolution occurred and may be also influenced by nonequilibrium crystallization effects related to the kinetics of the 

system (Stewart and Fowler, 2001), but it could never return 

to the most calcic composition unless all plagioclase dissolved (neglecting the effect of volumetrically less significant 

calcic phases, for example hornblende or augite, on Ca/Na 

ratios). Subsequent growth would rapidly decrease An content 

to lower values as the system increased in crystallinity. This 

plagioclase response is illustrated schematically in figure 11 

as profile A, which is rarely seen in plagioclase of the current 

RIM CORE

Profile A—Rare

Profile B—Common

Profile C—Sporadic

Crystallinity 10% 8% 25% 21% 40% GROWTH RESORPTION GROWTH RESORPTION FINAL GROWTH

Anorthite,

 mole percent

EXPLANATION

50-58

42-50

35-42

Crystallinity 

Crystallinity sequence not shown 

40% 25% 40% 25% 40%

>58

<35

Size of box = relative reservoir mass

RELATIVE DISTANCE

TO RIM

RELATIVE

ANORTHITE CONTENT

RELATIVE

ANORTHITE CONTENT

RELATIVE

ANORTHITE CONTENT

RIM CORE

RIM CORE

Dissolution

surface

Dissolution

surface

Dissolution

surface

Figure 11. Schematic scenarios to produce common An profiles observed in 2004–5 plagioclase. Given An ranges 

shown in legend and crystallinity proportions are approximate values. To generate profile B, often observed in 

plagioclase of 2004–5 dacites, probably requires influx of melt (that is, an increase in reservoir mass) to buffer Ca/

Na. If a closed system were to fluctuate in its degree of crystallinity due to temperature variations, profile A would be 

expected to result, which is only rarely observed. We therefore propose that profile B and, even more so, profile C, 

are likely the signs of a recharged system rather than one that is closed for magma input.
806 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

eruption but is common elsewhere (for example, Tatara dacite, 

see Singer and others, 1995). A common An profile among 

2004–5 phenocrysts, profile B (fig. 11), shows An oscillations 

with minima and maxima immediately inward and outward 

of the dissolution surface, respectively, and the range of An 

content is similar over the course of crystallization (fig. 10). 

This constant An range is important because, as plagioclase on 

the whole crystallizes, the system can only produce similar An 

maxima if there is an influx of melt to buffer melt Ca/Na. The 

influx needed depends on the Ca/Na ratio in the melt that is 

recharged and other phases that may be crystallizing. In general, this requires an open system in which the overall crystal 

mass increases but where melts are buffered to produce the 

same An range during crystallization of multiple bands (fig. 

10). Plagioclase simply convecting through a thermal regime 

(Singer and others, 1995) would not work because, if balanced, the amount of crystallization on a cooling path and the 

amount of resorption during the heating stage would not yield 

an overall growing crystal with return to similar An minima 

and maxima unless some phenocrysts grow while others are 

increasingly dissolved, for which evidence is lacking. Shifting 

crystallization to progressively lower pressures over the course 

of growth of multiple bands would make it increasingly more 

difficult to return to comparable An maxima. Plagioclase crystals with compositional profiles similar to schematic profile 

C (fig. 11)—with a baseline composition that is more calcic 

towards rim—carry even stronger evidence for an open-system 

behavior concurrently and (or) preceding the 2004–5 extrusion 

of dacite magma. The only alternate explanation for crystals to 

oscillate compositionally around comparable An minima and 

maxima or around an increasing An baseline with time during 

~30 percent of plagioclase crystallization (fig. 10) appears to 

be decompression-driven crystallization in a heating environment (Blundy and others, 2006). However, the lack of a rim 

overgrowth of high An on all crystals during the last ascent 

of 1,000 m (that is, where this affect would be strongest; fig. 

2a in Blundy and others, 2006) and the observed low An of 

tiny microphenocrysts of this study (for example, figs. 6, 8A) 

rather excludes the “latent-heat” explanation. We therefore 

take the compositional features of profiles B and C (fig. 11) as 

evidence for magma recharge either of hotter and less crystalline dacitic magma of similar bulk composition as the 2004–5 

eruption or of more mafic, Ca-richer magma that has not yet 

been sampled as extruded lava. Dissolution of plagioclase 

from gabbroic xenoliths may help buffer Ca/Na as well, but 

only if the system is sufficiently hot.

Plagioclase View of the Mount St. Helens 

Eruption

On the basis of the discussion above, we infer the following. Subsequent to the last extrusive event in 1986, remaining 

dacite magma began cooling. Crystal growth was intermittent as crystals cycled through the interior, hotter parts of the 

reservoir/conduit system, causing resorption. Temperature 

gradients across the reservoir were possibly maintained by 

minor recharge events of deeper dacite magma, which also 

supplied some plagioclase phenocrysts that were subjected to 

the same subsequent cycles of resorption and growth as the 

resident crystals. Seismic evidence also suggests that recharge 

may have occurred in the middle to late 1990s (Moran and 

others, this volume, chap. 2). Prior to the renewed eruptive 

activity in 2004, a relatively widespread crystallization event 

caused growth and incorporation of acicular orthopyroxene 

inclusions as some magma experienced rapid cooling, possibly 

associated with ascent. Subsequent arrival and mixing of new 

dacite magma remobilized more crystalline parts of the system, 

causing resorption as new magma mixed with resident magma, 

ultimately causing the final ascent to the surface immediately 

or within weeks. During the final ascent, crystals grew euhedral 

rims. We attribute the last growth cycle—in which plagioclase 

of lowest An content incorporates acicular orthopyroxene 

inclusions, is partially resorbed, overgrown with higher An 

content, and subsequently normally zoned—to magmatic processes that triggered the onset of extrusion in October 2004; but 

it predated the eruption by months to weeks. Some (as many as 

three?) of the older, more interior dissolution surfaces, but still 

within the outer ~80 μm of crystals, may postdate the last eruption in 1986 but predate the last growth cycle. Simply zoned 

crystals with continuously decreasing An content (that is, Type 

3 crystals) were carried by recharging dacite, grew largely in 

response to decompression, and thus may track the amount of 

recharged magma. They were stirred into the residing magma 

and, together with already existing phenocrysts, underwent 

crystallization to develop a common growth history during final 

ascent. Earlier recharge events may have carried similar plagioclase but, unless erupted immediately, that plagioclase was 

texturally modified by resorption-and-growth events, thereby 

losing its distinct textural character. This model is compatible 

with results from other petrologic studies of 2004–5 dacite (this 

volume: Pallister and others, chap. 30; Rutherford and Devine, 

chap. 31; Thornber and others, chap. 32; Cooper and Donnelly, 

chap. 36; Reagan and others, chap. 37).

Conclusions

We investigated plagioclase, focusing on near-rim areas 

of phenocrysts in dacite of the 2004–5 eruption, and compared 

those to plagioclase in selected dacite samples from the dome 

that grew in the 1980s. The results of our textural and compositional study are the following:

1. On the basis of distinguishing several plagioclase textural 

types in combination with a newly developed crystal mapping procedure, we show that 2004–5 dacite contains a 

different crystal population than 1980s dacite.

2. Observed An range of plagioclase near rim (≤80 μm) in 

2004–5 dacite is An57–35, and minima correspond with 

compositions of smallest microphenocrysts (≤20 μm) 

at overall crystal contents of ~50 percent. Location of 
34. Plagioclase Populations and Zoning in Dacite of the 2004–2005 Mount St. Helens Eruption 807

a given An composition is largely independent of the 

distance to the rim of crystals, except in texturally simple 

crystals without dissolution surfaces that are normally 

zoned (that is, highest An content in innermost reach and 

lowest An content at the rim). Crystals with dissolution 

surfaces possess normal zonation within bands bounded 

by dissolution surfaces, so that highest An content is the 

immediate overgrowth on a dissolution surface and lowest 

An content within a band is last to crystallize, leading to 

one or several An oscillations toward the rim.

3. Acicular orthopyroxene inclusions in plagioclase are 

characteristic of 2004–5 dacite, and they are consistently embedded in plagioclase with lowest An content. 

Orthopyroxene inclusions have textural counterparts in 

glassy groundmass of a rare 2004 sample and also in 

samples from the 1980s. The orthopyroxene occurrences 

suggest that inclusions in plagioclase track previous magmatic conditions at which acicular orthopyroxene crystals 

were temporarily abundant in interstitial melt.

4. We attribute An oscillations with comparable An minima 

and maxima mostly to changes in crystallinity as induced 

largely by temperature fluctuations in combination with 

recharge to maintain Ca/Na as crystallization generally 

progresses. Plagioclase of lowest An content and with 

acicular orthopyroxene inclusions may record higher 

crystallinity and lower temperature conditions (lower Ca/

Na in melt), whereas plagioclase of highest An content 

overgrowing dissolution surfaces crystals may track conditions of lower crystallinity and higher temperature (higher 

Ca/Na in melt). Convection along a pressure gradient may 

aid the process through ascent-induced crystallization and 

descent-induced resorption. Crystals with normal An zonation and essentially no dissolution surface likely crystallized entirely in a decompressing environment.

5. The assemblage of phenocrysts with one or more growth 

bands composed of similar An range requires localized 

and repetitive processes, including mixing, to juxtapose 

crystals with different growth histories. Crystallization 

of the outermost ~5–15 μm developed during a common 

growth history among phenocrysts.

6. We envision a scenario in which residual 1980s magma 

continued to evolve subsequent to the last extrusion in 

1986 and was maintained by small recharge events. A 

cycle of pronounced cooling/heating/cooling (possibly 

associated with ascent) shortly (months?) before the 

current eruptive activity can be inferred from the characteristic textural features of 2004–5 plagioclase, namely 

acicular orthopyroxene inclusions embedded prior to the 

last dissolution surface, which in turn was overgrown by 

euhedral rims. Our data are compatible with rechargedriven initiation of the current eruption and that recharged 

magma may also be dacitic but poorer in crystals and hotter than resident dacite into which it has been blended.

Acknowledgments

We acknowledge support for the remote-access electron 

microprobe laboratory at Portland State University through 

a National Science Foundation grant, EAR-0320863. M.J. 

Streck acknowledges support from the Eidgenössische 

Technische Hochschule (ETH) Zürich as Gastdozent (guest 

professor) during preparation of this manuscript and thanks 

W. Halter and C. Heinrich at the ETH and H.-J. Massone and 

T. Theye at the University of Stuttgart for their hospitality. 

Reviews by Maggie Mangan and Frank Tepley were very helpful and improved the paper significantly.

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Mandeville, C.W., and Meeker, G.P., 2008a, Chemistry, 

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Thornber, C.R., Pallister, J.S., Rowe, M.C., McConnell, S., 

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Catalog of Mount St. Helens 2004–2007 dome samples with 

major- and trace-element chemistry: U.S. Geological Survey 

Open-File Report 2008–1130, 9 p., with digital database.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 35

1

 Department of Geosciences, 104 Wilkinson Hall, Oregon State University, 

Corvallis, OR 97331

2

 Department of Geosciences, 104 Wilkinson Hall, Oregon State University, 

Corvallis, OR 97331; now at Department of Geoscience, 121 Trowbridge Hall, 

University of Iowa, Iowa City, IA 52242

3

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Trace Element and Pb Isotope Composition of Plagioclase 

from Dome Samples from the 2004–2005 Eruption of 

Mount St. Helens, Washington

By Adam J.R. Kent1, Michael C. Rowe2, Carl. R. Thornber3, and John S. Pallister3

Abstract

We report the results of in-situ laser ablation ICP–MS 

analyses of anorthite content, trace-element (Li, Ti, Sr, Ba, La, 

Pr, Ce, Nd, Eu, Pb) concentrations, and Pb-isotope compositions 

in plagioclase from eight dome-dacite samples collected from 

the 2004–5 eruption of Mount St. Helens and, for comparison, 

from three dome samples from 1981–85. For 2004–5 samples, 

plagioclase phenocrysts range in composition from An30 to 

An80, with the majority An42–An65. With the exception of Li, the 

range of trace-element abundances in plagioclase phenocrysts is 

largely constant in material erupted between October 2004 and 

April 2005 and is broadly consistent with the 1983–85 dome 

samples. Anomalously high Li contents in the early stage of the 

eruption are thought to reflect addition of Li to the upper part 

of the magma chamber immediately before eruption (within 

~1 year) by transfer of an alkali-enriched, exsolved vapor from 

deep within the magma chamber. Other trace elements show 

significant correlations (at >99 percent confidence limits) with 

anorthite content in plagioclase phenocrysts—Ba, light rareearth elements (LREE), and Pb show positive correlations, 

whereas Ti and Sr correlate negatively. Variations in plagioclase-melt partitioning as a function of anorthite content cannot 

explain trace-element variations—in particular predicting 

trends for Ti and Sr opposite to those observed. A simple model 

involving closed-system fractional crystallization of plagioclase 

+ hypersthene + amphibole + oxides largely reproduces the 

observed trends. The model requires no gain or loss of plagioclase and is consistent with the lack of europium anomalies 

in bulk dacite samples. Analytical traverses within individual 

plagioclase phenocrysts support this model but also point to a 

diversity of melt compositions present within the magma storage zone in which plagioclase crystallized.

Plagioclase crystals from gabbronorite inclusions in three 

dacite samples have markedly different trace-element and Pbisotope compositions from those of plagioclase phenocrysts, 

despite having a similar range of anorthite contents. Inclusions show some systematic differences from each other but 

typically have higher Ti, Ba, LREE, and Pb and lower Sr and 

have lower 208Pb/206Pb and 207Pb/206Pb ratios than coexisting 

plagioclase phenocrysts. The compositions of plagioclase from 

inclusions cannot be related to phenocryst compositions by 

any reasonable petrologic model. From this we suggest that 

they are unlikely to represent magmatic cumulates or restite 

inclusions but instead are samples of mafic Tertiary basement 

from beneath the volcano.

Introduction

Samples obtained from the 2004 eruption of Mount St. 

Helens provide an invaluable sample suite for application 

of petrological and geochemical techniques to examine the 

eruption of a silicic volcano. A wide range of approaches are 

detailed in this volume, and they provide valuable insight for 

monitoring of active and erupting volcanoes and for elucidating past eruptive histories on the basis of examination and 

analysis of eruptive materials.

This study reports measurements of trace-element 

abundances and Pb-isotope compositions in plagioclase from 

dome material erupted at Mount St. Helens from October 

2004 to April 2005. The goals of the study are to apply the 

techniques of trace-element and isotope geochemistry to better 
810 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

understand origin and evolution of magma produced by the 

current eruption, as well as to demonstrate the utility of in-situ 

trace-element analyses by using laser-ablation inductively 

coupled plasma mass spectrometry (LA–ICP–MS) in petrological volcano-monitoring applications. The rapidity of this 

technique and relatively simple sample-processing requirements mean that, in the future, LA–ICP–MS could yield traceelement analyses of volcanic products on short time scales, 

providing an additional tool for monitoring anticipated and 

ongoing volcanic eruptions. Rowe and others (this volume, 

chap. 29) report on the application of these same techniques to 

document the chemistry of ash samples.

Samples and Methods

Samples

We have analyzed plagioclase in eight different dome 

samples collected between October 2004 and April 2005 and in 

a smaller number of plagioclase phenocrysts from three samples 

of earlier dome material erupted in September 1981 (SH100), 

May–June 1983 (SH141), and May 1985 (SH187) (table 1). 

Textural and compositional information from plagioclase is 

commonly used to elucidate the compositions of volcanic 

rocks (for example, Pearce and Kolisnik, 1990; Zellmer and 

others, 2003; Triebold and others, 2005; Browne and others, 

2006; Streck and others, this volume, chap. 34). Our decision to 

concentrate on analysis of plagioclase reflects both the ubiquity 

of this mineral in all erupted products (typically ~80 percent of 

all crystalline phases and ~40 percent of the rock as a whole), 

as well as the observed textural diversity of plagioclase in dome 

samples. Plagioclase analyses, in conjunction with information 

on plagioclase-melt partition coefficients, also have the potential 

to act as a monitor of magma chemistry during crystallization 

(for example, Bindeman and others, 1998; Browne and others, 2006). In most of these samples, widespread groundmass 

crystallization makes direct analysis of liquid compositions difficult or impossible. Most dome samples contain little glass for 

analysis and, even where glass is found, it generally occurs only 

in restricted interstitial locations and has been clearly affected 

by late crystallization of groundmass phases. Our data supplement the extensive whole-rock geochemical datasets available 

for these samples (Pallister and others, this volume, chap. 30).

Textural Classification

Studies of the Mount St. Helens dacite have shown 

some textural complexity in plagioclase and other crystalline 

phases (for example, Streck and others, this volume, chap. 

34). However, in general, many of these features are observed 

at scales smaller than the 50–70-μm spatial resolution of the 

laser-ablation analysis used in this study. For this reason we 

have adopted a simplified textural classification for use with 

Sample

No.

Collection dateEstimated 

eruption date Sample type Number of analyses

2004–2005

SH300-1A 10/20/04 <1986 Dacite dome 13

SH304-2A 11/4/04 10/18/04 Dacite dome 351

SH304-2C 11/4/04 10/18/04 Dacite dome 17

Inclusion Gabbronorite 15

SH305-1 1/3/05 11/20/04 Dacite dome 28

SH305-2 1/3/05 11/20/04 Dacite dome 10

Inclusion Gabbronorite 10

SH306-1 1/14/05 12/15/04 Dacite dome 22

Inclusion Gabbronorite 17

SH311-1 1/19/05 01/16/05 Dacite dome 17

SH315-1 4/19/05 04/1/05 Dacite dome 19

1981–85

SH 100 9/8/81 September 1981 Dacite dome 6

SH 141 6/27/83 May–June 1983 Dacite dome 

lobe

8

SH 187 6/6/89 May 1985 Dacite dome 14

1

 Does not include 54 analyses from grain traverses.

2004–5 dome samples. This classification is also applicable to 

grains when polished sample mounts were viewed in reflected 

light. Most plagioclase occurs as euhedral or subhedral, equant 

to tabular phenocrysts less than 2 mm long. These display many 

of the features common in plagioclase phenocrysts from silicic 

magma, such as oscillatory zoning, spongy-textured zones, 

entrapment of melt inclusions, and some breakage of grains due 

to flow processes; but rounding of grains and resorbed zones are 

relatively uncommon. We refer herein to these crystals as plagioclase phenocrysts and interpret them as the result of crystallization of plagioclase directly from melt.

Plagioclase also is present in numerous crystal-rich inclusions in many recent Mount St. Helens lavas (Heliker, 1995). 

These generally are referred to as gabbroic inclusions, although 

many actually have gabbronoritic or noritic compositions and 

may be hornblende rich. Inclusions from the 2004–5 dacite 

appear similar to gabbroic and noritic inclusions from the 

1980–86 dacite documented by Heliker (1995). In the 2004–5 

dome material, inclusions occur in most recovered samples, 

although they are more abundant in some than others (for 

example, SH304-2C). We analyzed three hornblende-bearing 

gabbronorite inclusions from different samples (table 1) for 

this study. In these, plagioclase is relatively coarse (long axis as 

much as several millimeters) and subhedral to anhedral in form. 

Evidence of melt reaction and resorption is common along 

grain boundaries, suggesting that the inclusions are melting or 

reacting with the melt that transports them (Heliker, 1995). In 

thin sections, the inclusions can be seen to be in the process of 

actively disaggregating. Disaggregation results in a third textural 

type of plagioclase that we refer to as disaggregated incluTable 1. Details of samples analyzed for this study.

[Estimates of eruption date are discussed in Pallister and others (this volume, 

chap. 30). Inclusions are from sample listed immediately above.]
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 811

sions. Plagioclase crystals of this type occur isolated within the 

groundmass and are typically much larger than phenocrysts, as 

large as several millimeters, are subhedral or anhedral, and commonly show indications of significant resorption, disruption of 

zoning, and/or growth of new plagioclase on rims.

Analytical Methods

Samples were prepared for analysis by selecting small 

(~5–20 mm) pieces of dome lava from each sample and setting 

these in 25-mm-diameter epoxy mounts. In most cases, material 

was chosen because it was macroscopically representative of 

the larger sample, although in some cases pieces were selected 

to sample coarse-grained gabbroic inclusions. Once mounted in 

epoxy, samples were ground down several millimeters to expose 

the inside of the selected dacite piece by using coarse (240) grit 

paper before polishing with 300–600 grit paper and, finally, 

a 1-μm alumina powder and water slurry. Samples were then 

cleaned for ~5 minutes in distilled water in an ultrasonic bath, 

and mounts were photographed in reflected light and examined 

using backscattered electron imagery before analysis.

Trace-element measurements were made by LA–ICP–MS. 

All measurements were made in the W.M. Keck Collaboratory 

for Plasma Mass Spectrometry at Oregon State University using 

a DUV 193 nm ArF Excimer laser and a VG ExCell quadrupole 

ICP–MS. A general outline of the analytical techniques used for 

this instrument is given in Kent and others (2004a) and Kent and 

Ungerer (2006). Ablation was conducted under a He atmosphere, and He also was used to sweep resulting particulate into 

the ICP–MS at a flow rate of ~0.75 L/min.

During LA–ICP–MS, ~20–25 trace-element and internalstandard isotope masses were monitored (Kent and others, 

2004a). Before each ablation, count rates were measured at each 

mass for 45 s to determine background count rates; these were 

then directly subtracted from the rates measured during ablation 

to account for the instrumental background. Following ablation, 

45 s was allowed to elapse for signal washout before starting 

the next analysis. For all trace-element analyses the laser was 

held stationary relative to the sample so progressive ablation 

produced a circular crater. Ablation rates in silicate glasses and 

minerals are on the order of 0.1–0.2 μm per pulse, and a full 

ablation thus produced a crater 20–30 μm deep and resulted in 

ablation of ~100–400 ng of material. Two different analytical 

approaches were taken. First, in order to document trace-element variations between plagioclase phenocrysts and between 

these and plagioclase in gabbroic inclusions, a large number of 

analyses of plagioclase from multiple samples were made using 

50–80-μm-diameter laser spots. During LA–ICP–MS analysis, 

detection limits depend largely on the rate at which material is 

removed, and these spot sizes provided sufficiently low detection limits to enable analysis of a range of trace elements (see 

below). Samples underwent ablation for 30–40 s with a laser 

pulse rate of 4 Hz. In addition, for sample SH304-2A a number 

of selected grains were analyzed along core-rim traverses using 

a smaller spot size (20 μm) and 7-Hz pulse rate. The smaller 

spot size allowed consistent detection of only the most abundant trace elements (Li, Ti, Sr, Ba, La, and Ce), but it provided 

higher spatial resolution to allow study of trace-element variations during progressive plagioclase crystallization. Plagioclase 

textures and major-element compositions of the same crystals 

analyzed in the trace-element traverse also were investigated 

using techniques similar to those detailed in Rowe and others 

(this volume, chap. 29) with backscattered electron imagery and 

wavelength-dispersive analysis using a Cameca SX-100 electron 

microprobe (EMPA).

Before calculation of trace-element abundances, background intensities for each mass were subtracted directly from 

those measured during ablation, and signals during ablation 

were only considered to be above background if they were 

greater than the background count rate plus three standard 

deviations (calculated from counting statistics). Estimates of 

minimum detection limits are: <1.5 μg/g for Ti and Li; <0.4 

μg/g for Ba, and Rb; <0.1 μg/g for V, Sr, and Nd; and <0.05 

μg/g for Y, Zr, Nb, Cs, La, Ce, Pr, Sm, Eu, Pb, Th, and U. In 

general, only Li, Ti, Sr, Ba, the light rare-earth elements (LREE: 

La, Ce, Pr, Nd), Eu, and Pb were consistently detectable, and 

we have restricted our study to these elements. Trace-element 

abundances were calculated with reference to NIST 612 glass 

with 29Si as the internal standardizing isotope. Calculation of 

trace-element abundances required independent knowledge of 

the average SiO2 content of each ablation volume, although in 

zoned plagioclase this may vary on spatial scales smaller than 

the laser spot size. For this reason, SiO2 contents were determined directly from LA–ICP–MS analysis by exploiting the 

stoichiometric relation between anorthite content, CaO/SiO2, 

and SiO2 contents in plagioclase. Measured 43Ca+/

29Si+

 ratios 

from plagioclase were converted to CaO/SiO2 ratios by using 

the measurement of NIST 612 and application of equation 1:

, (1)

where

is the calculated CaO/SiO2 of the unknown

 plagioclase (with oxides in weight percent),

 is the known CaO/SiO2 of NIST 612 (we

 used 0.1594; Pearce and others 1997),

is the measured 43Ca+/

29Si+

 ratio from

 NIST 612, and

is the measured 43Ca+/

29Si+

 ratio in the

 unknown plagioclase.
812 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

We then calculated anorthite and SiO2 contents of plagioclase 

from measured CaO/SiO2 using equations 2 and 3:

, (2)

and

. (3)

Equations 2 and 3 are based on empirical relations observed for 

Mount St. Helens plagioclase from electron microprobe analyses (M.C. Rowe and A.J.R. Kent, unpub. data, 2005) and are 

close to the those expected from stoichiometric considerations.

Results of measurement of USGS BCR-2G standard glass 

using the above protocol are shown in appendix 1. We also 

show CaO/SiO2

 ratios for these glasses calculated from equation 1, and these ratios are within uncertainty of the reported 

composition of this glass. In general, measured values agree 

closely with accepted values, and nearly all elements are 

within ± 5–10 percent of the accepted values.

One test of this analytical procedure is the comparison 

between anorthite contents measured by EMPA and LA–ICP–

MS along the traverses in five plagioclase grains in sample 

SH304-2A (figs. 1, 2; appendix 2). Overall, there is good 

agreement between the two techniques, and where differences 

are apparent, they are most likely because the analyzed volume 

for LA–ICP–MS is significantly greater than the analyzed volume for EMPA. On the basis of reproducibility of CaO/SiO2

ratios in BCR-2G, we believe that uncertainties in measured 

anorthite contents are ±5 percent mole fraction. Overall, we 

estimate uncertainties in trace-element measurements in plagioclase as ≤10 percent (at 2σ) for Sr, Ti, Ba, La, and Ce and 

≤15–20 percent for Nd, Eu, Pb, and Li. Spot compositions of 

other Mount St. Helens plagioclase from 1981–85 and 2004–5 

are given in appendix 3.

Lead isotope compositions of plagioclase phenocrysts, 

plagioclase in gabbroic inclusions, and groundmass were 

analyzed by laser-ablation multicollector ICP–MS (LA–

MC–ICP–MS), using a NuPlasma multicollector ICP–MS 

at Oregon State University. Measurements were made using 

the same laser systems described above and using 80–100 

μm laser spot size, pulse rate of 10–15 Hz, and a lateral 

translation rate of 5 μm/s. Individual measurements involved 

from three to five blocks, with each block consisting of two 

separate 10-s measurements, generally resulting in ablation along a track length of 300–500 μm. For measurement 

of groundmass composition, a relatively finely crystallized 

region was chosen and the laser spot simply was translated 

across the region while using the same ablation conditions. 

Although this produced some variation in signal intensity, the 

generally higher Pb content of this material resulted in data of 

relatively high quality. For most samples, signal intensity was 

too low (<<100 mV total Pb) to enable sufficiently precise 

measurement of the minor isotope 204Pb, and thus we report 

only 208Pb/206Pb and 207Pb/206Pb ratios (appendix 1).

All ion beams were measured using Faraday collectors. 

Instrument mass bias was corrected by frequent measurement 

of NIST 612 and NIST 610 glass and by using the measured 

208Pb/206Pb ratio to apply an exponential mass-bias correction 

to Pb-isotope ratios measured in unknown samples. Ratios of 

2.1694 and 2.1651 for 208Pb/206Pb were measured by Baker and 

others (2004) in NIST 610 and NIST 612 and were used as the 

correct composition of this glass. Backgrounds were corrected 

by on-peak zero measurements for 30 s without the laser firing 

prior to analysis, with measured signals directly subtracted from 

signals measurement during ablation. Precision for individual 

analyses is strongly dependent on signal size and Pb abundance but is generally better than 0.2 percent for 208Pb/206Pb and 207Pb/206Pb ratios. Analysis of the BCR-2G glass standard gave 

results that are well within uncertainty of measurements made 

using solution techniques (Paul and others, 2005).

Results

Measured anorthite content for all plagioclase types ranges 

between An30 and An80, with most between An42 and An68 (figs. 

2, 3). Anorthite contents of plagioclase phenocrysts and those 

present within gabbroic inclusions show considerable overlap, 

although inclusions appear to show a somewhat more restricted 

range of compositions and may range to more anorthite-rich 

compositions (for example, grain 1 in fig. 2). Significant variations are evident in trace-element abundances in plagioclase, 

even where these are from the same sample, although the range 

of trace-element abundances from plagioclase phenocrysts 

from each sample is broadly similar, with the exception of Li. 

Lithium contents for dome samples erupted in October and 

November 2004 are consistently higher than those from subsequently erupted material (fig. 4; Kent and others, 2007; Rowe 

and others, this volume, chap. 29). Our results for plagioclase 

phenocrysts are indistinguishable from the composition of a 

smaller number of samples from the 1981–86 dome complex 

(fig. 3) and from sample SH300-1A, which is also considered 

to be a part of the 1980s dome complex pushed ahead of new 

erupting magma (Pallister and others, this volume, chap 30).

A correlation matrix for anorthite and trace-element 

abundances in plagioclase phenocrysts from all 2004–5 

samples (a total of 138 analyses) shows significant correlations between anorthite content and the abundances of all trace 

elements other than Li in plagioclase phenocrysts (table 2; fig. 

3): Sr and Ti are positively correlated and Ba, LREE, Eu, and 

Pb are negatively correlated with anorthite. For this relatively 

large number of analyses, correlation coefficients >0.17 and 

>0.23 are significant at the 95-percent and 99-percent confidence levels, respectively. Lithium shows no significant 

correlations with other elements, whereas Ba, LREE, and Pb 

are strongly positively correlated with each other. Strontium 

shows relatively poor correlations with Ti and La but signifi-
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 813

Grain 3

250 μm

250 μm 250 μm

250 μm

250 μm

Grain 2

Grain 1

Grain 4

Grain 5

cant correlations with Ba, Eu, and Pb. Titanium shows significant negative correlations with Ba, LREE, Eu, and Pb.

There are clear differences in trace-element abundances 

between plagioclase phenocrysts and plagioclase from gabbroic inclusions. Plagioclase from inclusions typically has 

higher Ti, Ba, LREE, and Pb and lower Sr contents than 

phenocrysts with the same anorthite contents, and there also 

are some systematic differences between different inclusions 

(figs. 3, 4).

In general, analyses of groundmass and phenocrysts 

typically show similar Pb isotope composition, whereas 

plagioclase from gabbronorite inclusions have significantly 

different Pb-isotope compositions with lower 208Pb/206Pb and 207Pb/206Pb (fig. 5; appendix 1). In cases where groundmass 

was analyzed directly adjacent to disaggregating and reacting 

inclusions (for example, SH304-2C), groundmass compositions appear to lie along mixing lines between inclusions and 

groundmass from other samples.

Figure 1. Backscattered electron images of five plagioclase grains from Mount St. Helens sample SH304-2A analyzed for 

major- and trace-element abundances along traverses. White line in each photograph shows location of analytical traverse.
814 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

200

400

600

0.2

0.4

0.6

0.8

304-2A

Grain 1

304-2A

Grain 2 

0.2

0.4

30

50

70

800

1,000

1,200

50

150

250

2

6

10

0.2

0.4

0.6

0.8

0.2

0.4

30

50

70

800

1,000

1,200

50

150

250

2

6

10

Ce (μg/g) Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN FeO* (wt. %)

Ce (μg/g) Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN FeO* (wt. %)

DISTANCE FROM RIM, IN MICROMETERS DISTANCE FROM RIM, IN MICROMETERS

0

(RIM) (RIM)

100 200 300 400 500 600 700 0 100 200 300 400 500

EMPA

LA-ICP-MS

50

100

150

200

Figure 2. Mole-fraction anorthite (XAN), FeO, Li, Ti, Sr, Ba, and Ce concentrations measured along traverses in five 

plagioclase grains from Mount St. Helens dacite sample SH304-2A (fig. 1). Symbol width for LA–ICP–MS data is the same size 

as the laser spot used. 0 is rim. All data except FeO are in appendix 2.

Discussion

Compositional Variations Among Plagioclase 

Phenocrysts

Plagioclase phenocrysts from the 2004–5 dome samples 

show considerable variation in trace-element compositions, 

with abundances ranging by factors of ~3–5 over the range 

of anorthite contents recorded. Even at constant anorthite 

there is a factor of ~2–3 variation in trace-element abundances (fig. 3), which is far outside analytical errors. With the 

exception of Li, there is no systematic variation in the range 

of trace-element (fig. 4) or anorthite content in plagioclase 

phenocrysts over the course of the eruption, consistent with 

the lack of variation in whole-rock compositions (Pallister 

and others, this volume, chap. 30). In addition, there is no 

clear difference in trace-element composition of plagioclase 
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 815

0 100 200 300 400 0 100 200 300

304-2A

Grain 3

304-2A

Grain 4

50

100

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200

0.2

0.4

0.6

0.8

0.2

0.4

30

50

70

800

1,000

1,200

50

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200

2

6

10

50

100

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200

0.2

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30

50

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800

1,000

1,200

1,400

50

150

250

2

6

10

DISTANCE FROM RIM, IN MICROMETERS DISTANCE FROM RIM, IN MICROMETERS

(RIM) (RIM)

Ce (μg/g) Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN FeO* (wt. %)

Ce (μg/g) Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN FeO* (wt. %)

Figure 2.—Continued.

phenocrysts analyzed from 1981–85 dome samples and the 

2004–5 samples (fig. 3), although there are suggestions that 

these samples originate from distinct magma batches (for 

example, Pallister and others, this volume, chap. 30; Blundy 

and others, this volume, chap. 33).

Abundances of trace elements other than Li show significant negative (Ba, LREE, Pb) or positive (Sr, Ti) correlations 

with anorthite content (table 2; fig. 3). One explanation for this 

is that the correlations simply reflect changes in plagioclasemelt partitioning with changing anorthite content (Blundy 

and Wood, 1994; Bindeman and others, 1998). However, 

although calculated plagioclase compositions (determined 

using the partitioning models in the preceding references 

and the bulk composition of 2004 dacite as a proxy for melt 

composition) are a reasonable match for the more incompatible elements (REE, Ba and Pb), the predicted trends for 

Sr and Ti are reverse to those evident in plagioclase phenocrysts (fig. 3).
816 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

304-2A

Grain 5

0 100 200 300

Ce (μg/g) Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN FeO* (wt. %)

50

100

150

200

0.2

0.4

0.6

0.8

0.2

0.4

20

40

60

80

800

1,000

1,200

50

150

250

2

6

10

DISTANCE FROM RIM, IN MICROMETERS

(RIM)

Figure 2.—Continued.

The simplest explanation for this observation is that 

decreases in Ti and Sr reflect removal of these elements from 

the melt during progressive closed-system crystallization of 

phenocryst phases. Because formation of crystals will sequester elements according to their partition coefficients, we do not 

need subsequent separation of crystals and liquid (as required 

to change the bulk composition) in order to change the composition of residual melt and subsequent crystals. The lack of 

europium anomalies in bulk dacite compositions (Pallister and 

others, this volume, chap. 30), together with simple calculations based on the Nd/Eu ratios of bulk dacite and feldspar, 

suggest that the 2004–5 dacite has lost <<5–10 percent 

plagioclase. The modal abundances of plagioclase and other 

phenocrysts observed in dacite samples thus probably approximate the overall proportions in which they have crystallized, 

a view broadly consistent with experimental investigations 

(Rutherford and Devine, this volume, chap. 31).

We have calculated changes in plagioclase composition using a simple model involving progressive crystallization of the observed phenocryst phases in the proportions in 

which they occur within the dacite from a melt with an initial 

composition that is the same as the bulk composition of dome 

dacite (fig. 6, table 3). Of the trace elements analyzed, Sr and 

Ti are the most compatible in the crystallizing assemblage of 

plagioclase, hypersthene, hornblende, and magnetite-ilmenite, 

with Sr entering plagioclase and Ti entering hornblende and 

ilmenite. We also note that our calculations are unlikely to 

fully explain all variations, given that we assume constant 

phase proportions and starting melt composition and that no 

accounting is made for differences in temperature, pressure, 

or reequilibration between plagioclase and melt (which may 

be significant for fast-diffusing elements such as Sr, according to Cherniak and Watson, 1994). However, the calculations 

do suggest that positive trends between anorthite contents and 

Sr and Ti, at least below ~An60, could result from sequestration of these elements from melt during phenocryst growth 

and thus are linked to total crystallinity. Strontium/barium 

ratios in plagioclase also depart significantly from the trends 

expected from partitioning and are also largely replicated by 

our model. The negative correlations with anorthite shown 

by more incompatible elements (Pb, La, Ba) also are reproduced by our model, although these correlations do not differ 

significantly from those predicted by partitioning alone (fig. 

3). Overall, we suggest that crystallization of the observed 

phenocryst phases, together with the control of anorthite content on crystal-liquid partitioning, exerts strong control over 

plagioclase trace-element abundances.

Control of mineral compositions by closed-system 

equilibrium crystallization is recognized in other crystal-rich, 

silicic magma systems (for example, Zellmer and others, 2003; 

Treibold and others, 2006), including Mount St Helens. Blundy 

and others (this volume, chap. 33) also suggest that closedsystem crystallization controls glass inclusion compositions 

in 1980–86 Mount St. Helens lavas. Streck and others (this 

volume, chap. 34) argue that increases in crystallinity are the 

primary control on anorthite zoning in the outer ~80 μm of 

plagioclase phenocrysts in the 2004–5 eruption. Moreover this 

simple model does not necessarily conflict with more complex 

crystallization histories, as long as we consider trends shown 

by plagioclase to represent the range of histories experienced 

by individual plagioclase grains. Note that we do not visualize a simple “freezing” model in which a single homogenous 

melt crystallizes progressively until the current crystallinity is 

reached. Rather, as suggested by zoning and resorption features 
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 817

50

0.2 0.3 0.4 0.5 0.6 0.7 0.8

100

150

200

10

20

30

40 10

8

6

4

2

2.0

1.5

1.0

0.5

12

10

8

6

4

2

100

200

300

400

500

500

1,000

1,500

r = 0.52

r = 0.37

r = -0 .76

r = -0 .61

r = -0 .63

r = -0 .76

Pb (μg/g) Eu (μg/g) La (μg/g)

Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

XAN

0.2 0.3 0.4 0.5 0.6 0.7 0.8

XAN

2σ

2σ

2σ 2σ

2σ

2σ

2σ

Phenocrysts Gabbroic inclusions

EXPLANATION

Partitioning model

SH304-2A SH304-2C

SH304-2C

SH305-1

SH305-2

SH306-1

SH311-1

SH315-1

SH300-1

1981–85

BW ‘94 KD

Max, min

B ‘98 KD

Max, min

SH305-2

SH306-1

Figure 3. Trace-element abundances and anorthite content measured in plagioclase samples from Mount St. 

Helens dacite of 2004–5. Data shown for plagioclase present as phenocrysts and within gabbroic inclusions. Thick 

dashed red lines show linear least-squares fit to phenocryst data, with the value of the correlation coefficient r

listed. No correlation shown for Li. Solid red and black lines show predicted elemental abundances in plagioclase 

calculated using plagioclase-melt partitioning models of Bindeman and others (1998; B ’98) and, for Sr and Ba, 

Blundy and Wood (1994; BW ’94), and the trace-element abundances measured in dacite sample SH304-2A 

(Pallister and others, this volume, chap. 30) as the estimated bulk melt composition (see text for explanation). 

All partition coefficients calculated at 850°C. Results for three dome samples (SH100, SH141, and SH187) from 

1981–85 eruptions shown for comparison. Representative error bars (2σ) shown for each plot.
818 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

0.5

1.0

1.5

0

2

4

6

8

0

2

4

6

8

10

0

10

20

30

EXPLANATION

October 18 and later

Phenocrysts

Inclusions

SH300-1 (1985?)

DAYS AFTER OCTOBER 1, 2004

0

50

100

150

200

0

0 50 100 150 200 250

DAYS AFTER OCTOBER 1, 2004

0 50 100 150 200 250

50

100

150

0

400

800

1,200

1,600

Pb (μg/g) La (μg/g) Eu (μg/g)

Ba (μg/g) Sr (μg/g) Ti (μg/g) Li (μg/g)

Anorthite Li Ti Sr Ba La Eu Pb

Anorthite 1 -0.04 0.52 0.37 -0.76 -0.61 -0.63 -0.76

Li -0.04 1 0.14 -0.05 0.06 0.06 -0.04 -0.05

Ti 0.52 0.14 1 0.07 -0.23 -0.38 -0.31 -0.57

Sr 0.37 -0.05 0.07 1 -0.23 -0.11 -0.20 -0.24

Ba -0.76 0.06 -0.23 -0.23 1 0.83 0.80 0.69

La -0.61 0.06 -0.38 -0.11 0.83 1 0.74 0.66

Eu -0.63 -0.04 -0.31 -0.20 0.80 0.74 1 0.68

Pb -0.76 -0.05 -0.57 -0.24 0.69 0.66 0.68 1

Table 2. Correlation matrix for anorthite and selected trace elements for plagioclase phenocrysts from 

2004–2005 Mount St. Helens dome samples.

[Values shown are the calculated linear correlation coefficient r. For the number of analyses (n = 138) |r | > 0.17 is significant at 

95-percent confidence, |r | > 0.23 is significant at 99-percent confidence.]

Figure 4. Average trace-element compositions of plagioclase phenocrysts and within gabbronorite inclusions from 

Mount St. Helens dacite collected between October 2004 and April 2005 as a function of estimated eruption date 

(Pallister and others, this volume, chap. 30). Symbols represent average composition ±1σ.
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 819

SH304-2C Groundmass

SH304-2C Inclusion

SH304-2C Phenocrysts

SH304-2A Phenocrysts

SH306-1 Phenocrysts

SH306-1 Inclusion

SH306-1 Groundmass

SH304-2A Groundmass

2.056

2.052

2.048

2.044

2.040

2.036

2.032

0.820

207Pb / 206Pb

EXPLANATION

208Pb / 206Pb

0.824 0.828 0.832

Figure 5. Measured 208Pb/207Pb and 207Pb/206Pb ratios of 

groundmass and plagioclase from phenocrysts and inclusions 

in three samples of 2004–5 Mount St. Helens dacite. Error bars 

represent 2σ error.

in plagioclase and other phenocrysts (see below; Rutherford and 

Devine, this volume, chap. 31; Streck and others, this volume, 

chap. 34), plagioclase phenocrysts experienced a range of pressure and temperature conditions and coexisted with a range of 

melt compositions within a convecting and self-mixing magma 

chamber (Couch and others, 2001; Rutherford and Devine, this 

volume, chap. 31). We discuss this further below.

Departures of Li and Eu in plagioclase from the predicted equilibrium partitioning values (fig. 3) are related 

to other factors. Variations in Li contents are discussed in 

further detail below and in Kent and others (2007). For Eu, 

the plagioclase-melt partition coefficients used in figure 3 

are from the study of Bindeman and others (1998) and were 

measured from experiments conducted in the presence of 

atmospheric oxygen. In this case all Eu was present in the trivalent form, and thus Eu behaves consistently with the other 

moderately incompatible middle REE (KD = ~0.1). In most 

volcanic systems, conditions are more reducing, significant 

amounts of more compatible divalent Eu are present, and Eu 

typically is more compatible in plagioclase than other REE 

(for example, Rollinson, 1993). If we use a general partition 

coefficient value of 2 for the plagioclase-melt partition 

coefficient (Rollinson, 1993), then we calculate more reasonable plagioclase Eu contents of 1–2 μg/g.

Compositional Variations Within Plagioclase 

Crystals

Anorthite, FeO, and trace elements were measured along 

traverses in selected plagioclase phenocrysts from SH304-2A 

to investigate changes in trace-element composition during 

progressive crystallization (figs. 1, 2). The grains chosen represent the general classes of phenocryst types recognized by 

Rutherford and Devine (this volume, chap. 31) and Streck and 

others (this volume, chap. 34).

The most common compositional feature of plagioclase 

phenocrysts is cyclic zoning, starting with an abrupt increase in 

An contents, sometimes also in conjunction with indications of 

mineral dissolution or resorption (for example, grain 5 in fig. 

1), followed by more gradual decrease in An content until the 

start of the next cycle. Variations in the composition, number, 

and width of cycles are apparent in individual phenocrysts but, 

in general, the first plagioclase is ≥An50, decreasing outward to 

~An40–20. This pattern recurs frequently; grain 4 shows multiple cycles, and grains 2, 3, and 5 show one or two cycles. 

Such cyclic compositional variation is common in plagioclase 

phenocrysts from the 2004–5 Mount St Helens dome, although 

there are subtle differences of opinion regarding their origin. 

Rutherford and others (this volume, chap. 31) suggest that 

these cycles represent plagioclase crystallizing during cycles 

of convection from deeper, hotter parts of the magma chamber 

to shallower and cooler conditions, whereas Streck and others 

(this volume, chap. 34) argue that anorthite cycles in the outer 

80 μm of phenocrysts largely represent changes in crystallinity 

rather than changes in external conditions.

Although LA–ICP–MS analyses do not have sufficient 

spatial resolution to examine individual anorthite cycles in 

detail, trace-element abundances measured along traverses 

through individual grains can be examined in light of the 

variations evident between phenocrysts (figs. 3, 6). We have 

used the covariation between An contents and trace-element 

abundance as the basis for comparisons (fig. 7).

With some exceptions, trace-element variations within 

individual phenocrysts broadly follow the same general trends 

evident between plagioclase phenocrysts (fig. 2). Thus, within 

individual grains, Ti and Sr are positively correlated with 

anorthite, and An contents are broadly anticorrelated with 

Ba and La (for example, grains 2, 5). This relation also supports a model in which largely closed-system crystallization 

controls trace-element composition in coexisting liquid and 

subsequently formed plagioclase. However, in detail, individual 

crystals define separate subparallel trends (fig. 7). For example, 

Sr contents in grain 1 form a linear trend with positive slope 

that is offset to lower Sr contents at a given An than are similar 

trends in grains 2 and 4. In some cases, as with grain 4, analyses 

from a short segment of the traverse lie off the trend defined 

by the other analyses—specifically, the two outermost points 

measured on the rim have lower Sr and slightly higher Ti than 

the interior of the grain (figs. 2, 7). It is unlikely that changes in 

element partitioning related solely to anorthite or temperature 
820 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0

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5

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35 Fractionation Model

BW ‘94 K D

Max, Min

B ‘98 K D

Max, Min

Partitioning

Plagioclase phenocrysts

SH304-2A

SH304-2C

SH305-2

EXPLANATION

Pb (μg/g) La (μg/g)

Ba (μg/g)

Sr/Ba

Sr (μg/g) Ti (μg/g)

XAN XAN

Figure 6. Comparison of selected trace elements and Sr/Ba ratios in plagioclase phenocrysts with 

trends predicted by a simple closed-system crystallization model. Only Mount St. Helens samples 

SH304-2A, SH304-2C, and SH305-2 are shown, for clarity, although these are broadly representative of 

other 2004–5 dome samples. Potential liquid lines of descent were calculated by incrementally removing 

plagioclase, hornblende, hypersthene, and oxides (magnetite and ilmenite in ratio 1: 5) from an initial melt 

of bulk composition similar to SH304-2. All phases were removed in the proportions in which they occur 

as phenocrysts within the dacite (table 3). Trace-element contents of ferromagnesian and oxide phases 

are given in table 3 and are based on preliminary analyses of these phases by electron microprobe (for 

Ti) and LA–ICP–MS. Note that the calculation is not particularly sensitive to variations in ferromagnesian 

and oxide compositions owing to the low modal proportions of these phases and their generally low traceelement contents. The exception is Ti, which is relatively abundant in hornblende (TiO2

 ~2–3 wt. percent) 

and in ilmenite (TiO2

 ~50 wt. percent). For every 2-percent increment of crystallization, trace-element 

abundances of equilibrium plagioclase were calculated using the partitioning models of Bindeman and 

others (1998; black lines) and, for Sr and Ba, Blundy and Wood (1994; red lines). All partition coefficients 

were calculated at 850°C, consistent with current estimates of magmatic temperatures for 2004–5 

(Pallister and others, this volume, chap. 30), although the model is not significantly different at higher 

temperatures. Uncertainties in partition coefficients shown by dashed lines associated with predicted 

trends from each partitioning model. Because the anorthite content of crystallizing plagioclase is dictated 

by several factors (for example, volatile contents, pressure, temperature, and composition), we chose to 

crystallize the first plagioclase at An70 composition and to have the anorthite content decrease linearly 

and progressively until the final plagioclase has a composition of An35 after crystallization of 54 weight 

percent, the total amount of phenocrysts observed in sample SH304-2A on a void-free basis (Pallister 

and others, this volume, chap. 30). The change in Sr/Ba ratio predicted by partitioning (using the partition 

coefficients of Bindeman and others, 1998) is shown by the purple line.
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 821

Phase Crystal

proportion Ti Sr Ba La Ce Eu Pb

Hypersthene 0.10 900 10 1 1 2 0.1 0.1

Hornblende 0.06 30,000 50 70 1 2 0.3 1.7

Oxide 0.02 100,000 0 0 0 0 0 0

Plagioclase 0.82

Starting magma 4,000 400 336 15 30 1.3 9.8

Table 3. Mineral and magma compositions used for crystal-fractionation model shown in figure 6.

[Column labeled Crystal proportion is that observed in dacite sample SH304-2A (Pallister and others this volume, chap. 

30), calculated on a groundmass-free basis and not including plagioclase microphenocrysts in groundmass. For oxide 

minerals, Ti content represents crystallization of magnetite:ilmenite in ratio of ~5:1 (Pallister and others, this volume, 

chap. 30). Not shown are plagioclase trace-element contents, which were calculated separately for each increment of 

plagioclase crytallized from the coexisting melt composition, using the partitioning relations of Bindeman and others, 

(1998). Starting magma composition based on whole-rock analyses of 2004–2005 dome material (Pallister and others, 

this volume, chap. 30).]

variations are sufficient to produce the variations found between 

crystals. We suggest that plagioclase compositions also reflect 

variations in the composition of coexisting melt from which 

plagioclase crystallized. Although compositional variations 

could be inherited partly from the mid or lower crustal source 

of the dacitic magmas, the near-constant composition of 2004–5 

bulk dacite samples argues against large source-derived variation (Pallister and others, this volume, chap. 30). We suggest 

instead that localized variations in P and T, crystallinity, and 

proportions of crystallizing phases within the magma storage 

zone result in localized variations in melt composition. Crystals 

reflect the composition of localized melt(s) from which they 

grow, and thus individual crystals record variable composition 

trends. Occasional juxtaposition of crystals with new melt compositions, possibly related to convective stirring, also results in 

crystals with compositions that depart from well-defined trends 

on variation diagrams, as seen in Sr and Ti in grain 4. We note 

that melt inclusions from 1980–86 and 2004–5 samples also 

show large differences in Ba and REE abundances at similar 

SiO2

, consistent with diverse melt compositions during mineral 

growth (Blundy and others, this volume, chap. 33).

Grain 1 is slightly different from the other grains analyzed. It is part of a large, complex plagioclase glomerocryst 

and has a complex anorthite-rich core (as high as An75) with 

an adjacent low-anorthite central part and then two cycles of 

An55–60 to An38–40 (fig. 2). This grain has Ti contents as high as 

~600 μg/g in its inner region (from ~200 to 700 μm from the 

rim), which are considerably higher than the Ti contents of 

other plagioclase phenocrysts (fig. 3). High Ti suggests that the 

core of this crystal might have been derived from a disaggregated gabbroic inclusion, although other trace elements do not 

have inclusion-like compositions. In particular, Sr abundances 

in the core of this grain are ~1,000–1,200 μg/g, unlike the ~500 

μg/g contents seen in plagioclase within gabbroic inclusions 

(fig. 3). Although the origin of this grain is uncertain, the high 

Sr might reflect further heterogeneity among inclusions (fig. 3), 

or it could be due to reequilibration of Sr in plagioclase with 

melt after disaggregation, with more slowly diffusing Ti not yet 

reequilibrated (compare with Zellmer and others, 2003). The 

latter explanation requires a residence time of several thousand 

years to reequilibrate over distances of hundreds of microns.

Lithium Variations in Plagioclase Phenocrysts

Variations in the Li contents in 2004–5 dome and ash 

samples are of particular interest and have been discussed 

in detail by Kent and others (2007) and Rowe and others 

(this volume, chap. 29). As shown in figure 4, Li contents in 

plagioclase phenocrysts erupted at the onset of dome extrusion 

are anomalously high (samples SH304-2A, SH304-2C). These 

concentrations are higher, by about a factor of two, than those 

seen in subsequently erupted material and in sample SH300-1 

(interpreted to be a remnant of 1980s dome material), and 

this difference is significant to >99 percent confidence. The 

increase in Li in October–November 2004 samples mirrors 

that reported by Berlo and others (2004) in plagioclase from 

the 1980 eruption of Mount St. Helens, in which high Li 

contents in plagioclase phenocrysts erupted prior to May 18 

and cryptodome material erupted May 18 are thought to reflect 

vapor-phase transport of alkali metals during degassing from 

deep within the magma storage zone before eruption (Berlo 

and others 2004; Blundy and others, this volume, chap. 33). 

Kent and others (2007) suggest a similar model to explain the 

2004–5 material. Although Li contents in the bulk magma 

from October and November 2004 are not elevated (Pallister 

and others, this volume, chap. 30), this is probably as a result 

of vapor loss during late shallow (< ~10 MPa) degassing 

(Kent and others, 2007). Lithium diffusion in plagioclase is 

slower than in melt, and thus plagioclase preserves the high 

Li signature, provided decompression and cooling are sufficiently rapid. This model is supported by the measurement of 

Li concentrations as high as 207 μg/g in melt inclusions from 

SH304-2A (Kent and others, 2007) and in Li-loss profiles 

evident in the outer 100–300 μm of plagioclase transects (fig. 

2). High Li signatures also are evident in ~20 percent of pla-
822 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Figure 7. Ti, Sr, Ba, and La versus mole fraction anorthite 

(XAN) content measured along plagioclase traverses (see figs. 

1, 2). Fractionation models from figure 6 are shown as black 

and red lines (see fig. 6 caption for details). Fields show the 

range of compositions measured in plagioclase phenocrysts 

(gray field, dashed line) and gabbroic inclusions (blue field, 

solid line), from figure 3.

gioclase in ash from explosive eruptions on October 1–5, 2004, 

as noted by Rowe and others (this volume, chap. 29), who 

suggest that these explosions were phreatomagmatic in origin. 

Finally, diffusion modeling based on differences between Li 

contents in plagioclase phenocrysts and gabbroic inclusions 

in SH304-2C (fig. 4) suggests that Li enrichment occurred 

relatively recently—probably within a year of eruption (Kent 

and others, 2007). One plausible scenario is that the accumulation and phase separation of vapor in the apical part of the 

shallow Mount St. Helens magma chamber (~5 km) resulted in 

increased fluid pressure before eruption. Rupturing of wallrocks in late September 2004, perhaps induced by increases 

in fluid pressure, resulted in loss of the low-density vapor and 

reequilibration between a high-density, Li-bearing vapor phase 

and magma. After upward movement of magma commenced 

in late September or early October, Li-enriched apical magma 

contributed to initial phreatomagmatic explosions in early 

October (Rowe and others, this volume; chap. 29) and was the 

first material erupted once extrusion commenced. Comparison 

with eruption rates suggests that Li-enriched magma represents 

~15–20 percent of the volume of the total material erupted.

Composition and Origin of Gabbronorite 

Inclusions

Mafic plutonic inclusions are a common feature of Mount 

St. Helens lavas from the last ~3,000 years and are relatively 

common in 2004–5 dome samples (Pallister and others, this 

volume, chap. 30). The potential sources for mafic inclusions 

include crystalline cumulates and crystal-rich wallrock zones 

(Heliker, 1995; Cooper and Donnelly, this volume, chap. 36) 

and fragments of basement rock removed from conduit walls 

and transported within the magmatic system. One key issue is 

whether these inclusions represent parts of the same magmatic 

system responsible for generation and transport of Mount St. 

Helens magma or represent unrelated rocks from the Tertiary 

Cascade crust beneath the volcano. Heliker (1995) argued 

that the former might be the case. However, preliminary U-Pb 

zircon ages of ~25 Ma from three gabbronorite inclusions from 

the 1980–86 dome complex (Pallister and others, this volume, 

chap. 30) suggest that these are derived from Tertiary basement incorporated into the Mount St. Helens magmatic system, 

and trace-element and Pb-isotope analyses of inclusions and 

plagioclase phenocrysts support this model. Plagioclase crystals 

analyzed from three gabbronorite inclusions have significantly 

different trace-element compositions than coexisting plagioclase 

phenocrysts, with lower Sr and generally higher Ti, Ba, REE, 

and Pb (fig. 3). There are also small but consistent differences 

among inclusions, suggesting that some heterogeneity exists in 

their source (figs. 3, 4). Differences in trace-element abundances 

between phenocrysts and inclusions suggest that inclusions have 

equilibrated with a melt of significantly different composition 

than that from which coexisting phenocryst phases crystallized. 

Overall, the compositions of plagioclase in gabbroic inclusions 

0

400

800

1,200

0

50

100

150

200

250

0

200

400

600

800

Ti, IN MICROGRAMS

PER GRAM

Sr, IN MICROGRAMS

PER GRAM

Ba, IN MICROGRAMS

PER GRAM

La, IN MICROGRAMS

PER GRAM

2

0

0.2 0.3 0.4 0.5 0.6 0.7 0.8

4

6

8

10

XAN

Inclusions

Phenocryst

cores SH304-2A

EXPLANATION

Grain 1

Grain 2

Grain 3

Grain 4

Grain 5
35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 823

are difficult to relate to those of phenocrysts and to bulk magma 

composition by any reasonable petrologic model.

The Pb-isotope compositions of plagioclase from inclusions also are consistently different from those of plagioclase 

phenocrysts and groundmass (fig. 5), with inclusions having 

lower 208Pb/206Pb and 207Pb/206Pb than groundmass and phenocrysts. This difference is inconsistent with a simple petrologic relation between inclusions and the host melt, and it 

specifically argues against inclusions representing a “restite”-

like material remaining from melting of metabasalt to produce 

dacite. Such restite would be expected to have the same Pbisotope composition as melts unless significant mixing with 

additional melt or crustal assimilation had occurred.

Lead-isotope composition of groundmass material shows 

that disaggregation of inclusions also effects the composition of 

the groundmass on relatively small spatial scales. This phenomenon is particularly evident in SH304-2C, where measurements of the groundmass composition made adjacent to the 

inclusion lie on an apparent mixing line between the compositions of phenocrysts and plagioclase within gabbroic inclusions 

(fig. 5). Contributions from disaggregating inclusions also 

may influence variations in measured (230Th)/(232Th) in plagioclase separates (Cooper and Donnelly, this volume, chap. 36). 

Such contributions from disaggregation of gabbroic inclusions 

probably limit the utility of conventional bulk rock or mineral 

separate-based isotopic measurements for estimating magmatic composition, as these will invariably represent a mixture 

between the true isotopic composition of the magma and various admixtures of disaggregated and remelted inclusions.

Conclusions

Laser ablation ICP–MS analyses of anorthite content, 

trace-element (Li, Ti, Sr, Ba, LREE, Pb) concentrations, and Pbisotope compositions in plagioclase from dacite of the 2004–5 

and 1981–86 eruptions of Mount St. Helens provide insight 

into the petrologic processes leading to formation and eruption 

of these magmas. Anomalously high Li contents in the early 

stage of the eruption are thought to reflect addition of Li to the 

upper part of the magma chamber immediately before eruption 

(within ~1 year) by transfer of an alkali-enriched exsolved vapor 

from deep within the mama chamber. Accumulation of Li-rich 

vapor in the apical part of the magma storage zone may have 

increased fluid pressures, perhaps helping to initiate eruption. 

The compositional ranges of other trace elements in plagioclase 

phenocrysts remain largely constant in material erupted between 

October 2004 and April 2005 and are broadly similar to those 

measured in 1981–85 dome samples. These elements show 

significant correlations with anorthite content that, particularly for Sr and Ti, cannot be described solely by variations 

in plagioclase–melt partitioning. A simple model involving 

closed-system fractional crystallization of plagioclase + hypersthene + amphibole + oxides largely reproduces the observed 

trends, suggesting that plagioclase compositions are predominantly controlled by the degree of crystallinity of the magma 

and sequestration of compatible elements during crystallization. Analyses from traverses within individual plagioclase 

phenocrysts generally support this model but also suggest that 

localized variations exist in the composition of melt from which 

individual plagioclase crystallize. These analytical differences 

probably reflect localized variations in phase proportions and 

crystallinity during magma residence.

Plagioclase from gabbronorite inclusions in three samples 

has markedly different trace-element and Pb-isotope compositions compared to plagioclase phenocrysts. Inclusions 

typically have higher Ti, Ba, REE, and Pb and lower Sr and 

have lower 208Pb/206Pb and 207Pb/206Pb ratios than coexisting 

plagioclase phenocrysts. The compositions of plagioclase 

from gabbroic inclusions appear to be unrelated to phenocryst 

compositions. We suggest that gabbroic inclusions are samples 

of the mafic Tertiary basement from beneath the volcano.

Acknowledgments

The facilities of the W.M. Keck Collaboratory for Plasma 

Mass Spectrometry were used in this work. We would also 

like to thank the staff of the U.S. Geological Survey David 

A. Johnston Cascades Volcano Observatory for providing a 

unique sample set, and the Mount St Helens petrology working group (you know who you are!) for feedback and comments. This work was funded in part by National Science 

Foundation grant EAR 0440382.

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35. Trace Element and Pb Isotope Composition of Plagioclase from Dome Samples from the 2004–2005 Eruption 825

Appendix 1. Supplementary Analytical 

Data for Glass Standard BCR-2G 

and for Pb-Isotope Compositions of 

Plagioclase and Groundmass in 2004–

2005 Mount St. Helens Dome Samples

Data for seven replicate analyses of glass standard BCR2G were determined during this study, and a comparison 

with accepted concentrations is listed in table 4. Table 5 lists 

208Pb/206Pb and 207Pb/206Pb ratios of plagioclase phenocrysts, 

groundmass, and plagioclase within gabbroic inclusions from 

three samples of the Mount St. Helens dome erupted in 2004–

2005 measured by laser-ablation multicollector ICP–MS. The 

average of four replicate isotopic analyses from glass standard 

BCR-2G is included in table 5 for reference.

BCR-2G ---------- Accepted

(μg/g)

Measured

(μg/g) ± 2σ

Li 10.5 11.2 3.0

Ti 13,500 12,047 490

Sr 346 339 13

Ba 674 624 25

La 25.3 23.1 0.9

Ce 53.7 50.8 2.5

Pr 6.9 6.3 0.5

Nd 28.8 26.2 1.8

Eu 1.9 1.8 0.2

Pb 11.0 11.8 0.7

CaO/SiO2 0.13 0.12 0.02

208Pb/206Pb 207Pb/206Pb

SH304-2A

Phenocryst 2.0495 (62) 0.8309 (29)

Phenocryst 2.0507 (48) 0.8311 (22)

Phenocryst 2.0443 (35) 0.8275 (16)

Phenocryst 2.0443 (27) 0.8260 (18)

Phenocryst 2.0466 (22) 0.8255 (14)

Groundmass 2.0478 (11) 0.8283 (6)

Groundmass 2.0489 (9) 0.8285 (5)

Groundmass 2.0474 (11) 0.8276 (5)

Groundmass 2.0466 (5) 0.8277 (5)

Groundmass 2.0477 (8) 0.8271 (7)

SH304-2C

Phenocryst 2.0459 (10) 0.8277 (4)

Phenocryst 2.0505 (38) 0.8310 (12)

Groundmass 2.0451 (15) 0.8275 (9)

Groundmass 2.0465 (11) 0.8278 (8)

Groundmass 2.0451 (9) 0.8275 (1)

Groundmass 2.0465 (8) 0.8278 (1)

Groundmass 2.0429 (12) 0.8234 (1)

Groundmass 2.0420 (4) 0.8245 (1)

Groundmass 2.0398 (12) 0.8234 (1)

Inclusion 2.0407 (9) 0.8209 (4)

Inclusion 2.0397 (11) 0.8206 (4)

Inclusion 2.0403 (17) 0.8216 (7)

Inclusion 2.0397 (14) 0.8209 (6)

SH306-1

Phenocryst 2.0478 (33) 0.8288 (19)

Phenocryst 2.0457 (49) 0.8293 (33)

Groundmass 2.0468 (7) 0.8269 (6)

Groundmass 2.0470 (14) 0.8272 (5)

Groundmass 2.0472 (15) 0.8287 (9)

Inclusion 2.0420 (7) 0.8204 (3)

Inclusion 2.0369 (15) 0.8193 (5)

Inclusion 2.0362 (9) 0.8192 (6)

Inclusion 2.0396 (12) 0.8198 (7)

Inclusion 2.0378 (23) 0.8213 (11)

Inclusion 2.0369 (17) 0.8208 (9)

Inclusion 2.0397 (17) 0.8234 (9)

BCR-2G (n=4) 2.0637 (10) 0.8316 (5)

Table 4. Analytical data from replicate analysis of USGS 

glass standard BCR-2G.

[Average of seven analyses. CaO/SiO2 ratios calculated from equation 1 (see text). “Accepted values” compiled from these sources: 

rare-earth elements from Kent and others (2004b); Li from Kent and 

Ungerer (2006); Si from Rocholl (1998); all other elements from Kelley and others (2003).]

Table 5. Pb-isotope compositions of plagioclase and 

groundmass in 2004–2005 Mount St. Helens dome samples.

[Analysis by laser-ablation multicollector ICP–MS at Oregon State University. 

Values are isotopic ratios and, parenthetically, their 1σ uncertainties expressed 

in the final decimal places. “Inclusions” refers to plagioclase in gabbroic 

inclusions. Average of four replicate isotopic analyses from glass standard 

BCR-2G included for reference.]
826 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 2. Anorthite Content and Trace-Element Abundances in Five 

Plagioclase Grains from Dome Sample SH304-2A

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web: at http://pubs.usgs.gov/pp/1750. ]

Appendix 2 is a spreadsheet that lists the mole fraction anorthite and measured concentrations for six elements (Li, Ti, Sr, 

Ba, La, Ce) from one traverse across each of five plagioclase grains. Traverses range in length from 375 to 775 μm.

Appendix 3. Anorthite and Trace Element Abundances in Plagioclase from 

2004–2005 and 1981–1985 Dome Samples from Mount St. Helens

[This appendix appears only in the digital versions of this work—in the DVD-ROM that accompanies the printed volume 

and as a separate file accompanying this chapter on the Web at http://pubs.usgs.gov/pp/1750. ]

Appendix 3 is a spreadsheet that lists the mole fraction anorthite and analytical data for 10 elements (Li, Ti, Sr, Ba, La, Ce, 

Pr, Nd, Eu, Pb) from 203 analyses of plagioclase phenocryst cores, margins, and inclusions in eight sample erupted in 

2004–2005. Also shown are 28 analyses of plagioclase phenocryst cores in three samples erupted during the period 1981–85.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 36

1

Department of Earth and Space Sciences, University of Washington, Box 

351310, Seattle, WA 98195; now at Geology Department, University of California, One Shields Avenue, Davis, CA 95616

2

 Department of Earth and Space Sciences, University of Washington, Box 

351310, Seattle, WA 98195

238U-230Th-226Ra Disequilibria in Dacite and Plagioclase 

from the 2004–2005 Eruption of Mount St. Helens

By Kari M. Cooper1 and Carrie T. Donnelly2

Abstract

Uranium-series disequilibria in whole-rock samples and 

mineral separates provide unique insights into the time scales 

and processes of magma mixing, storage, and crystallization. 

We present 238U230Th-226Ra data for whole-rock dacite and 

gouge samples and for plagioclase separated from two dacite 

samples, all erupted from Mount St. Helens between October 

2004 and April 2005. We also present new 238U-230Th disequilibria for a suite of four reference samples from the 1980–86 

eruption of Mount St. Helens. We use the U-series data to 

evaluate the origin of the 2004–5 magma, its relation to the 

1980–86 magma, and the relation of 2004–5 phenocrysts to 

their host magmas. Dacite samples from 2004–5 show variable 

(

230Th)/(238U), ranging from 238U-enriched to 230Th-enriched. 

(230Th)/(232Th) ratios in 2004–5 dacite and gouge samples do 

not vary outside of analytical error and are within the range of 

(230Th)/(232Th) measured for the 1980s reference suite. However, (230Th)/(232Th) for plagioclase separates for dome samples 

erupted during October and November 2004 are significantly 

different from corresponding whole-rock values, which suggests that a large fraction (>30 percent) of crystals in each 

sample are foreign to the host liquid. Furthermore, plagioclase 

in the two 2004 samples have U-series characteristics distinct 

from each other and from plagioclase in dacite erupted in 

1982, indicating that (1) the current eruption must include a 

component of crystals (and potentially associated magma) 

that were not sampled by the 1980–86 eruption, and (2) dacite 

magmas erupted only a month apart in 2004 contain different 

populations of crystals, indicating that this foreign component 

is highly heterogeneous within the 2004–5 magma reservoir.

Introduction

After an 18-year hiatus, Mount St. Helens erupted in 

October 2004, beginning a period of dacite dome extrusion 

that produced more than 70 million cubic meters of crystalrich, gas-poor dacite during the first 18 months of eruption 

(Pallister and others, this volume, chap. 30). Extrusion continues at the time of this writing (2007). The dome has erupted 

as essentially a solid plug, extruding a series of spines mantled 

by meter-thick fault gouge (for example, Pallister and others, 

this volume, chap. 30; Cashman and others, this volume, chap. 

19). Samples have been collected by the staff at the U.S. Geological Survey Cascades Volcano Observatory and distributed 

to a group of ~25 investigators working on various petrologic 

and geochemical investigations aimed at understanding the 

origin and characteristics of the newly erupted material and 

the relation of the currently erupting magma to that erupted in 

the 1980–86 eruptive cycle. Major- and trace-element compositions of the 2004–5 dacite whole-rock samples are similar to 

those of the 1980–86 dacite dome (for example, Pallister and 

others, 2005; and this volume, chap. 30), suggesting the possibility that the current eruption is tapping magma that remained 

in the reservoir after the 1980s. In this chapter, we present 

238U-230Th-226Ra data for a suite of samples from the first six 

months (October 2004–April 2005) of the ongoing eruption of 

Mount St. Helens, and 238U-230Th data for a suite of reference 

samples from the 1980–86 eruption. We focus here on wholerock data for dacite dome and gouge samples for the 2004–5 

eruption; we also present data for plagioclase separated from 

two dacite samples. These data complement other analyses by 

the Mount St. Helens petrology working group presented in 

this volume. We use these data to address the following questions: (1) Is the current eruption fed by a new batch of magma 

previously unseen at Mount St. Helens, and/or is this eruption 

tapping a remnant magma body from the 1980–86 eruption? 

(2) What is the relation of plagioclase phenocrysts to the host 

magmas in the 2004 dome samples?
828 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Background: U-Series Disequilibria in 

Older Mount St. Helens Lavas

U-series Disequilibria in Volcanic Rocks

U-series disequilibria in volcanic rocks have been used 

for decades as a tool for understanding magmatic processes 

(see reviews by Bourdon and others, 2003; Ivanovich 

and Harmon, 1992, and references therein). Disequilibria 

measured in whole-rock, glass, and mineral separates can 

be used both as tracers of chemically distinct magmas or 

crystals (for example, Cooper and Reid, 2003; Turner and 

others, 2003c) and to provide temporal and chemical information about magmatic processes ranging from melt generation (Bourdon and Sims, 2003; Lundstrom, 2003; Sims and 

others, 2003; Sims and others, 2002; Turner and others, 

2003a) to magma residence and differentiation within the 

crust (Blake and Rogers, 2005; Condomines and others, 

2003; Cooper and others, 2001; Hawkesworth and others, 

2004; Reid, 2003; Reid and others, 1997; Rogers and others, 

2004; Turner and others, 2003b; Turner and others, 2003c; 

Vazquez and Reid, 2004) to degassing and related crystal 

growth (Berlo and others, 2004; Reagan and others, this 

volume, chap. 37; Turner and others, 2004). A unique aspect 

of U-series studies is that they allow connection between 

chemical variations within crystals and absolute ages (for 

example, Vazquez and Reid, 2004).

U-series studies of magmatic processes are based on the 

fact that 238U decays to 206Pb through a series of intermediate daughter isotopes with half-lives ranging from less than 

one second to as much as 0.245 m.y. (234U; Cheng and others, 

2000). Any U-bearing system, given sufficient time as a closed 

system, will attain a state of radioactive or secular equilibrium 

where the rate of disintegration of each of the nuclides in the 

chain is the same. Magmatic processes, such as melt generation, crystallization, and degassing, can fractionate the different nuclides in the chain, creating disequilibria between individual parent-daughter pairs. Such disequilibria will return to 

radioactive equilibrium at a rate dictated by the half-life of the 

daughter isotope; thus the degree of disequilibrium preserved 

in erupted lavas contains information about the time since the 

last fractionation event, and each parent-daughter pair is sensitive to a different time scale. In particular, 230Th-238U disequilibria can record events that occurred within the past ~10,000 

years to ~350,000 years, whereas 230Th-226Ra disequilibria 

are sensitive to time scales of a few hundred years to ~10,000 

years. This range of hundreds to hundreds of thousands of 

years is commensurate with time scales of magma differentiation and storage in many systems (for example, recent reviews 

by Condomines and others, 2003; Hawkesworth and others, 

2004; Reid, 2003; Zellmer and others, 2005), including that of 

older Mount St. Helens lavas (Cooper and Reid, 2003; Volpe 

and Hammond, 1991).

U-series Disequilibria in Older Mount St. Helens 

Lavas

U-series data in older Mount St. Helens lavas have been 

discussed previously (Bennett and others, 1982; Cooper and 

Reid, 2003; Volpe and Hammond, 1991), which provides an 

excellent context for our analyses of the 2004–5 samples. 

Volpe and Hammond (1991) analyzed 238U-230Th-226Ra disequilibria in whole-rock samples, groundmass, and mineral separates for seven lavas erupted from Mount St. Helens within 

the past ~2,000 years (figs. 1A, 2A). Data for whole-rock and 

groundmass samples span an analytically significant range of 

230Th/232Th ratios (fig. 1A), indicating that the different magmas had different source regions (Volpe and Hammond, 1991). 

These samples are also significantly enriched in 226Ra compared to radioactive equilibrium with 230Th (that is, (226Ra)/

(

230Th) > 1, where by convention parentheses around the chemical symbols represent activities) with (226Ra)/(230Th) ranging 

from 1.07 in an andesite of Castle Creek age (MSH 90-4) to 

1.55 in a sample of dacite erupted in 1982 (MSH 90-9; fig. 

2A). The preservation of these 226Ra excesses requires that the 

time between the last fractionation event (likely to be melt 

generation) and eruption was less than about 10,000 years.

Volpe and Hammond’s (1991) 238U-230Th data for mineral separates and whole-rock samples define linear arrays on 

a 238U-230Th isochron diagram, with slopes corresponding to 

apparent ages of 2–6 ka for most samples (fig. 1A). In contrast, 

two samples of Castle Creek age (erupted ~2 ka) yield apparent ages of 34–27 ka (Volpe and Hammond, 1991). At the same 

time, all of the mineral separates preserve 226Ra-230Th disequilibria indicating fractionation of Ra from Th within the past 

10,000 years (fig. 2A), which is in conflict with the 230Th-238U 

apparent ages for Castle Creek samples. In detail, the 226Ra-230Th 

data for minerals and whole-rock analyses of the same samples 

do not define linear arrays on a Ba-normalized 226Ra-230Th 

isochron diagram (fig. 2A), which Volpe and Hammond (1991) 

interpreted as evidence of addition of a 226Ra-enriched fluid to 

the liquid after crystallization of the minerals. Cooper and Reid 

(2003) reexamined Volpe and Hammond’s data by using a new 

method of calculating 226Ra-230Th ages that accounts for impurities in the mineral separates and for differences in partitioning 

behavior of Ra and Ba. After applying this method, all of the 

Mount St. Helens data can be explained by crystallization and 

aging alone, without requiring late-stage addition of Ra to the 

system. All of the revised 226Ra-230Th ages are consistent with 

recent (less than a few thousand years) crystal growth, and, with 

the exception of the two Castle Creek samples, are concordant 

with 230Th-238U ages. 226Ra-230Th and 230Th-238U disequilibria 

return to secular equilibrium on different time scales; thus average ages for bulk separates with a protracted growth history 

will weight the old and young components differently for the 

two parent-daughter pairs, resulting in different average ages. 

Cooper and Reid (2003) interpreted the discordant 226Ra-230Th 

and 230Th-238U ages in the Castle Creek samples to reflect multiple phases of crystallization within the bulk mineral separate, 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 829

1.00

1.10

1.20

1.30

1.40

(230Th)/(232Th) 

1.00

1.10

1.20

1.30

1.40

0.4 0.6 0.8 1.0 1.2 1.4 1.6

(

238U)/(232Th)

(230Th)/(232Th)

1.00

1.10

1.20

1.30

(230Th)/(232Th)

B

A

C

MSH90-9 (1982)

1980–86 reference suite

MSH90-5 (Castle Creek)

MSH90-7 (Sugar Bowl)

MSH90-3 (Kalama)

MSH90-8 (Kalama)

MSH90-2 (Castle Creek)

gm wr px pl ol

MSH90-9 (1982)

SH305-1 (Nov. 2004)

SH304-2A8 (Oct. 2004)

SH304-2A1 (Oct. 2004)

Pre-1980 whole-rock

MSH90-9 (1982)

1980–86 reference suite

SH305-1 (Nov. 2004)

SH304-2A8 (Oct. 2004)

SH304-2A1 (Oct. 2004)

SH307 and 320 (gouge)

(Jan.–Apr. 2005)

SH306 to 317; DRS_3_9_4

Figure 1. 238U-230Th isochron diagrams for Mount 

St. Helens samples. Solid bold line with slope 1 on 

all panels is the equiline (representing radioactive 

equilibrium). Thin solid and dashed lines are fits to 

the mineral-separate data for each sample. A, Pre2004 samples. Black squares are analyses of 1980–

86 samples from this study; error bars are shown 

for ±1-percent uncertainties. “MSH90-X” data are 

from Volpe and Hammond (1991); error bars for 

those samples are omitted for clarity but are similar 

in size to those shown for the 1980–86 referencesuite samples. Symbols for different phases are 

as follows: diamonds, groundmass (gm); squares, 

whole-rock (wr); circles, pyroxene (px); triangles, 

plagioclase (pl); cross, olivine (ol). B, Data for 

1980–2005 samples. Shown for reference are pre1980 samples (whole-rock data only; open squares) 

and 1982 dacite (MSH90-9; gray symbols; from Volpe 

and Hammond, 1991). Error bars for 2004–5 samples 

are shown only where larger than plotted symbol. 

Symbols for plagioclase and whole-rock data as in 

panel A; gouge samples indicated by “x” symbols. 

C, Whole-rock data and plagioclase separates for 

October 2004 dome (SH304-2A) and November 2004 

dome (SH305-1). The 1982 dacite sample is shown 

for reference (Volpe and Hammond, 1991). Symbols 

as in panels A and B.

either crystal populations of differing ages or old cores with 

younger overgrowths. In addition, plagioclase in the 1982 dacite 

(MSH90-9) has anomalously high Ra/Ba compared to what 

would be expected for equilibrium partitioning, which may 

reflect crystallization that was rapid enough to prevent chemical equilibrium between liquid and crystals (Cooper and Reid, 

2003). These observations would predict that if the 2004–5 

magmas were simply remnant magma from the 1980–86 

eruption, then (1) whole-rock isotopic compositions (including (230Th)/(232Th)) would be identical in 2004–5 and 1980–86 

samples, and (2) crystals within the 2004–5 dacite would have 

238U-230Th-226Ra characteristics like those in crystals erupted in 

the 1980s. Our new data for 2004–5 samples provide a test of 

these predictions.

Analytical Methods

Chemical separation of U-series elements and isotopic 

analysis by multicollector inductively coupled plasma mass 

spectrometry (MC-ICP-MS) were conducted at the University 

of Washington. Mass-spectrometry protocols and 

results for rock and solution standards are detailed 

in appendix 1. Chemical separation procedures 

were modified from published procedures (Goldstein and Stirling, 2003; Pietruszka and others, 

2002) and are also detailed in the appendix. Briefly, 

the elements of interest were separated from the 

rock or mineral matrix using ion-exchange chromatography, conducted in class-100 laboratory facilities (low-particulate facilities, designed to minimize 

trace-metal contamination). Rock chips or mineral 

separates (0.5–1.5 g) were dissolved in concentrated hydrofluoric+nitric acids, followed by evaporation and dissolution in hydrochloric acid. Resulting sample solutions were split into three aliquots: 
830 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

1,000

500

0

1,500

2,000

2,500

3,000

1,000

500

0

1,500

2,000

2,500

3,000

1,000

500

0

1,500

2,000

2,500

3,000

0 500 1,000 1,500 2,000 2,500 3,000

(

230Th)/[Ba], IN DISINTEGRATIONS PER

YEAR PER MICROGRAM (dpy/μg)

(226Ra)/[Ba], IN DISINTEGRATIONS PER YEAR PER MICROGRAM (dpy/μg)

B

C

A

Pre-1980 samples

MSH90-9 (1982)

SH305-1 (Nov. 2004)

SH304-2A1 (Oct. 2004)

SH307 and 320 (gouge)

MSH90-9 (1982)

MSH90-5 (Castle Creek)

MSH90-7 (Sugar Bowl)

MSH90-3 (Kalama)

MSH90-8 (Kalama)

MSH90-2 (Castle Creek)

gm wr px pl ol

MSH90-9 (1982)

SH305-1 (Nov. 2004)

SH304-2A1 (Oct. 2004)

(Jan.–Apr. 2005)

SH306 to 317; DRS_3_9_4

Figure 2. 226Ra-230Th-Ba isochron 

diagrams for Mount St. Helens samples. 

Solid line with slope 1 (the equiline, 

representing radioactive equilibrium) 

is shown on all panels. A, Pre-2004 

samples. Symbols as in figure 1 and 

shown in the legend. B, Data for 1980–

2005 samples. Pre-1980 samples (wholerock and plagioclase data only; open 

symbols) are shown for reference. Error 

bars are shown where they are larger 

than the size of the symbols. C, Wholerock data and plagioclase separates 

for October 2004 dome (SH304-2A) and 

November 2004 dome (SH305-1). The 

1982 dacite sample (Volpe and Hammond, 

1991) is shown for reference (gray 

symbols). Symbols as in panels A and B.

one small aliquot was spiked with 233U and 229Th for U and Th 

concentration measurement by isotope dilution, a second small 

aliquot was spiked with 135Ba for Ba concentration measurement 

by isotope dilution, and the largest aliquot was spiked with 228Ra 

and used for measurement of Ra concentration by isotope dilution and U and Th isotopic compositions. Uranium and thorium 

(both for measurement of isotopic composition and for isotope 

dilution) were separated from the rock matrix and from each 

other using Eichrom TRUTM resin. Radium was subsequently 

separated from the rock matrix washed from the TRU column 

of the Th and U isotopic-composition aliquot using cationexchange resin and was purified using Eichrom Sr-specTMresin 

and a final cation-exchange microcolumn. Isotopic ratios were 

measured using a Nu Plasma MC-ICP-MS equipped with three 

ion counters and an energy filter on the high-mass ion counter 

for increased abundance sensitivity.

Results

238U-230Th Disequilibria

2004–2005 Dome Samples

230Th-238U disequilibria for 2004–5 

Mount St. Helens dome dacites are 

within the range of those measured in 

older Mount St. Helens lavas (Volpe 

and Hammond, 1991; table 1, fig. 1B). 

(

230Th)/(232Th) ratios measured in 2004–5 

whole-rock dacite samples span a range 

of ~2 percent (from 1.257 to 1.282), 

analytically indistinguishable from 

each other given our reproducibility of 

±1 percent. The dome samples show 

some variation in U/Th ratio, ((238U)/

(

232Th) = 1.159–1.366), resulting in 

a variation in 230Th-238U disequilibria 

from 238U-enriched ((230Th)/(238U) = 

0.929) to 230Th-enriched ((230Th)/(238U) = 

1.103). The majority of arc lavas are 238U 

enriched, which is generally attributed 

to the influence of subduction-zone 

fluids, but 230Th-enriched lavas have been 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 831

observed in some arc settings (Turner and others, 2003a). It is 

also possible, given the presence of zircon in a sample of 2004–6 

dome material (C. Miller, written commun., 2006), that some 

of the variation in U/Th ratios is due to differences in percentage of zircon and/or incomplete dissolution of zircon in different samples. No systematic decrease in U/Th ratio or 230Th-238U 

disequilibria over time is observed (not shown), although the 

two earliest samples (October–November 2004 lava spines) 

have the highest U/Th ratios and are the only samples that are 

238U-enriched.

1980–1986 Reference Suite

The four reference samples from the 1980–86 eruption 

have thorium isotopic compositions that overlap 230Th/232Th 

measured in the 2004–5 samples (fig. 1B). Three samples 

(SH32, SH131, and SH157, erupted in 1980, 1982, and 1984, 

respectively) have 230Th/232Th ratios at the high end of those 

measured in 2004–5 dome samples, though they overlap 

within measurement error. The 1986 sample has (230Th)/(232Th) 

~4 percent lower than the other three 1980s samples, but it 

is within error of the lowest (230Th)/(232Th) ratio measured 

in 2004–5 dome samples. Our measurement of 230Th-238U 

disequilibria in SH131 (1982 dome) is not within error of the 

measurement of the 1982 dome sample (MSH90-9) reported 

by Volpe and Hammond (1991). This discrepancy could be 

due to thorium isotopic heterogeneity within the 1982 dome, 

or it could indicate that one of the measurements is in error. 

Thorium and U concentrations and Th/U ratios in the two 

samples of the 1982 dome are similar but not identical, which 

may support the former interpretation. For the purposes of this 

paper, we will focus on comparisons of 2004–5 samples with 

our measurements of the 1980s reference suite because they 

were performed in the same lab using the same methods.

Plagioclase Separates for 2004 Dome Samples

Plagioclase separates from two subsamples of dome rock 

erupted in October 2004 (SH304-2A1 and SH304-2A8) have 

U/Th ratios significantly lower than the whole-rock values 

for the same samples, qualitatively consistent with relative 

partitioning of U and Th in plagioclase (Blundy and Wood, 

2003). 230Th/232Th ratios in the plagioclase separates are also 

low compared to the corresponding whole-rock values (table 

1; fig. 1), and both plagioclase separates have large 230Thenrichments ((230Th)/(238U) = 2.087 and 1.404, respectively). A 

plagioclase separate from the dome erupted in November 2004 

(SH305-1) is also 230Th enriched ((230Th)/(238U) = 1.117), with 

U/Th ratio lower than the corresponding whole-rock measurement but higher than that in plagioclase from SH304-2A. 

The 232Th/230Th ratios in plagioclase and whole-rock samples 

of SH305-1 differ by about 4 percent, outside of analytical 

uncertainty. The U/Th ratios measured in plagioclase separates 

for the October 2004 dome (SH304-2A1 and SH304-2A8) are 

significantly lower than U/Th ratios measured by Volpe and 

Hammond (1991) in plagioclase separated from older Mount 

St. Helens samples, whereas plagioclase in the November 

2004 dome sample (SH305-1) has U/Th ratio similar to those 

measured in plagioclase by Volpe and Hammond (1991). 

These differences in U/Th most likely reflect differences in 

the percentage or nature of impurities (groundmass and/or 

inclusions of glass or other mineral phases) in the plagioclase 

separates, although some part of the variation may reflect differences in U/Th of the magmas from which they crystallized.

2004–2005 Gouge Samples

The (230Th)/(232Th) ratios measured in gouge samples 

(SH307-2A and SH320-1) are within error of those measured 

in the 2004–5 dacites, although (238U)/(232Th) ratios are significantly lower than those measured in dome samples. Concentrations of both uranium and thorium are lower in gouge 

samples than in dacite samples (table 1). Sample SH320-1 also 

has lower barium and radium concentrations than do dome 

samples, although SH307-2A has barium and radium concentrations within the range of dome whole-rock samples.

226Ra-230Th Disequilibria

All whole-rock dome and gouge samples are enriched 

in 226Ra relative to equilibrium with 230Th, with 226Ra excesses 

of 40–55 percent for whole-rock samples (similar to those 

measured by Volpe and Hammond (1991) for older whole-rock 

samples) and 60–77 percent for gouge samples (table 1; fig. 

2B). Gouge samples have (230Th)/[Ba] ratios lower than those 

of dacite dome samples, but broadly similar (226Ra)/[Ba] ratios 

(fig. 2). (Brackets around the chemical symbol, by convention, 

indicate concentration.)

A bulk plagioclase separate from the October 2004 dome 

sample (SH304-2A1) has 226Ra-230Th-Ba characteristics similar 

to plagioclase separates from older Mount St. Helens samples, 

but it has lower (226Ra)/[Ba] and (230Th)/[Ba] than Volpe and 

Hammond’s (1991) measurements of plagioclase in the 1982 

dome (MSH90-9; fig. 2). The plagioclase separate for the 

November 2004 dome (SH305-1), in contrast, is within error 

of radioactive equilibrium ((226Ra)/(230Th) = 1.039), with significantly higher (230Th)/[Ba] than any of the other Mount St. 

Helens plagioclase separates.

Discussion

Origin of the 2004–2005 Dacite: Remnant 

Magma from 1980–1986?

Whole-Rock Data

Major- and trace-element composition of the 2004–5 

dacite whole-rock samples are similar to those of the 1980–86 

dacite dome (for example, Pallister and others, 2005; this 

volume, chap. 30), suggesting the possibility that the cur-
Sample No. Eruption date Ba, in ppm Ra, in fg/g Ra %SE Th, in ppm U, in ppm 230Th/ 232Th (10-6) (234U)/ (238U) (238U)/ (232Th) (230Th)/ (232Th) (230Th)/ (238U) (226R

(230

1980–1986 samples SH52 1980 -- -- -- 2.049 0.911 7.010 -- 1.353 1.301 0.962 -

SH131 1982 -- -- -- 1.974 0.885 7.002 -- 1.364 1.300 0.953 -

SH157 1984 -- -- -- 2.390 1.063 6.928 -- 1.353 1.286 0.950 -

SH226 1986 -- -- -- 2.496 1.084 6.752 -- 1.322 1.253 0.948 -

2004–2005 samples SH304-2A1 WR #1 10/18/04 338.6 555.01 0.61 2.685 1.162 6.906 1.002 1.317 1.282 0.974 1.4

SH304-2A1 WR #2 10/18/04 332.2 575.12 2.4 2.598 1.166 6.839 1.002 1.366 1.270 0.930 1.5

SH304-2A1 Pl #1 10/18/04 117.9 30.11 1.6 0.175 0.028 5.529 1.004 0.492 1.026 2.087 1.5

SH304-2A1 Pl #2 10/18/04 116.4 -- -- -- 0.029 5.528 1.001 -- 1.026 -- -

SH304-2A8 WR 10/18/04 327.8 -- -- 2.581 1.134 6.826 1.001 1.337 1.267 0.948 -

SH304-2A8 Pl 10/18/04 118.1 -- -- 0.198 0.052 6.015 1.001 0.795 1.117 1.404 -

SH305-1 WR 11/20/04 330.9 521.21 1.8 2.585 1.131 -- 1.005 -- -- -- -

SH305-1 WR replicate 11/20/04 -- 2.465 1.083 6.770 -- 1.338 1.257 0.939 1.5

SH305-1 Pl 11/20/04 113.4 79.81 1.3 0.531 0.204 7.034 1.005 1.169 1.306 1.117 1.0

SH306 12/15/04 329.6 538.02 0.43 2.464 0.965 6.8543 1.002 1.191 1.272 1.068 1.5

SH307-2A gouge 315.9 514.02 1.4 2.065 0.663 6.836 1.003 0.977 1.269 1.299 1.7

SH307-2A gouge replicate 332.6 -- -- 2.137 0.693 -- -- 0.987 -- -- -

SH311-1B WR 01/16/05 327.9 496.31 2.4 2.504 0.953 6.888 1.007 1.159 1.279 1.104 1.3

MSH05DRS_3_9_4 WR 03/08/05 326.9 515.02 1.1 2.513 1.036 6.776 1.004 1.255 1.258 1.002 1.4SH316-1A WR 04/15/05 331.5 505.72 0.44 2.544 1.061 6.842 1.002 1.269 1.270 1.001 1.4SH317-1A WR 05/01/05 331.5 520.22 0.49 2.504 1.023 6.830 1.000 1.244 1.268 1.019 1.4SH320-1 gouge 283.2 406.02 1.2 1.783 0.625 6.849 1.000 1.066 1.271 1.193 1.6

Rock standards

TML -- 3627 1 0.19 31.780 11.201 5.763 -- 1.073 1.070 0.997 0.9

TML replicate -- 3707 2 1.3 31.119 10.944 5.806 1.0025 1.070 1.078 1.007 0.9

BHVO-1 -- 163.32 1.1 1.214 0.421 5.784 1.0035 1.057 1.074 1.016 1.1BHVO-1 replicate -- -- -- 1.213 0.421 -- -- 1.056 -- -- -

BCR-2 672.8 -- -- 5.879 1.701 4.745 -- 0.880 0.881 1.001 -

1 Ra analyses by static, multi-ion-counting routine, as described in appendix 1.

2 Ra analyses by dynamic routine using only IC0, as described in appendix 1.3 Average of two measurements of the same solution during the same day. Individual measurements differed by 0.9 percent.4 Calculated using Ra concentration measured for SH305-1 WR.5 (234U)/(238U) of rock standards represents the mean of eight measurements of the same solution over four analytical sessions. Relative standard deviation of replicate measurements 

and 0.24 percent for BHVO-1.

Table 1. Concentrations, isotopic compositions, and activity ratios measured in Mount St. Helens samples.[See appendix 1 for analytical details. WR, whole-rock; Pl, plagioclase. For duplicate sample names, #1 or #2 indicate chemical separation and analysis of splits of the same solutiodissolution and analysis of splits of the same rock or mineral separate. Accuracy and reproducibility of measurements was <0.5 percent for Th and U concentrations, <2 percent for for 230Th/232Th, and <0.6 percent for 234U/238U (see appendix 1 for details). Uncertainties in Ra concentrations were limited by in-run errors during mass spectrometry, which are quo

of the mean (relative). Decay constants used in calculations were λ234=2.82629×10-6 (Cheng and others, 2000), λ230=9.158×10-6 (Cheng and others, 2000), λ232=4.933×10-11, λ238=1.55

Dashes indicate no data.]
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 833

rent eruption is tapping magma that remained in the reservoir 

after the 1980s. If the current eruption is tapping magma that 

is either unmodified from the 1980s or has been modified 

only by fractional crystallization, one prediction would be 

that isotopic compositions of the 1980–86 dacite would be 

identical to those in the 2004–5 dacite. In the case that the two 

magmas are related by cooling and fractional crystallization, 

but without significant influence of open-system processes 

such as wall-rock assimilation or mixing with compositionally 

distinct batches of magma, some variation in trace-element 

concentrations of the residual liquid would also be predicted. 

Kent and others (this volume, chap. 35) propose a model in 

which the 1980s and 2004–5 magmas are related by fractional 

crystallization that occurred primarily by chemical isolation 

of grain interiors from the host magma, without volumetrically significant gain or loss of plagioclase crystals, but with a 

small volume of an additional crystal component added to the 

magma through disaggregation of xenoliths. In this scenario, 

because crystals are not lost from or appreciably added to 

the bulk magma, whole-rock trace-element ratios would be 

unmodified during crystallization. In detail, however, a simple 

model of fractional crystallization would predict higher concentrations of incompatible elements (such as Ra, Th, U, and 

Ba) in the 2004–5 samples compared to the 1980–86 samples. 

Although Ra, Th, and U concentrations are indeed higher in 

the 2004–5 samples (table 1; Volpe and Hammond, 1991), Ba 

concentrations are lower, which suggests that a more complex 

explanation is necessary.

Nevertheless, the similarities between magmas erupted 

during the 1980s and the 2004–5 dacite (for example, majorand trace-element data in Pallister and others, this volume, 

chap. 30) suggest that the 2004–5 eruption is tapping at least 

some component of magma that remained in the reservoir 

after the 1980–86 eruption. The U-series data are consistent 

with this hypothesis but require some additional component(s) 

in the plagioclase populations (see below). Thorium isotopic 

compositions that we measured in whole-rock samples from 

1980–86 samples span a small, but analytically significant, 

range (230Th/232Th ratios vary by ~4 percent), overlapping the 

thorium isotopic compositions that we measured in the 2004–5 

samples, with (238U)/(232Th) similar to the October–November 

2004 dome (SH304-2A1 and SH305-1; fig. 1B). Later-erupted 

2004–5 samples have lower (238U)/(232Th), with no change in 

(

230Th)/(232Th), suggesting that the 2004–5 magma has some 

trace-element heterogeneity, but that different magmas erupted 

are all closely related. 226Ra-232Th disequilibria are also similar 

in all 2004–5 and 1982 dacites, consistent with the similarity 

in thorium isotopic compositions. Interestingly, there appears 

to have been a slight decrease in 230Th/232Th ratios during the 

course of the 1980–86 eruption (fig. 1; table 1), suggesting 

that at least two magmas with different origins were involved 

in that eruption (and perhaps more, if the much higher 

230Th/232Th measured in a 1982 dacite sample by Volpe and 

Hammond (1991) also reflects compositional heterogeneity).

Volpe and Hammond’s (1991) measurements of (226Ra)/

[Ba] and (230Th)/[Ba] in their sample of 1982 dacite do not 

overlap with our measurements of 2004–5 samples (fig. 2B). 

Whether there is significant heterogeneity within the 1980–86 

and 2004–5 suites, as suggested above for 230Th-238U disequilibria, cannot be assessed at this time because we do not yet 

have 226Ra data for the 1980–86 reference suite. However, 

the 2004–5 samples fall in a relatively tight cluster with the 

exception of the October 2004 dome (SH304-2A1), which has 

similar (226Ra)/[Ba] but slightly higher (230Th)/[Ba], possibly 

due to more plagioclase fractionation from this sample than 

seen in the other dacites.

Plagioclase Separates

When crystal compositions are considered in addition to 

whole-rock compositions, the 2004–5 magma must contain a 

component not observed in the 1980–86 magmas. The 2004–5 

magmas contain plagioclase with compositions unlike those in 

older Mount St. Helens samples, which would require addition 

of crystals to a putative remnant 1980–86 magma. This is most 

dramatically illustrated by 230Th/232Th measured in plagioclase separated from the October 2004 dome (SH304-2A1), 

which is substantially lower than 230Th/232Th measured in any 

recently erupted samples of Mount St. Helens lavas (figs. 1B, 

C). Whether this lower 230Th/232Th ratio is attributed to an 

old component in the plagioclase separate or to the addition 

of a population of plagioclase unrelated to the host magma, 

the low-230Th/232Th component is not present to any significant degree in the plagioclase separate from the 1982 dacite 

measured by Volpe and Hammond (1991). Plagioclase in the 

November 2004 dome (SH305-1) has 230Th/232Th higher than 

the whole rock, also consistent with a plagioclase component 

that is foreign to the host liquid, although the foreign crystal 

component must be of a significantly different composition than that in SH304-2A1. In contrast to differences in 

their 230Th-238U signatures, the 226Ra-230Th-Ba composition of 

plagioclase in the October 2004 dome (SH304-2A1) is similar 

to plagioclase in samples erupted before 1980, but distinct 

from plagioclase in the 1982 dome sample (MSH90-9; fig. 2). 

In addition, the 226Ra-230Th-Ba signature of plagioclase of the 

November 2004 dome (SH305-1) is different from all measurements of plagioclase in older Mount St. Helens samples, 

including MSH90-9 (Volpe and Hammond, 1991) and SH3

04-2A1 (fig. 2). 

The presence of a foreign plagioclase component in 

the 2004–5 lavas is supported by Pb isotopic heterogeneity 

between plagioclase crystals and groundmass (Kent and others, this volume, chap. 35), by high Al in amphibole crystals 

that is not in equilibrium with the host liquid at pressures 

<300 MPa (Rutherford and Devine, this volume, chap. 31), by 

variable pressure-temperature conditions of crystallization of 

amphibole (Thornber and others, 2005), and by textures and 

zoning in plagioclase in the 2004–5 dacite that are distinct 

from the patterns observed in 1980–86 dacites (Streck and 

others, this volume, chap. 34). Therefore, the 2004–5 eruption 

clearly samples multiple crystal populations and, potentially, 

some associated liquid fractions that were not present in the 
834 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

eruptive products from 1980–86. However, given the similarity 

in whole-rock U-series compositions of the dacites, any liquid 

fraction associated with the crystals must be volumetrically 

minor, suggesting that the crystals were incorporated largely 

as a crystal mush or cumulate.

One potential scenario to account for these observations 

is that the 2004–5 magma is dominated volumetrically by 

remnant magma from 1980–86 but that it incorporated disaggregated crystal mush from within the reservoir system during the 

18-year hiatus between eruptions, consistent with the presence 

of gabbroic inclusions within many of the 2004–5 dacite samples. In this case, the crystal mush must be genetically distinct 

from the 1980–86 magmas in order to explain the difference 

between (230Th)/(232Th) measured in the 2004–5 plagioclase 

separates and that measured in 1980–86 and 2004–5 wholerock samples. It is also possible that the 2004–5 eruption is at 

least partially tapping new magma introduced into the reservoir 

system, perhaps associated with some of the deep earthquake 

swarms between 1986 and 2004 (Moran and others, 2005). Considering the similarity in major- and trace-element compositions 

of the 2004–5 dacite to dacites erupted earlier at Mount St. Helens, this new component may have originated through partial 

melting in the lower crust (see, for example, Pallister and others, 

1992; Smith and Leeman, 1987). The range in (230Th)/(238U) and 

(

226Ra)/(230Th) ratios observed in the 2004–5 samples is within 

the predicted range for incongruent dehydration melting of 

lower crustal amphibolite (Dufek and Cooper, 2005). However, 

we cannot rule out an origin through fractional crystallization of 

a garnet-bearing assemblage from a mantle-derived magma at 

lower crustal depths, as argued by Blundy and others (this volume, chap. 33). Finally, it is possible that our plagioclase separates contain a component of plagioclase from disaggregated 

xenoliths, either incorporated from Tertiary basement rocks or 

from the plutonic roots of the Mount St. Helens reservoir system 

(Kent and others, this volume, chap. 35; Pallister and others, this 

volume, chap. 30). In any case, the preservation of significant 

226Ra excesses in all of the 2004–5 whole-rock samples indicates 

that the time between melt generation and eruption for the bulk 

of the magma was short, less than one or two half-lives of 226Ra 

(that is, less than a few thousand years).

Plagioclase in 2004–2005 Dacite: Effects of 

Impurities and Crystal Ages

Plagioclase in October 2004 Dome Samples

Plagioclase separates for the October 2004 dome subsamples (SH304-2A1 and SH304-2A8) have strikingly different thorium isotopic compositions than any of the 2004–5 

whole-rock samples (fig. 1). Taken at face value, the slope of 

the two-point plagioclase-whole-rock isochron would correspond to an apparent crystallization age of ~42 ka. However, 

this apparent age is inconsistent with the preservation of 226Ra 

excess measured in the same separates, precluding a simple 

history of crystal growth over a short time interval (less than 

a few hundred years) that occurred tens of thousands of years 

ago. Instead, this anomalously old apparent age likely indicates that the plagioclase separates include some percentage of 

older and/or foreign crystals in addition to plagioclase crystallized recently from the host liquid.

We estimate, on the basis of examination with binocular 

microscope, that the bulk plagioclase separate for SH304-2A1 

is >95 percent pure. However, there are a number of components that could be present in small quantities in the bulk 

plagioclase separate, and the data plotted in figure 1 reflect 

the mixture of all of these; thus, in order to discriminate the 

effects of plagioclase aging from the effects of impurities, we 

must consider the U-series signature of each potential contributor. The bulk plagioclase separate could comprise a combination of plagioclase crystallized from the host liquid (at or 

near the time of eruption, and/or during storage in the crustal 

reservoir system); plagioclase that is foreign to the host liquid 

(either as a separate population of plagioclase crystals and/or 

as cores of crystals with multiple growth stages); groundmass 

adhering to the outside of plagioclase grains; and microscopic 

inclusions of melt or other phases within the plagioclase crystals. These in turn could include small amounts of the other 

major crystallizing phases, for example, amphibole, pyroxene, 

and Fe-Ti oxides (Pallister and others, this volume, chap. 30; 

Rutherford and Devine, this volume, chap. 31), and small mass 

fractions of accessory phases: apatite and zircon (C. Miller, 

written commun., 2006). Concentrations of Th, U, and Ra in 

pyroxene or Fe-Ti oxides will be low enough that a few percent of each would not significantly affect the 238U-230Th-226Ra 

systematics of the bulk plagioclase separate (see also Blundy 

and Wood, 2003; Cooper and others, 2001); therefore, we will 

not consider them further. We discuss the U-series signatures 

and effects of each of the other components below.

Zero-Age Plagioclase

Plagioclase textures and rim compositions that are in 

equilibrium with the host liquid indicate that some fraction 

of plagioclase in each of the dome samples grew within a 

short time of eruption (for example, Pallister and others, this 

volume, chap. 30; Streck and others, this volume, chap. 34), 

at least partly as a result of decompression-induced crystallization that occurred during ascent of the 2004–5 magma 

(Blundy and others, this volume, chap. 33; Pallister and others, this volume, chap. 30). In addition, the lack of europium 

anomalies in REE patterns for whole-rock samples indicates 

that gain or loss of plagioclase to the magma must have been 

volumetrically minor (Blundy and others, this volume, chap. 

33; Kent and others, this volume, chap. 35; Pallister and 

others, this volume, chap. 30), far less than the ~40 percent 

by volume of plagioclase crystals present in the magmas. 

In addition, plagioclase separates for 2004–5 samples have 

210Pb/226Ra ratios that are not in radioactive equilibrium, 

indicating that some component of the plagioclase grew 

within decades of eruption (Reagan, 2005; Reagan and 

others, this volume, chap. 37). Therefore, the plagioclase 

separates certainly contain some plagioclase that crystal-
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 835

lized from the host liquid at effectively zero age with respect 

to 238U-230Th-226Ra disequilibria (that is, less than ~10 ka 

for 238U-230Th and less than a few hundred years for 230Th- 226Ra). Such zero-age plagioclase would have 230Th/232Th 

equal to that in the host liquid, and 238U/232Th consistent with 

partitioning of Th and U between plagioclase and liquid 

at low pressure at ~850°C (based on oxide thermometry; 

Pallister and others, this volume, chap. 30; Rutherford and 

Devine, this volume, chap. 31). Blundy and Wood (2003) 

estimate DU/DTh in plagioclase to be ~0.18 at 900°C for 

An40, similar to the temperatures and compositions appropriate for 2004–5 Mount St. Helens lavas. Using these data, we 

calculate (238U)/(232Th) in plagioclase in equilibrium with 

the October 2004 dome sample to be ~0.24 (fig. 3, table 2). 

Absolute values for DU and DTh are less well constrained, but 

by using DU = 6.0±3.7×10-4 (measured in An58 plagioclase 

at 950°C; Blundy and Wood, 2003), we estimate Th and U 

concentrations of roughly 1–2 ppb in pure plagioclase (table 

2). These calculations provide minimum estimates of Th 

and U concentrations in plagioclase in equilibrium with the 

liquid fraction of the magma, as the whole rock contains 

~30–40 percent crystals and Th and U are incompatible in 

all of the major crystallizing phases. However, the Th/U 

ratio will likely be similar in the liquid and whole rock, and 

calculated (238U)/(232Th) is likely more robust to changes in 

mineral mode than are concentrations of U or Th. The bulk 

plagioclase separate, therefore, must contain some phase 

with (238U)/(232Th) higher than ~0.24 and (230Th)/(232Th) 

lower than that measured in the bulk separate (1.026).

Using the elastic-strain model and the latest fitting parameters for plagioclase (Blundy and Wood, 2003), we calculate 

(

226Ra)/[Ba] in plagioclase in equilibrium with a liquid of the 

October 2004 whole-rock composition to be ~330 dpy/μg (An40

at 850°C), and (230Th)/[Ba] of ~1–2 dpy/μg (disintegrations per 

year per microgram). Calculated (226Ra)/[Ba] is slightly higher 

than that measured in the bulk mineral separate, which could 

reflect some aging of the plagioclase since crystallization (as 

much as ~500 years), and/or the effects of other, low-Ra/Ba 

phases in the bulk separate. Th/Ba ratios could be significantly 

different in melt and whole-rock, considering that Ba is several 

orders of magnitude less incompatible than Th in plagioclase 

and amphibole, which dominate the crystal mode. Th/Ba in the 

liquid will therefore be higher than that in the whole-rock, leading to higher (230Th)/[Ba] in plagioclase in equilibrium with the 

liquid. This is consistent with ion microprobe measurements of 

Th and Ba in plagioclase in older Mount St. Helens samples, 

where (230Th)/[Ba] ratios were 10–50 for the 1982 dome 

and ~40–75 for older samples (Cooper and Reid, 2003). We 

estimate that (230Th)/[Ba] in plagioclase crystallizing from the 

October 2004 magma is likely ~10–100, lower than the ratio of 

~200 in the bulk separate.

Other Populations of Plagioclase

In addition to zero-age plagioclase crystallized from the 

host liquid, the bulk plagioclase separate may contain some 

older plagioclase, which would have lower (226Ra)/[Ba] and 

(if older than ~5–10 ka) lower (230Th)/(232Th) than zero-age 

plagioclase. There are three possible origins for older crystals 

in the October 2004 dome plagioclase separate: (1) crystals 

that precipitated from the host magma and subsequently aged; 

(2) plagioclase crystallized from a different magma within 

the Mount St. Helens system (“antecrysts” in the terminology 

of W. Hildreth, as cited in Charlier and others (2005)); or (3) 

crystals incorporated from wall rocks unrelated to the Mount 

St. Helens magmatic system (xenocrysts). Gabbroic to dioritic 

xenoliths have been observed both in the 1980s eruptive 

products (Heliker, 1995) and in the current eruption (Kent and 

others, this volume, chap. 35; Pallister and others, this volume, 

chap. 30). Note that these are termed “inclusions” by Kent and 

others (this volume) and Pallister and others (this volume), but 

we use the term “xenoliths” instead to avoid confusion with 

microscopic inclusions within plagioclase crystals. Pallister 

and others (this volume, chap. 30) estimate the abundance 

of gabbroic to dioritic xenoliths in the 2004–6 dacite at 1–5 

percent. They report ion microprobe U-Pb ages of 25 Ma for 

zircon separated from xenoliths in the 1980–86 lava dome, 

suggesting that the zircons are samples of Tertiary intrusions 

unrelated to the Mount St. Helens magmatic system. Kent 

and others (this volume, chap. 35) interpret the disaggregated 

xenoliths within the 2004–5 dome to have a similar origin. 

However, no xenoliths from the 2004–5 dome have been dated 

and they could, in theory, be of any age as old as the age of the 

basement rocks. In the case of Mount St. Helens, antecrysts 

could have ages as old as ~300 ka (Clynne and others, this 

volume, chap. 28), but older crystals (ranging back to Tertiary 

in age) would be xenocrysts. Regardless of origin, old plagioclase would likely have low (230Th)/(232Th) and low (238U)/

(

232Th), which makes it the only one of the likely components 

in the bulk plagioclase separate that could explain the low 

(

230Th)/(232Th) measured in that separate (fig. 3A).

If the older plagioclase crystallized from the host liquid, 

it would have (238U)/(232Th) like that in zero-age plagioclase 

but with lower (230Th)/(232Th). Assuming that thorium concentrations are the same in old and young populations, we 

calculate that (230Th)/(232Th) of the bulk separate could be 

produced by a mixture of ~70 percent zero-age plagioclase 

with ~30 percent plagioclase in 238U-230Th radioactive equilibrium. The percentage of old crystals in the mixture would be 

higher if they were not in radioactive equilibrium; for example, a mix with 200-ka crystals would require ~40 percent old 

crystals, whereas a mix with ~100-ka crystals would require 

~50 percent old crystals. Conversely, the percentage of old 

crystals may be lower if the thorium concentrations in the old 

crystals are higher, as suggested by the trace-element data for 

plagioclase interpreted by Kent and others (this volume, chap. 

35) as disaggregated gabbroic xenoliths. For example, if the 

old crystals had thorium concentrations twice as high as the 

zero-age crystals, which may be reasonable given the factor of 

2–4 enrichment in incompatible elements, such as Pb and La, 

relative to the zero-age crystals (Kent and others, this volume, chap. 35), the percentage of old crystals required would 
836 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Phase Ba, in 

ppm Ra, in fg/g Th, in ppm

U, in 

ppm

(

238U)/ 

(

232Th)

(

230Th)/ 

(

232Th)1

(

230Th)/ 

(

238U)

(

230Th)/

[Ba], in 

dpy/μg

(

226Ra)/

[Ba], in 

dpy/μg

Zero-age plagioclase 193 56 0.001 0.001 0.24–0.41 1.270 3.10–5.29 1.1 2 332

Zero-age amphibole 34 4.4 0.010 0.005 1.37 1.270 0.929 51 154

Zero-age apatite (L)3 100 no data 4.57 2.12 1.41 1.270 0.899 7.5×103 0 4

Zero-age apatite (M&S)3 149 no data 4.16 3.03 2.22 1.270 0.572 4.5×103 0 4

Zero age zircon 1.3 0.0 43.3 117 8.20 1.270 0.155 5.3×106 0 4

2000 yr apatite (L) 100 373 4.57 2.12 1.41 1.272 0.901 7.5×103 4.3×103

2000 yr apatite (M&S) 149 226 4.16 3.03 2.22 1.287 0.580 4.6×103 2.6×103

2000 yr zircon 1.3 3.5×103 543.3 117 8.20 1.395 0.170 5.8×106 3.1×106

1

 (230Th)/(232Th) for zero-age phases assumed to be equal to that in whole rock.

2

 (230Th)/Ba for zero-age plagioclase calculated from whole-rock composition; actual value may be somewhat higher in plagioclase in equilibrium with liquid 

(see text for discussion).

3

 Calculated using partition coefficients of Luhr and others, 1984 (L), and Mahood and Stimac, 1990 (M&S), respectively.

4

 (226Ra)/Ba assumed to be zero for apatite and zircon at time of crystallization; this assumption makes little difference in the values calculated for older crystals.

5

 Ra concentration in 2,000-yr-old zircon calculated by assuming constant (230Th), which gives a minimum value; because of the high U/Th ratio in zircon, 230Th activity increases by ~10 percent over 2,000 years.

Table 2. Calculated compositions of mineral phases in chemical equilibrium with Mount St. Helens sample SH304-2A1.

decrease by a factor of two. Kent and others (this volume, 

chap. 35) interpreted the lack of Eu anomalies in REE patterns 

in the whole-rock data for the 2004–5 dacite to indicate that 

only a small percentage of plagioclase can have been added 

to or lost from the magma, which would suggest that any 

older crystals present are near radioactive equilibrium (that 

is, hundreds of thousands of years old) and/or that they have 

high concentrations of thorium and presumably other trace 

elements. However, given the uncertainties in the thorium 

concentrations of different potential populations of plagioclase, coupled with the possibility that other phases are present 

in the bulk separates, we cannot uniquely determine whether 

the old crystals are older than ~300 ka and therefore must be 

xenocrystic.

Any old-plagioclase component, regardless of origin, 

with (230Th)/(232Th) lower than the bulk plagioclase separate 

is likely to be older than ~10 ka and therefore the 230Th-226Ra 

parent-daughter pair will be in radioactive equilibrium. If we 

assume that plagioclase in the bulk separate is a mixture of (1) 

plagioclase rims or crystals that grew immediately before or 

during eruption in chemical equilibrium with the host liquid 

and (2) an old plagioclase component (cores or crystals) that 

is >10 ka and therefore in 226Ra-230Th equilibrium, and if we 

further assume that both components have the same thorium 

and barium concentrations and (230Th)/[Ba] ratios, we calculate 

that the bulk separate contains 25–40 percent old plagioclase 

(where the range encompasses variations in crystallization 

temperature from 850°C to 900°C and plagioclase from An50

to An40). If the old component is <10 ka, or if the young component has (226Ra)/[Ba] higher than predicted by equilibrium 

(a potential consequence of rapid crystallization (Cooper and 

Reid, 2003)), the proportion of old crystals would be higher.

However, such calculations are sensitive to assumptions 

about (230Th)/[Ba] and barium concentrations in each population of plagioclase, and the mixing proportions are therefore 

only broadly constrained by this analysis. An additional complication is that radium concentrations in the bulk plagioclase 

separate will be sensitive to the effects of inclusions of other 

phases (for example, zircon, allanite, and groundmass) within 

the separate. Therefore, we do not attempt to interpret the percentage of old plagioclase crystals derived from 230Th-226Ra-Ba 

relations, except to note that the possible range in the percentage of old crystals is similar to the range that we calculated 

using thorium isotopic compositions.

Groundmass or Melt Inclusions

In addition to zero-age plagioclase and older plagioclase, 

some amount of groundmass or melt is likely to be present in 

the plagioclase separate as adherents to the outside of grains 

and/or as melt inclusions within plagioclase. The colinearity of the two plagioclase separates for subsamples of the 

October 2004 dome (SH304-2A1 and SH304-2A8) with the 

whole-rock data (fig. 1), together with higher concentrations 

of thorium, uranium, and barium in sample 2A8 compared to 

2A1, suggests that SH304-2A8 plagioclase contains a higher 

percentage of included groundmass or melt inclusions. The 

difference in concentrations of thorium, barium, and uranium 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 837

Figure 3. A, 238U-230Th; B, 230Th-226Ra isochron diagrams 

illustrating the effects of different types of impurities 

in the bulk plagioclase separates. Orange symbols are 

calculated values for zero-age apatite, zircon, plagioclase, 

and amphibole, in chemical equilibrium with October 2004 

whole-rock (SH304-2A1) and plagioclase in radioactive 

equilibrium; for calculated values see table 2. The 

explanation applies to both panels; multiple symbols are 

shown for zero-age plagioclase and apatite where a 

range of partition coefficients are considered. The inset 

in panel A shows detail of upper-left corner of the larger 

diagram. Mixing trajectories on both diagrams will be 

straight lines; arrows shown in panel B illustrate mixing 

lines between zero-age plagioclase (orange triangle) 

and either zero-age or 2,000-year-old apatite and zircon. 

Length of arrows schematically indicates that (because 

of its higher concentrations of Th and U) zircon will have 

more leverage on the composition of the bulk plagioclase 

separate than will apatite.

between the two subsamples is consistent with the presence 

of 1–2 percent more groundmass (with the composition of the 

whole rock) in SH304-2A8 plagioclase than in the separate 

from SH304-2A1, and it is possible that SH304-2A8 plagioclase itself contains a few percent by mass of groundmass.

The presence of groundmass will not affect the slope of 

the two-point plagioclase-whole rock 238U-230Th isochron and 

therefore will not change the apparent age of the plagioclase 

separate (~42 ka). As demonstrated by a comparison of the 

two plagioclase separates for the October dome samples, the 

presence of groundmass or melt inclusions will shift the bulk 

separate toward higher 230Th/232Th and 238U/232Th, and therefore 

the pure plagioclase will have 230Th/232Th and 238U/232Th even 

lower than those measured in the bulk separate (fig. 3). The 

presence of groundmass would also lead to higher (226Ra)/[Ba] 

and (230Th)/[Ba] in the bulk plagioclase separate compared to 

that in pure, zero-age plagioclase. This presence of groundmass is qualitatively consistent with the observations, but a 

mixing line on a 230Th-226Ra-Ba isochron diagram between 

zero-age plagioclase and the whole rock does not pass through 

the bulk separate (fig. 3) indicating that some other phase(s) 

must be present in the bulk separate.

 Amphibole

Thorium and uranium are moderately incompatible in 

amphibole (LaTourrette and others, 1995; Tiepolo and 

others, 2000), and although the absolute value of the 

partition coefficients vary with SiO2, the partitioning 

behavior of U is similar to that of Th in all compositions. Therefore, amphibole does not appreciably 

fractionate U from Th (Blundy and Wood, 2003, and 

references therein), and the effect of amphibole on the 

238U-230Th systematics of the plagioclase separate will 

be similar to, but of a much smaller magnitude than, 

that of groundmass (fig. 3A).

Radium and barium are incompatible in amphibole, and DRa /DBa in amphibole is ~0.08 regardless 

of crystal composition (Blundy and Wood, 2003). 

Absolute values of DBa range from 0.10 to 0.72, similar 

to the range of DBa in plagioclase. Therefore, at the 

time of crystallization, amphibole would have (226Ra)/

[Ba] slightly lower than that in coexisting plagioclase, 

with (230Th)/[Ba] similar to, or slightly higher than, 

that in plagioclase (fig. 3B). However, because of the 

Zero-age apatite

Zero-age zircon

Zero-age plagioclase

Equilibrium plagioclase

Zero-age amphibole

SH305-1 (Nov. 2004)

SH304-2A1 (Oct. 2004)

EXPLANATION

0.2

0.8

0.6

0.4

1.0

1.2

1.4

0 2.0 4.0 6.0 8.0

(

238U)/(232Th)

0 2,000

1,000

500

0

1,500

2,000

2,500

4,000 6,000 8,000

(230Th)/(232Th)

A

B

To zero-age apatite

To zero-age zircon

To zircon (2,000 yr)

To apatite (2,000 yr)

(

230Th)/[Ba], IN DISINTEGRATIONS PER

YEAR PER MICROGRAM (dpy/μg)

(226Ra)/[Ba], IN DISINTEGRATIONS

PER YEAR PER MICROGRAM (dpy/μg)

1.4

1.3

1.2

1.1

1.00 0.4 0.8 1.2
838 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

similarity of the magnitude of radium and barium partition 

coefficients in amphibole and plagioclase, the effects of a few 

percent amphibole on the 226Ra-230Th-Ba systematics of the 

bulk plagioclase separate would be negligible.

Apatite and Zircon

Thorium and uranium are compatible in apatite and zircon, which results in orders of magnitude difference between 

thorium or uranium concentrations in plagioclase and in these 

accessory phases. Therefore, even small mass fractions of 

apatite and zircon could have significant effects on 238U-230Th- 226Ra disequilibria in the bulk plagioclase separate.

No experimental data exist for partition coefficients of 

the U-series nuclides in apatite or zircon. However, studies of 

coexisting apatite and glass in natural systems yield DTh=1.6–

1.7, with uranium somewhat more compatible (DU=1.8–2.6; 

Luhr and others, 1984; Mahood and Stimac, 1990). Although 

some caution is warranted because these data were collected 

from trachytic to pantelleritic samples, the relative partitioning behavior when comparing two elements should be broadly 

similar in other compositions. This will lead to (238U)/(232Th) 

in apatite that is higher than that in the coexisting liquid at the 

time of crystallization (fig. 3A). On the basis of microbeam 

studies of natural systems, partition coefficients for thorium 

and uranium in zircon are higher than those for apatite, and 

DU /DTh will also be higher than in apatite (DU ~100 and DU /

DTh ~6; Blundy and Wood, 2003). Ion microprobe measurements of 238U-230Th disequilibria in magmatic zircon crystals confirm that (238U)/(232Th) is high, although the range 

of (238U)/(232Th) (~2 to >10, compared to ~1 for coexisting 

glass; for example, Bacon and Lowenstern, 2005; Charlier 

and others, 2003; Reid and others, 1997) is typically greater 

than predicted by equilibrium partitioning and may reflect 

local effects (for example, Bacon and Lowenstern, 2005). On 

the basis of partitioning data and previous measurements of 

magmatic zircon, (238U)/(232Th) in zircon coexisting with the 

October 2004 magma will likely be higher by a factor of 2–6 

than that in the whole rock, and therefore zircon or apatite 

inclusions would elevate (238U)/(232Th) of the bulk plagioclase 

separate compared to that in pure plagioclase. The effect on 

(230Th)/(232Th) will depend on the age of the accessory phases; 

zero-age zircon or apatite will not change (230Th)/(232Th) compared to zero-age plagioclase, but because of their high (238U)/

(232Th) ratios, both zircon and apatite will rapidly evolve to 

higher (230Th)/(232Th) with time. Therefore, addition of older 

zircon or apatite will increase (230Th)/(232Th) in the bulk 

separate. The only way that addition of zircon or apatite could 

explain the low (230Th)/(232Th) in the bulk separate would be 

if they were xenocrystic and had crystallized from a magma 

which itself had unusually low (230Th)/(232Th). 

In detail, this scenario is difficult to reconcile with the 

position of the bulk plagioclase on a 238U-230Th isochron diagram (fig. 3A). Global compilations of U-series disequilibria 

in volcanic rocks (for example, Lundstrom, 2003, and references therein) show that all samples measured to date have 

(

230Th)/(232Th) ratios above ~0.5. The high DU /DTh (and correspondingly high (238U)/(232Th)) for apatite and zircon means 

that even if they crystallized from a magma with (230Th)/(232Th) 

of approximately 0.5, a mixing line between pure plagioclase 

and these accessory phases would not pass through the point 

for the bulk plagioclase separate. Therefore, an additional 

phase or phases with low (230Th)/(232Th) and low (238U)/(232Th) 

(for example, old plagioclase crystals, as argued above) must 

be present in the bulk plagioclase separate. Even if apatite or 

zircon are present, they must make up a small percentage of 

the bulk plagioclase separate. A maximum of 1.4 percent or 

0.025 percent (by weight) of apatite or zircon, respectively, 

could be present if we assume that all of the uranium in the 

bulk plagioclase separate resides in apatite or zircon; the presence of other impurities (such as groundmass) would decrease 

the allowable percentage of these phases.

Although apatite or zircon alone cannot be controlling 

the 238U-230Th disequilibria in the bulk plagioclase separate, 

small mass fractions of these accessory phases may be present 

and could influence the 230Th-226Ra disequilibria of the bulk 

separate. Barium is incompatible in both apatite (DBa <0.3; 

Luhr and others, 1984) and zircon (DBa = 0.003–0.005; Blundy 

and Wood, 2003), and Blundy and Wood (2003) estimate that 

DRa in zircon will be ~10-6. Radium partitioning in apatite has 

not been measured or estimated, but it is likely to be more 

incompatible than barium, and both phases will likely have 

very low (226Ra)/[Ba] at the time of crystallization, effectively 

zero for zircon. However, considering that thorium is compatible in apatite and zircon, both phases will have extremely 

high (230Th)/[Ba] compared to coexisting plagioclase (or liquid), and they will evolve rapidly to high (226Ra)/[Ba]. Mixing 

trajectories for zero-age plagioclase with zircon and apatite are 

shown on figure 3B; contamination with zero-age apatite could 

reproduce the low (226Ra)/[Ba] of the plagioclase separate. 

Older apatite or zircon produce mixing trajectories that have 

slopes too shallow to pass through the data for the bulk plagioclase separate. The interpretation that apatite and zircon are 

not the dominant control on the 226Ra-230Th-Ba systematics of 

plagioclase is consistent with the 238U-230Th data, where apatite 

or zircon in small amounts could contribute to the U-series 

budget of the bulk plagioclase separate, but an additional component (in addition to zero-age plagioclase and zircon/apatite) 

is required to explain the data.

230Th-226Ra Age of October 2004 Plagioclase

The bulk plagioclase separate for the October 2004 

dome (SH305-1) has (226Ra)/(230Th)>1, indicating that at least 

some mass fraction of the crystals measured are significantly 

younger than 10 ka. However, curves for the evolution of melt 

in equilibrium with plagioclase and curves for the evolution 

of the whole rock on a (226Ra)/[Ba] evolution diagram for 

SH304-2A1 do not intersect, even when considering a range 

of preeruptive temperatures and plagioclase compositions and 

the potential effects of melt inclusions in the bulk separate 

(fig. 4). This pattern of data on an evolution diagram could be 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 839

explained if the melt from which (at least some of) the plagioclase precipitated had significantly lower (226Ra)/[Ba] than 

the SH304-2A1 whole rock, but such an explanation would 

require (226Ra)/[Ba] < ~900 dpy/μg in this melt (if we assume 

some impurities in the bulk separate), lower than that observed 

in any recent Mount St. Helens lavas (see fig. 2). It is more 

likely that the plagioclase history and the effects of multiple 

types of impurities are too complex to allow a meaningful age 

determination by this method.

Composition of the October 2004 Plagioclase Separate

On the basis of the mixing relations outlined above, the 

plagioclase separate for the October 2004 dome is composed 

of at least three components: (1) zero-age plagioclase crystals, (2) older plagioclase crystals (whether xenocrysts or 

0 1,000

2,000

0

4,000

6,000

8,000

10,000

2,000 3,000

YEARS BEFORE PRESENT

4,000 5,000 6,000

(226Ra)/[Ba], IN DISINTEGRATIONS

PER YEAR PER MICROGRAM (dpy/μg)

Whole rock

An40, 850°

An50, 850°

An50, 850°-wr

An40, 900°

Figure 4. 226Ra evolution diagram for October 2004 dacite 

(SH304-2A1). Curves shown are for evolution of (226Ra)/[Ba] over 

time in whole rock (bold line) and for melt in equilibrium with 

plagioclase with composition of An40–An50 at temperatures of 

850–900°C. In a case where plagioclase crystallized rapidly 

within the past ~10,000 years in equilibrium with a liquid having 

the composition of the whole-rock, the intersection of curves 

for melt in equilibrium with plagioclase and whole rock would 

indicate the crystallization age (see Cooper and others, 2001, 

for details of the technique). The fact that curves for melt in 

equilibrium with plagioclase in SH304-2A1 do not intersect the 

whole-rock curve indicates that the bulk plagioclase separate 

and melt were not in equilibrium, even considering a range 

of potential crystallization temperatures and compositions of 

plagioclase. Correcting the plagioclase composition for the 

effects of having impurities (groundmass, zircon, or apatite) in 

the bulk separate makes the disequilibrium between melt and 

plagioclase more extreme. An example is shown by the long-dash 

curve, calculated assuming 3 percent whole-rock in the bulk 

separate, which is broadly consistent with concentrations of Th 

in the plagioclase being below detection limit for laser-ablation 

ICP-MS analyses (Kent and others, this volume, chap. 35). The 

lack of equilibrium between plagioclase bulk separate and the 

likely melt composition suggests that at least some of the crystals 

are foreign to the host liquid (see text for discussion).

antecrysts), and (3) some groundmass or melt inclusions and/

or a minute mass fraction of apatite or zircon. Groundmass, 

apatite, and zircon will act in a similar way on the 238U-230Th 

disequilibria, shifting the bulk plagioclase separate toward 

high (238U)/(232Th) and high (230Th)/(232Th) ratios. The volume of crystals must be dominated by young plagioclase, but 

there must also be some component of older plagioclase with 

low (230Th)/(232Th) in order to explain the low (230Th)/(232Th) 

measured in the bulk separate. The exact age of these foreign 

crystals cannot be uniquely determined; therefore, they may 

represent xenocrysts or antecrysts.

Plagioclase in November 2004 Dacite

In contrast to the low (230Th)/(232Th) measured in the 

October 2004 plagioclase separate, the (230Th)/(232Th) ratio 

measured in plagioclase separated from the November 2004 

dome sample (SH305-1) is approximately 4 percent higher 

than that in the host magma. As a result, a two-point plagioclase-whole-rock isochron has a negative slope, yielding a 

negative apparent age. This observation is inconsistent with 

crystallization of plagioclase solely from the SH305-1 host 

liquid, but (as with plagioclase in the October 2004 dome 

sample) it is consistent with the presence of impurities in the 

bulk plagioclase separate, which may include an antecrystic or 

xenocrystic plagioclase component. Interestingly, this foreign 

component must have higher (230Th)/(232Th) than zero-age 

plagioclase and, therefore, is compositionally distinct from 

that present in the October 2004 dome sample, despite the 

fact that the two dome samples were erupted only a month 

apart. The plagioclase separate for the November 2005 dome 

(SH305-1) also has higher (226Ra)/[Ba] and (230Th)/[Ba] than 

other plagioclase from Mount St. Helens samples (with the 

exception of Volpe and Hammond’s (1991) measurement of 

MSH90-9; fig. 2), and those ratios are higher than those which 

would be in equilibrium with the host magma. As argued 

above for the October 2004 dome, it is likely that the bulk 

plagioclase separate for the November 2004 dome contains 

some zero-age plagioclase, and in addition, the separate 

could contain older plagioclase, zircon, apatite, amphibole, 

or groundmass in small percentages. The whole-rock October 

2004 and November 2004 dome samples have similar U-series 

characteristics; therefore, the effects of zero-age impurities in 

the bulk plagioclase separate for the November 2004 dome 

would be similar to those shown in figure 3 for the October 

2004 dome. In the case of the November 2004 plagioclase, the 

combined 238U-230Th-226Ra-Ba data for the bulk separate are 

consistent with the presence of a small percentage of apatite, 

zircon and/or melt inclusions (fig. 3). The presence of less 

than 1 percent apatite or zircon (±melt) inclusions within the 

plagioclase separate could explain high (238U)/(232Th), (230Th)/

[Ba], and thorium concentrations in the bulk separate; and if 

the zircon or apatite were a few thousand years old, it could 

also explain the high (230Th)/(232Th) and (226Ra)/[Ba]. The high 

(

230Th)/(232Th) and (226Ra)/[Ba] could still be consistent with the 
840 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

presence of some old plagioclase in the bulk separate, if offset 

by the presence of zircon or apatite, but the mixing proportions 

are difficult to determine because of the complexity of the data. 

Given the negative apparent 230Th-238U age and the near-equilibrium 226Ra-230Th-Ba data for the SH305-1 plagioclase separate, 

it is difficult to constrain the age of crystallization of any old 

component of plagioclase that may be in the bulk separate.

Implications for Magma Reservoir Processes

Age considerations aside, it is striking that plagioclase 

separates from the October and November 2004 dome samples 

have quite different U-series characteristics, despite having 

erupted only one month apart. Evidently, these two magmas 

entrained different populations of crystals, implying significant heterogeneity within the reservoir system or different 

transport paths intersecting different resident crystals.

In the case of the October 2004 dome, we can quantify the proportion of old and young crystals, but we cannot 

uniquely determine the age (and therefore the origin) of the 

old component. However, we speculate that the percentage 

of old crystals required (~15–50 percent of the plagioclase 

separate, depending on Th concentrations and age of the old 

component) might be higher than is easily explained by disaggregation of xenoliths that are present only in small amounts 

(1–5 percent of the bulk sample). Considering that 2004–5 

dacites have a total of 30–40 percent plagioclase by volume, 

the assumption that all old crystals represent disaggregated 

xenoliths would imply that ~5–20 percent of the volume of the 

whole-rock dacite is composed of these xenocrysts. This, in 

turn, would imply that at least 1–5 times the volume of xenoliths that are observed has been completely disaggregated and 

incorporated into the dacite, which would likely produce some 

trace-element signature of plagioclase addition. Furthermore, 

we suggest that incorporation of plagioclase from a partially 

molten crystal mush perhaps may be physically easier than 

disaggregation of completely solidified xenoliths, and that an 

antecrystic origin is likely for at least some of the old crystals 

within the bulk separate. It is likely that there are at least three 

plagioclase components within the October 2004 dome sample: (1) zero-age plagioclase, (2) antecrystic plagioclase disaggregated from crystal mush, and (3) xenocrystic plagioclase 

disaggregated from gabbroic xenoliths. The age and origin of 

the old component(s) may be better constrained by analysis of 

different size fractions of plagioclase within the dome samples 

(which could preferentially sample different crystal populations), and work is currently underway to analyze different 

size fractions of plagioclase in SH304-2A1.

We have found evidence for a xenocrystic or antecrystic 

component in plagioclase separates in both of the samples 

from the 2004–5 eruption for which U-series disequilibria in 

plagioclase have been measured, as well as in at least two of 

the older Mount St. Helens lavas of Castle Creek age (Cooper and Reid, 2003). Thus, entrainment and disaggregation 

of crystal mush or xenoliths appears to be a common, if not 

ubiquitous, process in the Mount St. Helens reservoir system. 

Furthermore, the heterogeneity of this foreign component 

within different samples from the 2004–5 eruption suggests 

that the reservoir system feeding the current eruption is complex and not well mixed.

U-Series Disequilibria in Gouge

The exterior of the Mount St. Helens dome is coated 

with meter-thick fault gouge, which appears to consist mainly 

of fragments of dacite dome, as large as several centimeters, 

within a matrix of finer fragments (Cashman and others, this 

volume, chap. 19). Gouge samples have 230Th/232Th indistinguishable from that of dacite dome samples (fig. 1B), consistent with an origin for the gouge as being dominated by 

fragmented juvenile dacite. However, lower U/Th and lower 

U and Th concentrations in gouge compared to dome samples 

suggests either loss of U and Th (with preferential loss of U) 

from the gouge material or dilution of dacite fragments within 

the gouge with some material lower in U and Th concentrations than dome rocks. Similarly, gouge samples have lower 

(

230Th)/[Ba] than dome samples, but they have (226Ra)/[Ba] 

similar to dome samples (fig. 2B), suggesting loss or dilution 

of Th rather than gain of Ba. Fluxing of gouge with a vapor 

or fluid could modify trace-element composition of the gouge 

either by precipitation of phases from the fluid or by leaching 

of fluid-mobile elements from gouge material. Fluids would be 

expected to carry U, Ra, and Ba, but not Th; thus, addition of 

a fluid-derived phase would be expected to add U, Ra, and Ba 

to the gouge rather than to decrease concentrations relative to 

the dacite. Leaching of dacite by a fluid could remove U, Ra, 

and Ba but could not explain the lower concentrations of Th in 

gouge. Therefore, the observations are most consistent with the 

presence of a component in the gouge (as yet unidentified) that 

has low concentrations of U and Th (and, in the case of SH320-

1, low concentrations of Ra and Ba) compared to dome dacite.

Conclusions

Juvenile material from the 1980–86 eruption shows heterogeneity in 230Th/232Th ratios. The total range in 230Th/232Th 

measured in 1980–86 samples overlaps but is larger than the 

range of 230Th/232Th ratios measured in 2004–5 dacite and dome 

samples. However, the differences in crystal compositions 

between 1980s and 2004–5 samples indicates substantial heterogeneity of plagioclase crystals sampled during the two eruptions 

and even within a short time period during the 2004–5 eruption. 

This heterogeneity precludes an origin for the 2004–5 magma 

solely by remobilization of remnant magma from the 1980–86 

eruptive cycle. Instead, the U-series data are consistent with 

a model in which the 2004–5 magma represents a mixture of 

some remnant 1980s magma with plagioclase crystals (and perhaps associated magma) having an isotopic composition previously unidentified in the erupted products of Mount St. Helens. 

The U-series data are consistent with an origin for the 2004–5 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 841

magma by partial melting of the lower crust, as has been proposed previously for Mount St. Helens dacites. The preservation 

of significant 226Ra excesses in all measured whole-rock samples 

of the 2004–5 eruption indicates that the magma was generated 

recently, likely within the past few thousand years.

In detail, plagioclase separates for October 2004 and 

November 2004 dome samples (SH304-2A1 and SH305-1) 

have different U-series characteristics. This suggests that the 

two dacites, erupted only about a month apart, sample different populations of plagioclase crystals and/or different types of 

impurities within the bulk plagioclase separates. Plagioclase in 

both samples is also different in 230Th-238U disequilibria from 

that measured in older Mount St. Helens samples. Because of 

the complexities within the crystal populations sampled in the 

2004–5 dacites, exact ages for plagioclase cannot be calculated. However, the preservation of 226Ra excess in plagioclase 

from the October 2004 dome (SH304-2A1) requires that some 

mass fraction is younger than ~10 ka, and other petrologic 

constraints suggest that at least some crystallization occurred 

shortly before or during magma ascent. Simple mass balance 

calculations suggest that the percentage of a zero-age component is ~60–80 percent, although it could be lower if the older 

plagioclase component is <10 ka, and/or if the young component has (226Ra)/[Ba] higher than predicted in equilibrium 

with the whole rock. The bulk plagioclase separate for the 

November 2004 dome (SH305-1) is within error of 226Ra-230Th 

equilibrium, and a young component (<10 ka) is not required 

by the 226Ra-230Th data. However, neither is it precluded by 

the data, and any older plagioclase is likely less than ~20 ka, 

if it crystallized from a magma with 230Th/232Th similar to that 

measured in Mount St. Helens whole-rock samples.

The combined whole-rock and plagioclase data for 

the 1980–86 and 2004–5 eruptions suggest that the crystal 

populations sampled by Mount St. Helens magmas commonly 

include a recycled component (whether antecrystic or xenocrystic) and that the magma-storage region beneath Mount St. 

Helens is complex, preserving chemical distinctions between 

magmas erupted only a short time apart. Thus, the crystals 

may contain a longer record of processes within the magma 

reservoir beneath Mount St. Helens than do the liquid fractions 

of these magmas. Detailed studies of the crystals may therefore provide insights into the temporal and spatial distribution 

of magmas and crystal mushes within the reservoir system.

Acknowledgments

We would like to thank the staff at the Cascades Volcano 

Observatory for providing samples, data, and information 

about the eruption and for organizing workshops and conference sessions as a venue to discuss recent results. We thank 

the other members of the Mount St. Helens petrology working 

group for helpful discussions and open sharing of data and 

ideas. J. Lowenstern and M. Reagan provided very constructive and helpful reviews. This work was partially supported by 

NSF EAR-0307691 to K.M. Cooper.

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844 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Appendix 1. Analytical Methods

Chemical Separations

Procedures for chemical separation were modified from 

those described by Goldstein and Stirling (2003) and Pietruszka 

and others (2002). Plagioclase separates were prepared using 

standard magnetic separation followed by hand-picking using a 

binocular microscope. Samples (0.5–1.5 g) of mineral separates 

and whole rocks were cleaned before dissolution using successive baths of distilled acetone, deionized, reverse-osmosis (RO) 

water, 0.1N HCl + 2 percent H2O2, RO water, and distilled 

acetone. Samples were dissolved in closed Savillex beakers on 

a hotplate using a 3:1 mixture of concentrated HF:HNO3. After 

dissolution, perchloric acid and saturated boric acid were added 

and samples were evaporated. Samples were treated a second 

time with perchloric and boric acids, evaporated again, dissolved in 3 N HCl + saturated boric acid, and centrifuged. The 

solution was decanted, and any residue was dissolved in additional 3N HCl + boric acid (in the proportions 100 mL HCl:5 

mL saturated boric acid); this process was repeated until no 

residue was visible after centrifuging. Sample solutions were 

split into three aliquots: one small aliquot (corresponding to 

~10–100 mg of rock or mineral) was spiked with 233U and 229Th 

for U and Th concentration measurement by isotope dilution, 

a second small aliquot (corresponding to ~5–10 mg of rock or 

mineral) was spiked with 135Ba for Ba concentration measurement by isotope dilution, and the largest aliquot (generally the 

remainder of the solution, corresponding to ~0.5–1.5 g of rock 

or mineral) was spiked with 228Ra and used for measurement of 

Ra concentration by isotope dilution and of U and Th isotopic 

compositions. All resins used in the chemical separations were 

precleaned with nitric and hydrochloric acids and RO water, 

and they were also cleaned and conditioned on the column with 

the acids used for sample loading and elution.

For the U-Th-Ra aliquot, radium (with major and most 

trace elements) was separated from Th and U using a column 

packed with 0.5 mL of Eichrom “pre-filter” (inert beads) and 

1–2 mL of Eichrom TRU resin. Samples were loaded in 2–8 

column volumes (cv) of 7M HNO3

; the load, together with 

subsequent washes with 2 cv of 7M HNO3 and 10 cv of 1.5M 

HNO3, was collected and saved for further purification of 

Ra. This was followed by a wash of 8 cv 3N HCl, which was 

discarded. Th was eluted from the column in 8 cv 0.2N HCl, 

followed by U elution from the column in 8 cv 0.1N HCl + 

0.05N HF. A second pass through a small TRU column (0.25 

mL pre-filter plus 0.5 mL TRU resin) was necessary to remove 

residual Th from the U fraction.

The thorium and uranium isotope dilution aliquots were 

spiked with ~1 ng each of 229Th and 233U tracers. Spiked 

samples were equilibrated by heating for ~48 hours followed 

by evaporation. Some samples (1980s samples) were treated 

with 1–2 drops of perchloric acid during equilibration in order 

to aid sample-spike equilibration. Thorium and uranium were 

separated using a column loaded with 0.25 mL Eichrom prefilter and 0.5 mL TRU resin. A single pass through the column 

was generally sufficient to separate U and Th from the matrix.

The radium fraction from the TRU column was dried and 

redissolved in 1N HCl. Radium was separated from the major 

elements and most trace elements using a 5–10 mL column 

loaded with cation-exchange resin (Dowex AG 50-X8). After 

loading in 1–3 cv 1N HCl, the column was washed with 0–2 

cv 1N HCl (adjusted so that the total load + wash was 3 cv), 5 

cv 2.0N HCl, 1.5 cv 2.5N HCl, and 1.5 cv 3.0N HCl. The Ra 

(+Ba) fraction was eluted from the column in 8 cv 6N HCl. For 

samples larger than 0.5 g, a second, 0.5–1 mL cation-exchange 

column was run following the same procedure. Radium was 

separated from barium using a Teflon column packed with 

0.45 mL Eichrom Sr-spec resin. The sample was dissolved in 

2 cv 2.5N HNO3

 for loading on the column. After loading, Ba 

was washed from the column with 0.75 cv 2.5N HNO3 and Ra 

was eluted from the column in 6 cv 2.5N HNO3. This column 

procedure was repeated an additional one to two times, followed 

by passing the Ra cut through a column loaded with 0.5 mL 

of Eichrom prefilter before evaporating. A final cation column 

(0.25 mL) was run 1–4 days prior to mass spectrometry in order 

to separate 228Th from 228Ra. The sample was loaded on this 

column in 1 cv 6N HCl; radium was eluted in this fraction plus 

an additional 6 cv of 6N HCl while Th remained on the column.

The barium isotope dilution aliquot was spiked with a 135Ba-enriched tracer, equilibrated on a hotplate for at least 48 

hours, and dried. Barium was separated using a column loaded 

with 0.65 mL of cation exchange resin (Dowex AG 50-X8). 

The sample was loaded on the column in 0.65 mL 1.5N HCl (1 

cv), followed by washes with 1.5N and 2N HCl. Barium was 

eluted in 2.5 N HNO3.

Total-process blanks during the time that the Mount St. 

Helens samples were run were <20 pg Th and <15 pg U for the 

large (isotopic composition) aliquots, and <4 pg Th and <6 pg 

U for the small (isotope dilution) aliquots (compared to typical 

sample sizes of 0.1–2 μg Th and 0.02–1 μg U in the isotopic 

composition aliquots and 10–50 ng Th and 2–20 ng U in isotope dilution aliquots). Barium blanks were 5–10 ng, compared 

to typical sample sizes of 1–2 μg Ba. Radium blanks were 

always below the limit of detection (~0.1 fg), and sample sizes 

were a few tens of femtograms Ra for minerals and 200–500 fg 

Ra for whole-rock samples. Blanks for each measurement were 

subtracted before calculation of concentrations; in most cases 

these had a negligible effect on calculated concentrations, but 

there were small corrections to Ba concentrations.

Mass Spectrometery

All measurements of isotopic ratios for chemically separated samples were performed with a Nu Plasma multicollector-ICP-MS (MC-ICP-MS) at the University of Washington. 
36. 238U-230Th-226Ra Disequilibria in Dacite and Plagioclase from the 2004–2005 Eruption of Mount St. Helens 845

This instrument is equipped with three ion counters, one of 

which (IC0; the high-mass ion counter) is equipped with an 

energy filter to reduce tailing effects of major peaks. All analyses with the exception of Ba measurements were performed 

using a desolvating nebulizer (dry plasma), which improves 

sensitivity by a factor of ~10 compared to use of a peristaltic 

pump (wet plasma). Ba abundances were high enough that 

the desolvating nebulizer was unnecessary, and samples were 

introduced as a wet plasma.

Uranium Isotopic Composition Measurements

234U/238U ratios were measured using a static analysis routine, with 238U measured on a Faraday cup and 234U measured 

on an ion counter. Instrumental mass bias and ion counter-faraday gain were calibrated using NIST uranium standard CRM112A with reference to the accepted 238U/235U and 234U/238U 

ratios of 137.88 and 5.286×10-5, respectively (Cheng and others, 2000). Rock and solution standards analyzed as unknowns 

were interspersed with unknowns during a day of analysis, and 

re-calibrations of mass bias and gain were performed every 

third or fourth sample. Rock standards known to be in radioactive equilibrium reproducibly yielded (234U)/(238U) ratios within 

4–6 per mil of 1.000 (see table 1), which we therefore consider 

to be our analytical uncertainty for unknowns.

Thorium Isotopic Composition Measurements

230Th/232Th ratios were measured using a static analysis routine with 232Th measured on a Faraday cup and 230Th measured 

on the high-mass ion counter equipped with an energy filter 

(IC0). At the beginning of each day of analysis of 230Th/232Th 

ratios, the voltage for the energy filter was tuned in order to 

optimize the abundance sensitivity, and the instrumental mass 

bias and ion counter-faraday gain were calibrated using NIST 

uranium standard CRM-112A. Abundance sensitivity was measured at 400–500 ppb at one AMU during the period of analyses. 

Correction for tailing of 232Th into the 230Th mass range was done 

using an exponential fit to the tail, which was measured each 

day using a Th solution standard (Th ‘U’ or WUN-1). Solution standards for Th measured as unknowns (Th ‘U’, WUN-1 

(sometimes called ZSR), and IRMM-035) were generally within 

1 percent of accepted or certified values (table 1). Rock standards (TML, BHVO-1, BCR-2), which went through the same 

chemical separation procedure as unknowns and were analyzed 

as unknowns, yielded (230Th)/(238U) within 1 percent of radioactive equilibrium, or within error of previously measured values 

(table 1). Therefore, we consider our measurement accuracy and 

reproducibility on unknowns to be ~1 percent.

Ba Concentration Measurements by Isotope 

Dilution

Isotopic compositions of sample-spike mixtures were 

measured using wet plasma and a static routine where all peaks 

were measured using Faraday cups. Ratios were normalized to 

the natural ratio of 136Ba/138Ba (=0.10954) to correct for instrumental mass bias. Normalized 137Ba/138Ba ratios were monitored, 

and they agreed with the natural ratio to within <0.5 percent 

for samples. NIST SRM 3104a (barium normal) was run as an 

unknown throughout the day and always yielded 137Ba/138Ba and 135Ba/138Ba ratios that were within 0.12 percent and 0.65 percent 

of the natural ratios, respectively. The barium spike concentration was calibrated with reference to NIST SRM 3104a, which 

has concentration certified to 0.3 percent at the 95-percent confidence level. Replicate spike calibration measurements agreed 

to within 0.3 percent. Measurement of USGS standard BCR-2 

agreed with the accepted value within 1.6 percent and replicates 

of the same unknowns agreed within 2 percent. Therefore, measurement precision is better than ~0.5 percent, whereas reproducibility and accuracy is better than 2 percent.

U and Th Concentrations by Isotope Dilution

233U/238U and 229Th/232Th ratios were analyzed separately 

using static routines on Faraday cups. Mass bias was calibrated before analysis and several times over the course of a 

day of measurements using NIST CRM112A. 233U and 229Th 

spike concentrations were calibrated using ClaritasTM Th and 

U standard solutions, certified to 0.5 percent, and calculations 

of concentrations in unknowns included corrections for totalprocess chemical blanks and evaporation corrections for the 

spikes. Accuracy was checked by analysis of rock standards 

BHVO-1, BCR-2 and TML. Our measurements of Th and 

U concentrations and (230Th)/(238U) of BHVO-1 agree within 

~1 percent with published measurements of KIL1919 (collected from the same sample locality as BHVO; Pietruszka 

and others 2002); both are within error of the recommended 

value for Th concentration of 1.2±0.3 ppm for USGS standard 

BHVO-2 and are likely within error of the recommended value 

for Th concentration in BHVO-1 (of 1.1 ppm), although no 

error is reported for this value. Certified concentrations of U in 

BHVO-1 or BHVO-2 are not available. Rock standard TML is 

known to be slightly heterogeneous in Th and U concentration 

but to be in radioactive equilibrium; BCR-2 is a sample of a 

Columbia River Basalt flow erupted before 6 Ma and, therefore, should also be in radioactive equilibrium. Our analyses 

of TML and of BCR-2 yielded (230Th)/(238U) ratios within 

~1 percent of equilibrium; therefore, the combined errors on 

measured U/Th and 230Th/232Th ratios must be better than 1 

percent. Thus, accuracy of our measurements of U and Th 

concentrations must also be better than 1 percent.

Ra Concentrations by Isotope Dilution

226Ra/228Ra measurements were made using two different routines during this study. 226Ra/228Ra ratios in the first 

samples run (including SH304-2A1 whole-rock and plagioclase, 

SH305-1 whole-rock and plagioclase, and SH311-1B wholerock) were measured using a static analysis routine in which 
846 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

228Ra and 226Ra peaks were measured on ion counters IC0 and 

IC1, respectively. Mass bias and relative ion counter gains were 

calibrated using NIST CRM112A. After the initial group of 

samples had high background counts at mass 225.5, we began 

scanning all samples in the mass range 224.5–229.5 for interferences and found that backgrounds measured at masses 225.5 

and 227.5 were significantly lower when using IC0 (which is 

equipped with an energy filter) than when using IC1. When 

using IC0, small peaks of ~10 cps are present in some regions of 

the mass range, but interferences and/or instrumental background at the half-masses used for background measurements 

were negligible (tenths of counts per second to a few cps, compared to count rates on peaks of hundreds of cps to thousands 

of cps for most samples). Furthermore, the ion counter gain 

for IC0 was more stable than that for IC1, so we switched to a 

dynamic routine where both peaks were measured on IC0. This 

also has the advantage that ion counter gain calibrations are less 

critical as long as gains are stable over the time scale of a sample run (~10–15 minutes). The disadvantages of this approach 

are that the total counting time for each peak is reduced and 

that the analyses are susceptible to intensity variations due to 

instability of the plasma; the latter effect was minimized by 

adding a small amount (a few tens of picograms) of 232Th (as a 

solution prepared from Ames Th metal) to each unknown and 

normalizing Ra count rates to the intensity of 232Th measured 

on a Faraday cup during the same cycle. This analysis routine 

produces better in-run precision (~0.5–1 percent SE, for peaks 

of ~200–3000 cps) than the initial routine (~2 percent SE). 

Concentration of 228Ra in the spike was calibrated with respect 

to NIST SRM 4965, which has Ra concentration certified to 

1.23 percent (3σ). (226Ra)/(230Th) measured in rock standard 

TML differs from equilibrium by less than 1 percent, and 

(226Ra)/(230Th) in USGS standard BHVO-1 is within ~2 percent 

of the previously measured value (Pietruszka and others, 

2002). Taking these measurements and the in-run precision 

into account, we consider our overall uncertainty at the 2σ

level to be ~3–5 percent for the early-run samples, and ~1–2 

percent for the later samples.
A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004–2006

Edited by David R. Sherrod, William E. Scott, and Peter H. Stauffer

U.S. Geological Survey Professional Paper 1750, 2008

Chapter 37

1

 Department of Geoscience, 121 Trowbridge Hall, University of Iowa, 

Iowa City, IA 52242

2

 Department of Earth and Space Sciences, University of Washington, Box 

351310, Seattle, WA 98195; now at Geology Department, University of 

California, One Shields Ave., Davis, CA 95616

3

 U.S. Geological Survey, 1300 SE Cardinal Court, Vancouver, WA 98683

Timing of Degassing and Plagioclase Growth in Lavas 

Erupted from Mount St. Helens, 2004–2005, from 

210Po-210Pb-226Ra Disequilibria

By Mark K. Reagan1, Kari M. Cooper2, John S. Pallister3, Carl R. Thornber3, and Matthew Wortel1

Abstract

Disequilibrium between 210Po, 210Pb, and 226Ra was 

measured on rocks and plagioclase mineral separates erupted 

during the first year of the ongoing eruption of Mount St. 

Helens. The purpose of this study was to monitor the volatile 

fluxing and crystal growth that occurred in the weeks, years, 

and decades leading up to eruption. Whole-rock samples were 

leached in dilute HCl to remove 210Po precipitated in open 

spaces. Before leaching, samples had variable initial (210Po) 

values, whereas after leaching, the groundmasses of nearly all 

juvenile samples were found to have had (210Po) ≈ 0 when they 

erupted. Thus, most samples degassed 210Po both before and 

after the magmas switched from open- to closed-system degassing. All juvenile samples have (210Pb)/(226Ra) ratios within 

2σ of equilibrium, suggesting that the magmas involved in the 

ongoing eruption did not have strong, persistent fluxes of 222Rn 

in or out of magmas during the decades and years leading to 

eruption. These equilibrium values also require a period of at 

least a century after magma generation and the last significant 

differentiation of the Mount St. Helens dacites. Despite this, 

the elevated (210Pb)/(226Ra) value measured in a plagioclase 

mineral separate from lava erupted in 2004 suggests that a 

significant proportion of this plagioclase grew within a few 

decades of eruption. The combined dataset suggests that for 

most 2004–5 lavas, the last stage of open-system degassing 

of the dacite magmas at Mount St. Helens is confined to the 

period between 1–2 years and 1–2 weeks before eruption, 

whereas plagioclase large enough to be included in the mineral 

separate grew around the time of the 1980s eruption or earlier.

Introduction

Open- and closed-system degassing occurring during 

the century before an eruption can be monitored by measuring the relative activities of short-lived radionuclides, such as 

polonium-210 (210Po; t

1/2 = 138.4 days), lead-210 (210Pb; t1/2 = 

22.6 years), and radium-226 (226Ra; t

1/2 = 1,599 years) in lavas. 

For example, Po partitions efficiently into exsolving volatile 

phases and commonly degasses from both mafic and silicic 

lavas, which results in (210Po)/(210Pb) << 1.0 in erupted lavas. 

Examples of lavas that have (210Po)/(210Pb) < 0.15 include those 

erupted from Etna (Lambert and others, 1985; Le Cloarec and 

Pennisi, 2001), Hawai‘i (Gill and others, 1985), Arenal (Gill 

and others, 1985; Reagan and others, 2006), Mount St. Helens 

in 1980 (Bennett and others, 1982), and Anatahan (Reagan and 

others, 2005). Excesses of 210Po over 210Pb have been observed 

in some phreatomagmatic tephras because of condensation of 

210Po from gases streaming through the shallow conduit system 

of a volcano before eruption (Reagan and others, 2005).

Radon-222 (222Rn), which lies between 226Ra and 210Pb in 

the 238U decay series and has a half-life of 3.82 days, strongly 

partitions into the gas phase (Gill and others, 1985), whereas 

about 99 percent of Pb stays within the melt (Gauthier and 

others, 2000). Nevertheless, significant deficits of 210Pb with 

respect to 226Ra have been observed in magmas with widely 

varying compositions. Such deficits can result from differential partitioning of these elements between coexisting 

mineral and melt phases (Williams and others, 1986; Gill and 

Williams, 1990; Condomines and others, 1995; Sigmarsson, 

1996), and/or from persistent (year- to decade-scale) losses of 

222Rn by degassing (Gauthier and Condomines, 1999; Gauthier 
848 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

and others, 2000; Le Cloarec and Pennisi, 2001; Le Cloarec 

and Gauthier, 2003; Turner and others, 2004; Berlo and others, 

2004). Deficits in (210Pb) with respect to (226Ra) in midoceanridge basalts have been attributed to lower partition coefficients for Ra compared to Pb during magma generation (Rubin 

and others, 2005).

Although large excesses of 210Pb with respect to 226Ra 

have been reported at some arc volcanoes, including the 1980s 

lavas and tephras from Mount St. Helens (Berlo and others, 

2004), (210Pb/226Ra) values for lavas erupting at many other arc 

volcanoes are within 2σ error of equilibrium when the error is 

propagated back to the age of eruption. These near-equilibrium 

lavas include the silicic andesite tephras erupted from Anatahan in 2004 (Reagan and others, 2005), basalts and andesites 

erupted from Klyuchevskoy and Bezymianny in Kamchatka 

(Turner and others, 2007), and most of the basaltic andesites 

that have erupted from Arenal since 1968 (Reagan and others, 

2006). The frequency of near-equilibrium (210Pb)/(226Ra) values 

in lavas probably reflects the special circumstances needed to 

supply enough 222Rn to accumulate significant 210Pb excesses 

(Reagan and others, 2006).

The regular sampling of the ongoing eruption of Mount 

St. Helens by the U.S. Geological Survey has provided an 

exceptional suite of samples for applying 226Ra-210Pb-210Po disequilibria to constrain the time scales of degassing and other 

magmatic processes that have occurred within the century 

before eruption. Silicic glass inclusions from lavas erupted 

at Mount St. Helens have as much as 6 weight percent H2O 

(Blundy and Cashman, 2005), and degassing of this water has 

the potential to generate significant disequilibrium between 

these short-lived nuclides.

Samples and Analytical Procedures

Whole-rock powders analyzed here span the period 

of the ongoing eruption from October 2004 to December 

2005. Most samples are solid juvenile dome samples, but 

the samples include the phreatomagmatic tephra erupted in 

October 4, 2004, and March 8, 2005, and two samples of the 

dome-coating gouge. All dome samples were collected within 

2 months of venting at the surface. Other chapters in this 

volume describe the mineral contents, textures, and chemical 

compositions of the samples analyzed here in detail (Pallister 

and others, this volume, chap. 30; Thornber and others, this 

volume, chap. 32). Briefly, all juvenile samples are dacites 

with approximately 65 percent SiO2

, 1.4 percent K2O, and 

FeO*/MgO ≈2. These samples are highly porphyritic, with 

approximately 50 percent plagioclase, 3–5 percent hornblende 

and orthopyroxene, and 1 percent or less each of magnetite and augite phenocrysts, typically in a microlite-choked 

groundmass. Gabbroic to granodioritic inclusions, some 

in varying stages of disaggregation, are present in varying 

amounts in the dacites. These inclusions were avoided in the 

sampling for this study.

The October 2004 tephra sample (MSH04E2A03_A2) 

consists of finely pulverized rock fragments, with a variety of 

mineral textures, reflecting sources for the particles in older 

Mount St. Helens rocks, as well as juvenile dacite (Rowe and 

others, this volume, chap. 29). This sample also had organic 

debris that was removed by hand picking before analysis. The 

March 2005 tephra sample (MSH05DRS_3_9_4) consisted 

entirely of fragmented young dome debris and had a high proportion of clasts with lengths from 2 to 6 mm. The coarseness 

of this sample allowed us to hand-pick fragments for additional fragmentation in a ceramic rotary-ring mill and analysis.

Dome sample preparation began by fragmentation to 

particle lengths of several millimeters to 2 cm. Unaltered 

fragments were hand picked for ultrasonic washing in purified water and coarse grinding in a ceramic rotary mill. Gouge 

sample SH307-2A was indurated, whereas SH314-1G consisted of a collection of particles ranging from silt to about 

2 mm in length. The SH307-2A sample was broken up and 

then coarsely ground. Sample SH314-1G was not additionally 

comminuted. Except for the initial tephra sample and the first 

dome sample, all samples were leached with 0.5 N HCl for 

10–15 minutes in an ultrasonic agitator, followed by centrifuging and rinsing and centrifuging twice in purified water. This 

was done to remove 210Po that condensed on surfaces of cracks 

and vesicles, which proved crucial for determining the initial 

groundmass (210Po) value.

Plagioclase was separated from sample SH304-2A using 

standard density and magnetic techniques after grinding and 

sieving to segregate particles between 170 mesh and 80 mesh. 

This separate was washed with a 0.5 N HCl–2 percent H2O2

mixture for 10 minutes, followed by washing in purified water 

before analysis.

Analytical techniques for 210Po are discussed in detail 

elsewhere (Reagan and others, 2005; Reagan and others, 

2006). Initial (210Po) and (210Pb) values and errors for samples 

were calculated using exponential regressions through multiple analyses employing a half-life for 210Po of 138.4 days 

(Holden, 1990). Rock standard RGM-1 was analyzed six times 

over the period of analysis. The weighted mean (210Po) value 

and standard deviation for these analyses are 4.20±0.04 (1σ) 

decays per minute per gram (dpm/g), which is similar to the 

measured values for (238U) for this standard (Le Fevre and Pin, 

2002). The concentrations of 226Ra in samples discussed here 

are from Cooper and Donnelly (this volume, chap. 36).

Other Time-Series Data

The variation in major-element, trace-element, and longlived radiogenic isotopic compositions for juvenile lavas from 

the ongoing eruption has been minimal (Pallister and others, 

this volume, chap. 30). However, (230Th)/(238U) values vary 

from 0.93 to 1.10 at a nearly constant Th isotopic composition 

in dome samples, which has been attributed to fractionation of 

U from Th during melting to generate the Mount St. Helens 
37. Timing of Degassing and Plagioclase Growth in Lavas Erupted from Mount St. Helens, 2004–2005 849

dacites (Cooper and Donnelly, this volume, chap. 36). This 

Th isotopic composition is similar to the values measured for 

lavas erupted from Mount St. Helens in the middle 1980s, suggesting that there is a genetic link between the lavas. However, 

the Th isotopic compositions of plagioclase separates from 

lavas from the ongoing eruption range from 1.026 to 1.306 

(whole-rock values are 1.26–1.28), and some of these values 

are different than any values measured for 1980s and 2004–5 

lavas, indicating that an additional older magma component is 

involved in the ongoing eruption (Cooper and Donnelly, this 

volume, chap. 36). A wide range for the average ages of plagioclase mineral separates is supported by their (226Ra)/(230Th) 

values, which range from 1.51 for plagioclase from sample 

SH304-2A to 1.04 for a plagioclase from SH305-1 (Cooper 

and Donnelly, this volume, chap. 36).

210Po-210Pb-226Ra disequilibria

Whole tephras erupted on October 4, 2004, were 

enriched in 210Po over 210Pb by more than a factor of 2, indicating that significant magmatic 210Po had condensed in the 

shallow hydrothermal system before the phreatomagmatic 

explosions began (Reagan and others, 2005). This contrasts 

with tephras ejected during the phreatomagmatic explosions 

of March 8, 2005 (MSH05DRS_3_9_4), which were largely 

degassed of 210Po when they erupted and had very little leachable 210Po (table 1; fig. 1).

The first lavas ejected at the vent (for example, sample 

SH301-1A7, table 1) had a small excess of 210Po over 210Pb 

values and (210Pb) that approximately matched (210Pb) values 

for samples erupted in the 1980s. We interpret these data to 

indicate that this sample was rock excavated from the dome 

materials erupted in the 1980s that were in the vent area. The 

small excess of 210Po for this sample was likely deposited in 

open spaces from fumarolic gases streaming through the vent 

area just before eruption. This observation was one of the first 

indications that this material was not juvenile 2004 material 

but was excavated from the 1980s dome (see Pallister and others, this volume, chap. 30).

All of the dome samples that were erupted and analyzed 

after sample SH301-1A7 represented new juvenile magma, 

and multiple analyses of leached samples produced radioactive equilibration curves that generally demonstrated complete 

degassing of 210Po before eruption (fig. 1). Plagioclase crystals 

in sample SH304-2A had initial (210Po) and (210Pb) values of 

about 0.29 dpm/g (table 1). Assuming that plagioclase in all 

of the 2004–5 samples have this same (210Po), that plagioclase 

crystals make up 50 percent of whole rock, and that other 

coarsely crystalline phases have near zero (210Po), samples 

with groundmasses that are entirely degassed of 210Po would 

have (210Po)≈ 0.14 dpm/g. Therefore, the “no-Po” line on figure 

1 is used to represent the approximate value of a whole rock 

whose groundmass had been entirely degassed of 210Po.

Representative values of (210Po) for whole rocks are 

compared with the values for leached samples in figure 2. In 

all cases, the whole-rock values exceed the leached values. 

For most samples, the values for unleached samples analyzed 

within two months of eruption are between 5 and 20 percent 

higher than for the leached samples. For sample SH305-1, the 

unleached values are more than 50 percent higher; for sample 

SH316-1A, these values are highly variable and as much as 

several times higher. These data indicate that a significant 

amount of 210Po resides in open pore spaces, which indicates 

that a fraction of the 210Po in all samples degassed into bubbles 

that never escaped from the magmas. This observation is 

important from an analytical point of view because it demonstrates the importance of leaching samples in dilute acid 

to obtain intrinsic 210Po activities in lavas after all gas is lost 

from the magma. It also is important because it can allow the 

change from open- to closed-system degassing to be constrained temporally.

Nearly all dome whole-rock samples analyzed had initial 

(

210Po)/(210Pb) ratios near the no-Po value when they vented. 

That is, the last day of complete degassing of 210Po from the 

Mount St. Helens dacite, which is represented by the date of 

intersection of the 210Po ingrowth curves with the no-Po line 

on figure 1, generally corresponds with their day of eruption 

within error of analysis and knowledge of eruption day (fig. 

3). An exception to this is sample SH305-1, which erupted in 

mid-November 2004 but apparently had last degassed Po in 

September 2004.

Like the dome samples, the lithified gouge sample 

SH307-2A degassed 210Po until the time that it vented to the 

surface. In contrast, the nonlithified gouge sample SH314-1G 

either ceased degassing 210Po about a month before it erupted, 

or gas derived from underlying magma added 210Po to this 

sample (fig. 3).

All of the calculated initial (210Pb)/(226Ra) values for 

whole rocks from the ongoing eruption are within 2σ analytical error of equilibrium (table 1). This is in marked contrast 

with samples erupted from Mount St. Helens in 1980, which 

had widely varying (210Po)/(226Ra) values (0.8 to 1.4) when 

they erupted (Berlo and others, 2004).

Discussion

210Po-210Pb Disequilibrium

The tephra sample ejected near the beginning of the 

ongoing eruption on October 4, 2004, was strongly enriched in 

210Po over 210Pb, demonstrating that magmatic Po had condensed on surfaces in the pore space of the shallow fumarolic 

system shortly before eruption and had erupted along with the 

pulverized rock fragments. In contrast, the hand-picked tephra 

fragments erupted on March 8, 2005, during active dome 

building had low initial (210Po) values, indicating that these 

fragments were mostly or entirely pieces of shallow dome 

material that had been emplaced about two weeks before their 

involvement in the phreatomagmatic explosions (fig. 3).
850 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

Sample No. Sample 

type

Eruption

day

Analytical 

treatment

Analysis 

day

(

210Po)

dpm/g ± 1σ (

210Pb) (210Pb)/(226Ra) ± 1σ

MSH04E2A03_A2 Tephra 10/4/04 whole 10 2.361 0.052 ~1.17

whole 99 1.927 0.044

whole 134 1.806 0.041

SH301-1A7 Dome 1980s whole 30 1.191 0.026 1.04

whole 150 1.131 0.026

whole 559 1.043 0.025

SH304-2A Dome 10/18/04 whole 42 0.359 0.018 1.19 0.958 0.027

whole 99 0.574 0.019

whole 126 0.642 0.030

whole 298 0.966 0.025

leached 126 0.593 0.021

leached 298 0.949 0.024

leached 369 1.015 0.024

SH304-2A

Plagioclase 10/18/04 leached 66 0.298 0.010 0.29 4.4 0.1

leached 188 0.285 0.011

SH305-1 Dome 11/20/04 whole 99 0.907 0.024 1.17 1.023 0.039

whole 112 0.845 0.040

whole 253 1.036 0.076

leached 112 0.621 0.021

leached 218 0.863 0.021

leached 253 0.913 0.025

leached 475 1.061 0.023

leached 475 1.084 0.024

SH306 Dome 12/15/04 whole 126 0.384 0.014 1.15 0.974 0.031

whole 141 0.486 0.015

whole 211 0.708 0.030

leached 141 0.445 0.014

leached 211 0.681 0.021

leached 396 0.938 0.021

SH307-2A Gouge 2/12/05 whole 159 0.288 0.030 1.17 1.038 0.033

whole 218 0.487 0.023

leached 159 0.251 0.010

leached 218 0.464 0.015

leached 475 0.980 0.024

SH311-1B Ballistic 

fragment

1/16/05 whole 159 0.467 0.030 1.15 1.056 0.033

leached 159 0.438 0.013

leached 216 0.611 0.019

leached 475 1.001 0.023

MSH05DRS_3_9_4 Handpicked 

clasts

3/8/05 whole 202 0.405 0.028 1.15 1.018 0.032

leached 202 0.394 0.011

leached 483 0.982 0.025

SH314-1G Gouge 4/17/05 whole 211 0.360 0.043 1.13

leached 211 0.311 0.012

leached 396 0.810 0.020

SH315-1 Dome 4/1/05 whole 211 0.370 0.070 1.13

leached 211 0.323 0.013

leached 396 0.812 0.020

leached 965 1.116 0.022

Table 1. Activities of 210Po and initial (210Pb)/(226Ra) values for samples erupted in 2004–2005 from Mount St. Helens, 

Washington.

[Eruption day (month/day/year) from Pallister and others (this volume, chap. 30). Day of analysis is number of days after October 1, 2004. Initial 

(

210Pb) values calculated by regressing a best-fit radioactive equilibration curve through leached sample analyses. Values for (226Ra) in dpm/g (disintegrations per minute per gram) from Cooper and Donnelly (this volume, chap. 36).]
37. Timing of Degassing and Plagioclase Growth in Lavas Erupted from Mount St. Helens, 2004–2005 851

Table 1. Activities of 210Po and initial (210Pb)/(226Ra) values for samples erupted in 2004–2005 from Mount St. Helens, 

Washington.—Continued

[Eruption day (month/day/year) from Pallister and others (this volume, chap. 30). Day of analysis is number of days after October 1, 2004. Initial 

(

210Pb) values calculated by regressing a best-fit radioactive equilibration curve through leached sample analyses. Values for (226Ra) in dpm/g (disintegrations per minute per gram) from Cooper and Donnelly (this volume, chap. 36).]

SH316-1A Dome 4/15/05 whole 253 3.642 0.121 1.14 1.028 0.033

whole 465 1.617 0.076

leached 334 0.671 0.018

leached 559 0.968 0.024

leached 965 1.143 0.023

SH317-1A Dome 5/1/05 whole 290 0.543 0.019 1.11 0.973 0.031

whole 465 1.005 0.068

leached 298 0.459 0.014

leached 483 0.846 0.022

leached 965 1.087 0.020

MSH319-1 Dome 5/15/05 whole 298 0.534 0.013 1.14

whole 465 0.818 0.042

leached 298 0.428 0.011

leached 516 0.895 0.022

leached 978 1.122 0.022

Sample No. Sample 

type

Eruption

day

Analytical 

treatment

Analysis 

day

(

210Po)

dpm/g ± 1σ (

210Pb) (210Pb)/(226Ra) ± 1σ

Although initial whole-rock 210Po activities are variable 

for dacite lavas from Mount St. Helens, these activities drop 

to near zero in nearly all samples when 210Po is removed from 

surfaces by leaching with dilute acid (table 1). Thus, like 

most other lavas (Gill and others, 1985; Lambert and others, 

1985; Le Cloarec and Pennisi, 2001; Bennett and others, 1982; 

Reagan and others, 2006), the intrinsic initial 210Po activities 

of Mount St. Helens dacites are near zero. The presence of 

significant 210Po in pore spaces of the Mount St. Helens dacites 

implies that the magmas continued to partition 210Po into the 

gas phase until the magmas arrived at the vent and cooled 

below the blocking temperature of 210Po.

The activities of 210Po in whole dacite samples were 

significantly more variable and higher on the day the samples 

vented to the surface compared with the activities of 210Po in 

leached samples. These differences probably reflect variations 

in the depth at which magmas ceased being able to vent gas 

to their surroundings; that is, when degassing changed from 

being open-system to closed-system. For most samples, the 

initial quantity of 210Po in pore spaces was small, which is 

consistent with closed-system degassing beginning no more 

than 1–2 weeks before eruption (fig. 2). At an eruption rate 

of 5–10 meters per day, which is the approximate ascent rate 

of magmas in the shallow conduit (LaHusen and others, this 

volume, chap. 16), this switch from open- to closed-system 

degassing must have typically occurred between about 35 and 

150 meters depth, which is similar to the depth where magmas 

are becoming rheological solids (Moran and others, this volume, chap. 2). However, sample SH305-1, which is unaltered 

and significantly more vesicular and more glassy than other 

samples, had much higher leachable 210Po values than less 

vesicular samples. This sample also has the greatest difference between its eruption day and last degassing day (fig. 3). 

Therefore, assuming that this sample’s velocity was the same 

as the velocity for other dome dacites, sample SH305-1 might 

have begun closed-system degassing as much as 200 days 

before eruption (corresponding to 1–2 km depth) and ceased 

degassing altogether about 2 months before it erupted (corresponding to about 0.5 km depth). If these assumptions are 

correct, this dacite chilled deeper within the edifice than other 

dacites from the ongoing eruption, which also might explain 

why sample SH305-1 contains fewer groundmass crystals than 

other samples from the 2004–5 eruption (see Cashman and 

others, this volume, chap. 19).

One sample with exceptionally high leachable 210Po was 

sample SH316-1A, which had a whole rock (210Po)/(210Pb) 

value significantly greater than unity when it erupted. This 

sample had significant orange staining on fracture surfaces; 

although these surfaces were avoided during hand picking 

before analysis, extraneous 210Po from underlying magmas 

appears to have been deposited heterogeneously in pore spaces 

throughout this sample. Nevertheless, this extraneous Po did 

not diffuse into the magma or solidified dacite, because the 

eruption and Po-degassing dates for sample SH316-1A are 

within error of each other.

Like the whole-dome samples, the lithified gouge sample 

SH307-2A degassed 210Po approximately until the day that 

it vented. Thus, despite its position along the conduit walls, 

where it quenched to a solid and was ground to gouge, this 

sample maintained a temperature that was high enough for 

210Po to degas completely until it vented. The loose gouge 

sample SH314-1G appears to have cooled below the 210Po 
852 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

0.0

0.5

1.0

1.5

2.0

2.5

0 100 200 300 400 500 600

DAYS SINCE OCTOBER 1, 2004

(210Po), IN DISINTEGRATIONS PER MINUTE PER GRAM (dpm/g)

Oct. 4, 2004, tephra

SH301-1A7

SH304-2A

SH305-1

SH306

SH307-2A, gouge

SH311-1B, ballistic fragment

MSH05DRS_3_9_4

SH314, 315-1

SH316-1A

SH317-1A

SH319-1

Dome samples unless indicated otherwise

EXPLANATION

no-Po

Figure 1. Plot of 210Po activities for leached samples recently erupted from Mount St. Helens, Washington. 210Po 

ingrowth curves drawn by choosing the day when (210Po) = 0 and the initial (210Pb) value that produced best-fit 

line through data points for each sample. The “no Po” line represents the approximate whole-rock value when 

groundmass was entirely degassed of 210Po. Analyses completed more than 600 days after October 1, 2004, are not 

shown, although they were used to determine the best-fit ingrowth curves. See text for further explanation.

blocking temperature about a month before eruption, which 

corresponds to a depth of about 150–300 m below the surface.

210Pb-226Ra Equilibrium

As demonstrated above, the dacites from the ongoing eruption of Mount St. Helens reach the surface nearly 

degassed of 210Po. Nevertheless, all lavas from the ongoing 

eruption had near-equilibrium (210Pb)/(226Ra) values when they 

erupted. Both because 222Rn is a noble gas and because 210Po 

appears to partition entirely into the gas phase at magmatic 

temperatures at Mount St. Helens, we assume that 222Rn also 

strongly partitions from magma into a separate gas phase. If 

so, then the lack of a measurable 210Pb deficit indicates that 

open-system degassing of dacites erupting at Mount St. Helens 

must be confined to the time period between about 1–2 years 

(the minimum time to produce a measurable 210Pb deficit) 

before eruption and the start of closed-system degassing near 

the surface. These equilibrium values also indicate that the 

lavas from the ongoing eruption did not experience persistent 

streaming of 222Rn from large volumes of underlying magma, 

which contrasts with lavas and tephras erupted from Mount 

St. Helens in 1980 (compare Berlo and others, 2004). Finally, 

these equilibrium (210Pb)/(226Ra) values indicate that the 

melting processes that created 226Ra excesses over 230Th (and 

probably 210Pb) in all of the dacites from the ongoing eruption 

(Cooper and Donnelly, this volume, chap. 36), as well as the 

majority of plagioclase fractionation, occurred more than a 

century before eruption.

The contrast between the lack of (210Pb)/(226Ra) disequilibria during this eruption and the significant disequilibria 

encountered in samples from the eruption in 1980 (Berlo and 

others, 2004) is noteworthy. This difference is in stark contrast to Li concentration data, which indicate that plagioclase 

is enriched in Li at the beginning of both eruptions (compare 

Berlo and others, 2004; Kent and others, 2007) compared 

with subsequently erupted samples. The Li data suggest that 

a Li-enriched fluid pervasively infiltrated the first-erupted 

magmas and raised Li concentrations in plagioclase through 
37. Timing of Degassing and Plagioclase Growth in Lavas Erupted from Mount St. Helens, 2004–2005 853

0.2

0

0.4

0.6

0.8

1.0

1.2

0 50 100 150 200 250 300

(210Po), IN DISINTEGRATIONS PER MINUTE PER GRAM (dpm/g)

DAYS SINCE OCTOBER 1, 2004

SH30

S

7

H306

SH305-1

SH304-2A

Open symbols, unleached samples

Filled symbols, leached samples

25

0

-25

75

125

175

225

-25 250 75 125 175 225

ERUPTION DAY (0 = OCTOBER 1, 2004)

LAST 210Po DEGASSING DAY

Dome sample

EXPLANATION

Gouge

Tephra

- 30

- 60

- 90

SH304-2A

SH307-2A SH314-1G

SH305-1

MSH05DRS_3_9_4

Figure 2. Plot of 210Po activity versus date of analysis in days after Oct. 1, 2004, for 

representative whole-rock powders from Mount St. Helens, Washington, before (open 

symbols) and after (filled symbols) leaching (symbols match those in figure 1).

Figure 3. Plot of the last day 

of 210Po degassing (date of 

intersection of 210Po ingrowth 

curves with no-Po line on figure 

1) versus the date of eruption for 

samples from the ongoing eruption 

of Mount St. Helens, Washington. 

Numbers for samples mentioned in 

the text are shown. Symbols show 

best guess for a sample’s eruption 

date, and lines indicate the range of 

potential eruption dates. The solid 

diagonal line marks equivalent last 

210Po degassing day and eruption 

day (210Po ceases degassing on 

eruption day). Dashed lines show 

calculated effect if 210Po ceased 

degassing 30, 60, or 90 days before 

eruption. Note that most samples of 

all types degas 210Po by closed- or 

open-system processes until the 

magmas are within two weeks of 

arriving at the surface.
854 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

diffusion, whereas this fluid either did not infiltrate subsequent magmas for both eruptions, or it did so for too short 

a time to affect Li concentrations in plagioclase (Kent and 

others, 2007). The 210Pb enrichment found in the cryptodome 

and pumice samples erupted in 1980 was attributed to a Rn 

flux accompanying the Li flux (Berlo and others, 2004). If 

so, and if the 222Rn excess required to ingrow 210Pb also was 

associated with Li-bearing fluid for the current eruption, then 

this fluxing must have occurred during a much shorter time 

period before the current eruption than was the case in 1980 

(Kent and others, 2007). An alternative possibility is that the 

222Rn flux required for the high (210Pb)/(226Ra) values in the 

1980 magmas was separate from the Li-bearing fluid. In that 

case, the 222Rn flux may have been associated with the orderof-magnitude higher rates of gas emission in 1980 compared 

with emission rates of the ongoing eruption (Gerlach and 

others, this volume, chap. 26).

The plagioclase mineral separate from sample SH304-2A 

has (210Po)/(210Pb)≈1, indicating that very little of this plagioclase grew within 1–2 years before eruption. However, comparing the (210Pb) value with the (226Ra) value reported in Cooper 

and Donnelly (this volume, chap. 36) gives a (210Pb)/(226Ra) 

value of about 4.4, indicating that a significant proportion of 

the plagioclase crystals in this lava grew within several decades 

of eruption. Assuming t = 850°C and an average plagioclase 

An value of 0.5, the DPb/DRa value in plagioclase would have 

been about 9.3 based on the lattice strain method of Blundy 

and Wood (1994) to calculate DRa and the partitioning model 

of Bindeman and others (1998) to calculate DPb. Reconciling 

the (210Pb)/(226Ra) values and DPb/DRa values produces a model 

average age of the plagioclase in the mineral separate of about 

30 years. Cooper and Donnelly (this volume, chap. 36) show 

that this plagioclase mineral separate has a (226Ra)/(230Th) value 

similar to that of the whole rock but a Th isotope ratio that is 

less than that of the whole rock by about 25 percent. These data 

were interpreted to suggest that 60–80 percent of the plagioclase is “zero-age” and the rest is older than 10 ka. Assuming 

that 70 percent of the plagioclase is young and that it grew in a 

single pulse, the (210Pb)/(226Ra) ratio in the newly grown material 

would have been about 5.9, which gives a model age of about 

17 years, consistent with a large proportion of the plagioclase 

growing at about the time of the 1980–86 eruption.

Conclusions

1. Leaching of powdered whole-rock samples by dilute 

acid to remove 210Po deposited in pore spaces is necessary 

to allow measurement of magmatic 210Po values.

2. 210Po degasses from lavas associated with the ongoing eruption of Mount St. Helens until they vent at the 

surface. The change from open- to closed-system degassing generally occurs within 150 m of the surface, which 

is about the depth at which the exterior dome samples 

become solid. Nevertheless, closed-system degassing 

began as deep as 1–2 km for some magmas.

3. Our sample of indurated gouge, collected from the 

surface of the lava dome, continued to degas 210Po until it 

reached the surface, whereas our sample of loose, more 

interior gouge appears to have ceased degassing 210Po at a 

depth of 150–300 m.

4. Magmas erupted in 2004–5 from Mount St. Helens 

were generated and underwent differentiation more than 

a century before eruption on the basis of equilibrium 

(210Pb)/(226Ra) values, but less than a few thousand years 

before eruption on the basis of disequilibrium (226Ra)/

(230Th) values (see Cooper and Donnelly, this volume, 

chap. 36).

5. The near-equilibrium (210Pb)/(226Ra) ratios for all of the 

dacites erupted in 2004–5 from Mount St. Helens suggest 

that the last stage of open-system degassing of these magmas was confined to the time period between about a year 

before eruption and the start of closed-system degassing 

near the surface. These values further indicate that persistent streaming of excess 222Rn gas through the magmas 

did not occur before this eruption, despite Li enrichments 

observed in plagioclase erupted in October and November 

2004 (Kent and others, 2007).

6. The combined dataset of short- and long-lived 

U-series nuclides suggests that a significant proportion 

of the coarse plagioclase in a sample erupted in November 2004 grew at about the time of the 1980s eruptions 

of Mount St. Helens, and not during this last stage of 

magma rise and degassing.

Acknowledgments

We thank the scientists at the U.S. Geological Survey’s Cascades Volcano Observatory and other members of 

the Mount St. Helens working group for sharing data and 

information about the ongoing eruption of Mount St. Helens. 

David Peate, Jennifer Garrison, and Kim Berlo are thanked 

for insightful discussions. This paper benefited greatly from 

constructive reviews by Pierre Gauthier and Todd Hinkley. 

National Science Foundation grants EAR 0504362 and EAR 

0609670 funded this research.
37. Timing of Degassing and Plagioclase Growth in Lavas Erupted from Mount St. Helens, 2004–2005 855

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128 A Volcano Rekindled: The Renewed Eruption of Mount St. Helens, 2004−2006

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Late September and early October 2004 saw intense 

seismic unrest, rapid ground deformation, and explosions 

heralding the renewal of lava-dome growth in the crater 

of Mount St. Helens, Washington. Thus ended the relative 

calm of two decades since the renowned 1980–86 eruptions. 

. . . Occurring more than a quarter century ago, the 

catastrophic 1980 eruption of Mount St. Helens may seem 

to some as just a historical footnote. This Professional Paper

documenting both the volcano’s recent activity and the 

need for coordinated warning and disaster response should 

serve the public well as a reminder of the challenges and 

consequences of living near an active volcano.

 From the Foreword by Mark D. Myers

Printed on recycled paper