U.S. Geological Survey Open-File Report 02–328

Geological Survey of Canada Open File 4350

August 2002

THE CASCADIA SUBDUCTION ZONE AND

RELATED SUBDUCTION SYSTEMS

Seismic Structure,

Intraslab Earthquakes and Processes,

and Earthquake Hazards

Editors: Stephen Kirby, Kelin Wang and Susan Dunlop

The Cascadia Subduction Zone and

Related Subduction Systems—

Seismic Structure, Intraslab Earthquakes and

Processes, and Earthquake Hazards

Edited by Stephen Kirby1, Kelin Wang2, and Susan Dunlop3

Released simultaneously as

U.S. GEOLOGICAL SURVEY OPEN-FILE REPORT 02–328 and

GEOLOGICAL SURVEY OF CANADA OPEN FILE 4350

Version 1.0

2002

This report is preliminary and has not been reviewed for conformity with U.S. Geological Survey editorial standards or with

the North American Stratigraphic Code. Any use of trade, firm, or product names is for descriptive purposes only and does

not imply endorsement by the U.S. Government. Although all data and software released on this CD-ROM have been used

by the USGS, no warranty, expressed or implied, is made by the USGS as to the accuracy of the data and related material

and (or) the functioning of the software.

U.S. DEPARTMENT OF THE INTERIOR

U.S. GEOLOGICAL SURVEY

1 U.S. Geological Survey MS 977, 345 Middlefield Road, Menlo Park, California, 94025, USA

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

3 Centre for Earth and Ocean Research, University of Victoria, P.O. Box 3055, STN CSC, Victoria, BC, V8W 3P6, Canada
ISBN: 0-607-99675-7

This publication and any updates to it can be found online at:

http://geopubs.wr.usgs.gov/open-file/of02-328/
The Cascadia Subduction Zone and Related Subduction Systems | iii

Preface

The following report is the principal product of an international workshop titled “Intraslab Earthquakes in the Cascadia

Subduction System: Science and Hazards” and was sponsored by the U.S. Geological Survey, the Geological Survey

of Canada and the University of Victoria. This meeting was held at the University of Victoria’s Dunsmuir Lodge,

Vancouver Island, British Columbia, Canada on September 18–21, 2000 and brought 46 participants from the U.S.,

Canada, Latin America and Japan. This gathering was organized to bring together active research investigators in the

science of subduction and intraslab earthquake hazards. Special emphasis was given to “warm-slab” subduction

systems, i.e., those systems involving young oceanic lithosphere subducting at moderate to slow rates, such as the

Cascadia system in the U.S. and Canada, and the Nankai system in Japan. All the speakers and poster presenters

provided abstracts of their presentations that were a made available in an abstract volume at the workshop. Most of

the authors subsequently provided full articles or extended abstracts for this volume on the topics that they discussed

at the workshop. Where updated versions were not provided, the original workshop abstracts have been included.

By organizing this workshop and assembling this volume, our aim is to provide a global perspective on the science of

warm-slab subduction, to thereby advance our understanding of internal slab processes and to use this understanding to improve appraisals of the hazards associated with large intraslab earthquakes in the Cascadia system. These

events have been the most frequent and damaging earthquakes in western Washington State over the last century. As

if to underscore this fact, just six months after this workshop was held, the magnitude 6.8 Nisqually earthquake

occurred on February 28th, 2001 at a depth of about 55 km in the Juan de Fuca slab beneath the southern Puget

Sound region of western Washington. The Governor’s Office of the State of Washington estimated damage at more

than US$2 billion, making it among the costliest earthquakes in U.S. history.
iv |
The Cascadia Subduction Zone and Related Subduction Systems | v

Table of Contents

SUMMARY AND INTRODUCTION

Workshop consensus findings and recommendations

Stephen H. Kirby, Kelin Wang, Workshop Steering Committee (Garry C. Rogers, Robert S. Crosson, Anne M.

Tréhu, Chris Goldfinger and Tom M. Brocher) and the workshop participants ......................................................................................1

Introduction to a global systems approach to Cascadia slab processes and associated earthquake hazards

Stephen H. Kirby and Kelin Wang ..................................................................................................................................................................................................5

CASCADIA SUBDUCTION SYSTEM: SETTING AND STRUCTURE

The Juan de Fuca plate and slab: Isochron structure and Cenozoic plate motions

Douglas S. Wilson ........................................................................................................................................................................................................................................9

Offshore structure of the Juan de Fuca plate from marine seismic and sonar studies

Chris Goldfinger, Robert P. Dziak and Chris Fox ............................................................................................................................................................. 13

Cascadia microplate models and within-slab earthquakes

Ray E. Wells, Richard J. Blakely and Craig S. Weaver................................................................................................................................................. 17

Geometry of the subducting Juan de Fuca plate: New constraints from SHIPS98

Anne M. Tréhu, Tom M. Brocher, Kenneth C. Creager, Mike A. Fisher, Leiph A. Preston, George Spence

and the SHIPS98 Working Group ............................................................................................................................................................................................... 25

Three-dimensional structure of the Cascadia forearc region from SHIPS active experiment

and earthquake observations: Tomographic inversion provides a high-resolution view into

the core of the Cascadia forearc complex

Robert S. Crosson, Neil P. Symons, Kenneth C. Creager, Leiph A. Preston, T. Van Wagoner, Tom M. Brocher,

Mike A. Fisher and the SHIPS98 Working Group .......................................................................................................................................................... 33

Velocity structure of western Washington from local earthquake tomography

Antonio Villaseñor, D. Stanley and Harley Benz ............................................................................................................................................................ 35

Three-dimensional reflection image of the subducting Juan de Fuca plate

Kenneth C. Creager, Leiph A. Preston, Robert L. Crosson, T. van Wagoner, Anne Tréhu

and the SHIPS98 Working Group ............................................................................................................................................................................................... 37

Modelling post-critical Moho reflections from intraslab earthquakes in the Cascadia subduction zone

Jeff Shragge, Michael Bostock and Garry C. Rogers.................................................................................................................................................... 43

Geometry and membrane deformation rate of the subducting Cascadia slab

Ling-Yun Chiao and Kenneth C. Creager............................................................................................................................................................................... 47

Crustal structure of northern Juan de Fuca plate and Cascadia subduction zone—new results, old data

Ron M. Clowes ........................................................................................................................................................................................................................................... 55

Seismic reflection imaging of the Cascadia plate boundary offshore Vancouver Island

Andrew Calvert........................................................................................................................................................................................................................................... 59

An image of the Cascadia subduction zone across central Oregon

Stephane Rondenay and Michael G. Bostock .................................................................................................................................................................. 63

CASCADIA SLAB EARTHQUAKES

Intraslab earthquakes beneath Georgia Strait/Puget Sound

Garry C. Rogers and Robert S. Crosson ................................................................................................................................................................................ 65
vi |

Precise relocations of slab seismicity in the northern Cascadia subduction zone

John F. Cassidy and Felix Waldhauser...................................................................................................................................................................................... 69

Large earthquakes under our feet and deep beneath the Puget Lowland and Georgia Strait: How often do they

occur, how big can they get and what are their likely effects?

Stephen H. Kirby, Kenneth C. Creager and Leif A. Preston ................................................................................................................................... 75

In-slab seismicity at the north end of Cascadia

Garry C. Rogers, Alison L. Bird, John F. Cassidy, Taimi L Mulder, Maiclaire K. Bolton and John P. Ristau ..................... 77

INTRASLAB EARTHQUAKES IN WARM SLAB SETTINGS WORLDWIDE

Warm-slab subduction as a global process

Stephen H. Kirby, E. Robert Engdahl and Antonio Villaseñor.............................................................................................................................. 79

The structure and intraslab seismicity of the Philippine Sea plate in Southwest Japan

Yoshiyuki Kaneda, Phil R. Cummins and Kouichi Uhira .......................................................................................................................................... 81

Intraslab earthquakes in the subducting oceanic plates below Mexico

Shri K. Singh, Vladimir Kostoglodov and J.F. Pacheco................................................................................................................................................ 87

Seismicity and seismotectonics of Central America with emphasis on intraslab seismic activity

Marino Protti ................................................................................................................................................................................................................................................. 93

Large, intermediate-depth earthquakes in Southwest Japan

Phil R. Cummins, Yoshiyuki Kaneda and Takane Hori .............................................................................................................................................. 99

Large (Mw≥ 5.5) intraplate slab earthquakes (1928–2000) of the Prince William Sound region, Alaska

Diane Doser, Monique Velasquez, Wesley Brown and Annette Veilleux ............................................................................................... 103

INTRASLAB AND INTRAPLATE RADIATED EARTHQUAKE ENERGIES WORLDWIDE

The radiated seismic energy and apparent stress of interplate and intraslab earthquakes at subduction-zone

environments: Implications for seismic hazard estimation

George L. Choy, John L. Boatwright and Stephen H. Kirby ............................................................................................................................... 107

Earthquakes having high apparent stress in oceanic intraplate lithosphere

Art McGarr and George L. Choy .............................................................................................................................................................................................. 115

Energy–to–moment ratios for damaging intraslab earthquakes: Preliminary results

Emile Okal and Stephen H. Kirby ............................................................................................................................................................................................ 117

MODELS

Thermal structure and metamorphism of subducting oceanic crust: Insight into Cascadia

intraslab earthquakes

Simon M. Peacock, Kelin Wang and Aaron M. McMahon ................................................................................................................................. 123

Heat sources in the subduction zones: Implications for slab seismicity and arc volcanism

Jason R. McKenna and David D. Blackwell ..................................................................................................................................................................... 127

Theoretical mineralogy, density, seismic wavespeeds and H2O content of the Cascadia subduction zone,

with implications for intermediate-depth seismicity and earthquake hazard

Bradley R. Hacker, Geoffrey A. Abers and Simon M. Peacock ........................................................................................................................ 133

Slab deformation at the edge of subduction zones: An example from Cascadia

Uri S. ten Brink ........................................................................................................................................................................................................................................ 139

Stresses in the Juan de Fuca plate and the role of mantle resistance to horizontal slab motion

Kelin Wang, Jiangheng He, Earl E. Davis and Chris Goldfinger ...................................................................................................................... 143
The Cascadia Subduction Zone and Related Subduction Systems | vii

Phase changes, fluids and the co-location of the deep and shallow seismicity beneath Puget Sound

and southern Georgia Strait

Garry C. Rogers....................................................................................................................................................................................................................................... 145

The role of phase changes in the development of forearc basins

Garry C. Rogers....................................................................................................................................................................................................................................... 147

GROUND MOTION PREDICTION AND HAZARD APPRAISAL

Knowledge of in-slab earthquakes needed to improve seismic hazard estimates for southwestern

British Columbia

John Adams and Stephen Halchuk ........................................................................................................................................................................................ 149

Empirical ground motion relations for subduction zone earthquakes

Gail M. Atkinson and David M. Boore ................................................................................................................................................................................ 155

Deep intraslab earthquakes and how they contribute to seismic hazard in the Pacific northwestern U.S.

Mark D. Petersen and Arthur D. Frankel ........................................................................................................................................................................... 159

RELEVANT INITIATIVES

The proposed NEPTUNE seafloor observatory on the Juan de Fuca plate: Applications to studies

of intraplate earthquakes

William S.D. Wilcock , Kelin Wang , Roy D. Hyndman , Anne M. Tréhu and Kenneth C. Creager .............................. 163

POLARIS: New opportunities in Cascadia

John F. Cassidy, Michael G. Bostock, Gail Atkinson, David W. Eaton, Isa Asudeh, David B. Snyder,

John Adams, Ian J. Ferguson, Alan G. Jones and Colin J. Thomson ........................................................................................................... 169
viii |
The Cascadia Subduction Zone and Related Subduction Systems | 1

Workshop consensus findings and recommendations

Stephen H. Kirby1, Kelin Wang2, Workshop Steering Committee (Garry C. Rogers2, Robert S. Crosson3,

Anne M. Tréhu4, Chris Goldfinger4 and Tom M. Brocher1) and the workshop participants

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

3 Geophysics Program, University of Washington, PO Box 351650, Seattle, Washington, 98195–1650, USA

4 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97331–5503, USA

skirby@usgs.gov, wang@pgc.nrcan.gc.ca, rogers@pgc.nrcan.gc.ca, crosson@u.washington.edu,

trehu@oce.orst.edu, gold@oce.orst.edu, brocher@usgs.gov

Intraslab earthquake science for Cascadia

and related subduction zones:

what can we presently agree upon?

Building upon the findings of the oral and poster presentations and the related discussions, we held a plenary

session to develop a consensus view of what we could

agree upon as to the primary characteristics of warmslab intraslab earthquakes and what interpretive models

are both consistent with these characteristics and suggestive of productive research directions toward understanding their origin. This section summarizes those

discussions.

1. Large intraslab earthquakes, by virtue of their frequency of occurrence, locations directly beneath

population centers and source characteristics, represent a major earthquake hazard in the Cascadia subduction zone.

2. There are two regions of observed deep seismicity in

the Cascadia system: one just north of the Mendocino

triple junction and another under the Puget lowlands/

Georgia Strait. This heterogeneous distribution probably reflects the stresses and seismic-strain-rate localization connected with the northward collision of the

Pacific plate along the Mendocino fracture zone and

the stress effects of the geometrical arch in the northern region, respectively. The absence of large, deep

intraslab earthquakes beneath western Oregon is

probably not due to inadequate sampling during the

instrumental seismic era, at least since the late 1950s.

3. Many of the fundamental characteristics of the Cascadia

subduction system are shared by numerous other

“warm” subduction systems involving young lithosphere

descending at slow rates. These characteristics include

usually shallow earthquake hypocentral depth (<100

+ 25 km), generally low average dip of the seismic

zone (<20°), focal mechanisms with in-slab T axes, and

feeble or absent arc volcanism. Trenches in these settings are also generally muted in relief and often have

neutral or reverse curvature.

4. Investigations to date of shallow intraslab earthquake

sources worldwide show that they tend to be enriched

in high-frequency energy compared to nearby interplate thrust earthquakes of comparable scalar seismic

moment. Subduction in some particular warm-slab

environments, such as Mexico, Peru and Chile has

produced many large (M>7.0–8.0) and very destructive slab earthquakes probably in the subducted slab

mantle. These events are sometimes exceptionally

enriched in high-frequency energy compared to

nearby interplate earthquakes, a factor that contributes to their destructiveness. These unusual regions

may involve downdip changes in the sign of slab curvature (implied by their flat-slab geometry that occurs in some of those localities). The normal-faulting

focal mechanisms of these events are consistent with

reverse-curvature flexure in the slab mantle but the

physical mechanisms for faulting in such settings are

not known.

5. Aftershocks of the largest earthquakes in this class are

usually smaller and fewer, and are not known to be

damaging. This contrasts with both shallow interplate

and forearc earthquakes, and argues for significant

differences in the mechanics of faulting.

6. Seismological observations in Cascadia, Nankai

(Southwest Japan), the Chilean Andes and the Aleutians indicate that small to intermediate-sized intraslab

earthquakes largely occur in a thin low-wavespeed

layer identified as subducting crust.

7. Thermo-mineralogical models for warm-slab subduction predict that the metamorphic reactions that lead

to dehydration and eclogite formation should occur

at depths shallower than 100 km in the oceanic crust.

This shallow depth distribution of intraslab earthquakes in many warm slabs suggests that they also

reflect the shallow metamorphism of subducted crust

predicted by thermo-mineralogical models.

8. Laboratory studies indicate that release of water by
2 | Kirby et al.: Workshop consensus findings and recommendations

dehydration of hydrous minerals under stress can permit brittle fracture by faulting, even at pressures corresponding to depths greater than a few tens of

kilometers, by elevating pore pressure and reducing

the frictional resistance to fault sliding, a process

termed dehydration embrittlement. Such a process

acting inside faults hydrated by water could also permit reactivation of such faults during the heating

accompanying slab descent.

9. Simplified theoretical models predict that densification

to eclogite of the slab crust by metamorphism should

produce internal slab stresses that may be important

in driving seismic deformation of the slab. These predicted deviatoric stresses also vary in sign across the

subducting crust–mantle boundary and hence can

cause sharp spatial variations in focal mechanisms, as

is sometimes observed. This phase-change source of

slab stress also should act in combination with those

associated with regional plate-scale forces, membrane

deformation associated with trenches having neutral

or reverse curvature, and with slab buoyancy forces.

Summary recommendations of research

directions for intraslab earthquake science

and hazard appraisal

These recommendations were arrived at based on discussions in three breakout groups and subsequently refined in plenary session and by the Workshop Steering

Committee.

Recommendation 1

Largely marine investigations of the incoming structure

of the Juan de Fuca plate with normal seafloor spreading

and regions of sheared oceanic lithosphere as described

by Wilson [this volume].

• Thermal structure: heat flow.

• Seismic wavespeed structure: seismic refraction.

• Identify faults and active faulting in subducting Juan

de Fuca plate: ocean bottom seismometer deployments for detecting near-trench earthquakes. Seismic

reflection and sonar surveys.

• Improve knowledge of slab motions: refine the Juan

de Fuca plate/Oregon Coast microplate and other

North America forearc microplate motions using GPS.

Integrate with longer-term geophysical and geologic

evidence to refine the earlier Cenozoic slab motions.

Possible payoffs include:

• Improved thermal/mineralogical models of slab evolution with descent.

• Identify possible future sources of intraslab earthquakes that can occur by reactivation during dehydration (fault dimensions and orientations).

Recommendation 2

Relocations of intraslab earthquakes using double difference, waveform correlation, three-dimensional wavespeed models and modern teleseismic methods.

• Integrate historical analogue and modern digital

teleseismic data using consistent wavespeed models.

• Integrate all analog and digital Pacific Northwest network data with recent three-dimensional wavespeed

models.

• Reconcile hypocenters from regional network and

teleseismic station data.

Possible payoffs:

• Identify possible subsurface faults and hence fixed

earthquake sources.

• Establish subsurface fault dimensions and hence place

possible limits on the maximum scalar moments and

magnitudes of earthquakes along them.

• Refine slab shape and, with slab motions, refine spatial rates of slab flexural deformation and compare with

distribution of slab seismicity moment release rates.

• Discriminate between upper plate and slab events.

Recommendation 3

Modern waveform analysis of pre-digital and historical

intraslab earthquakes in Cascadia for source properties

(targets: 1939, 1946 (?), 1949, 1965, 1976 slab events).

• Invert waveforms for centroid moment tensor (CMT)

Mij (focal mechanism), scalar moment Mo, and moment magnitude Mw. Reconcile with Pacific Northwest network first-motion focal mechanisms.

• Establish source shapes and directivity. Relate source

shapes to the geometry of seismogenic zone.

• Invert for radiated energies, Eo. Why does Eo/Mo tend

to be high compared with interplate thrust earthquakes?

Possible payoffs:

• Basic improvement of the magnitude estimates for

large and damaging slab earthquakes in Cascadia.

• Improved Gutenberg–Richter distribution and

seismicity-rate data.

• Improvement in models for slab seismic deformation

and deformation rates.

• Better constraints on ground-motion models in relation to the observed strong motions and earthquake

effects.

Recommendation 4

Active and passive seismic experiments and associated

modeling that image slab wavespeed anomalies (plus

forearc structure) and refine slab geometry vis a vis

intraslab earthquake hypocenters.

• Regional tomography from active sources and earthquakes.

• Reflections off slab interfaces.

• Detection of possible slow, high-frequency phases

trapped in a shallow crustal waveguide in the slab.
The Cascadia Subduction Zone and Related Subduction Systems | 3

• Teleseismic tomography and receiver-function analysis (including S waves).

• Body-wave conversions from teleseismic sources.

• Attenuation tomography?

Possible payoffs:

• Independent mapping of the spatial limits of the zone

of slab seismogenesis based on theoretical relationship between slab mineralogical changes, their effects

on wavespeed anomalies and on the release of water

triggering slab earthquake faulting.

• Improved absolute hypocenter locations using threedimensional wavespeed models and hence improved

slab geometry and deformation (see recommendation 2).

• Develop a physical basis for the low level of slab

seismicity beneath western Oregon, including Portland and its environs.

• Improve strong ground-motion models.

• Investigate possible slab “bank shot” ground motions

due to reflection of waves from shallow forearc sources

(e.g., along the Seattle fault) reflecting off the slab structure (e.g., effects of Moho reflections on ground motions from the 1989 Loma Prieta earthquake).

Recommendation 5

Develop two- and three-dimensional theoretical models for slab earthquake seismogenesis. Compare slab deformation rates inferred from models with those inferred

from observed slab seismic moment release rates and

their spatial distribution.

• Slab shape and kinematics, thermal structure, seismic

structure and lab findings as inputs.

• Mineralogy and its changes with P–T conditions, fluid

evolution and its effects on deformation, and deformation state and stress state as outputs.

Possible payoffs:

• Could establish upper bounds on slab earthquake-moment production rates and their spatial distribution.

• Could develop a physical basis for mapping the spatial limits of seismogenesis, including the low level of

slab seismicity beneath western Oregon, including

Portland and its environs.

Recommendation 6

Followup workshops: proposed workshops, mini-workshops or forums (some that could coincide with SSA or

AGU annual meetings).

• Planning for seismic experiments with slab targets

(Tacoma SHIPS and Oregon SHIPS; Canadian ocean

bottom seismometer and broadband experiments,

etc.) March 19–20, 2001, USGS/University of Washington support. Kirby, Brocher and Creager (see recommendation 4 above). Workshop held as planned.

• Historical slab earthquakes in Cascadia, southern

Alaska, the Caribbean and in analog foreign subduction systems (see recommendations 2 and 3). Coordinate the efficient collection and processing of

seismograms.

• Public forum on slab earthquake hazards, Seattle or

Olympia, 2 days. Improve public awareness of slab

earthquake hazards, develop partnerships between

USGS, universities, state and local government agencies, and support for slab investigations.

• Geological and geodetic expressions of slab earthquakes and slab deformation (uplift or subsidence with

respect to sea level, liquifaction, debris flows or debris

avalanches off Mount Ranier). University of Washington at Seattle. Include State of Washington representatives. Search for geological information that provides

a longer-term record of slab earthquake deformation

and effects than the short instrumental era provides.

Possible payoffs:

• Enhanced focus on intraslab earthquakes.

• Cooperative, multidisciplinary and complex projects

can be made possible (e.g., SHIPS).

Recommendation 7

International cooperation in countries having analog subduction systems and modern regional seismic networks.

Promote exchange of scientists and data.

• Canada: PGC/GSC, UBC, UVic. Rapid data exchange

of phase data/waveforms? Cross-border seismic experiments?

• Japan: JAMSTEC, Subduction Dynamics Frontiers Program (focus on Nankai). Planned 2003–2004 ocean

bottom seismometer deployment off the northern

Cascadia margin. Tohoku University: Tohoku seismology group.

• Mexico: UNAM (Instituto de Geofisica) is developing

a hazards program with internal funding for new work.

• Costa Rica: Universidad Nacional de Costa Rica.

Possible payoffs:

• Can leverage information on a greater number of slab

earthquakes in similar settings and thereby identify

global trends in seismicity rates, ground motions, and

seismic effects of slab earthquakes in warm-slab environments.

• Cross-fertilization of ideas and approaches.

References

Wilson, D.S. The Juan de F uca plate and slab: isochron

structure and Cenozoic plate motions, in The Cascadia

Subduction Zone and Related Subduction Systems,

edited by S.H. Kirby , K. Wang and S.G. Dunlop, pp.

9–12, U.S. Geological Survey Open-F ile Report 02–

328, Geological Survey of Canada Open File 4350,

2002.
4 |
The Cascadia Subduction Zone and Related Subduction Systems | 5

Introduction to a global systems approach to Cascadia

slab processes and associated earthquake hazards

Steve H. Kirby1 and Kelin Wang2

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

skirby@usgs.gov, wang@pgc.nrcan.gc.ca

Intraslab (Wadati–Benioff zone) earthquakes are often

under-appreciated as earthquake hazards. In Latin

America, such earthquakes have, in fact, had far greater

human impact in the 20th century than great interplate

earthquakes. Plate boundary earthquakes may have

larger magnitudes but they occur largely offshore. Shallow intraslab earthquakes in the Americas may cause

greater hazards because they usually occur inland where

population centers tend to be located. Other factors, such

as systematically higher energy yields at high frequency

[Choy et al., this volume] and the triggering of debris

flows and avalanches on steep slopes of nearby orogenic

and volcanic mountain belts also may contribute to the

larger impacts of intraslab earthquakes.

Intraslab earthquakes present a special challenge for

hazard appraisal because the causative faults do not rupture the ground surface, and, unlike plate-boundary earthquakes that occur on a single fault, their moment release

represents distributed deformation on many faults. Their

depths and localization in the subducting slab mute their

expressions on the surface. With a few exceptions, geologic and geodetic observations recording uplift or subsidence and other expressions of subsurface slip have

rarely been recognized.

The Cascadia subduction system has repeatedly sustained the effects of large, destructive intraslab shocks in

the last century [Rogers and Crosson, this volume]. After

this workshop was held, the M=6.8 Nisqually earthquake

occurred on February 28, 2001 at a depth of about 53

to 58 km beneath the southern Puget lowland. This event

was within the Wadati–Benioff zone defined by previous slab earthquakes, including five destructive M>5.5

events since 1945. All focal mechanism determinations

of the 2001 earthquake indicate that it was a normal

faulting event and only four small aftershocks, with magnitudes ranging from 1.0 to 4.3, occurred during the year

following the main shock. The governor of the State of

Washington estimated that the monetary losses connected with this earthquake exceeded US$2 billion. This

most recent shock is a reminder that such earthquakes

are a recurring expression of an ongoing process of slab

deformation in a seismically active belt including the

Willamette Valley in western Oregon, the Puget lowland

in western Washington, and the Georgia Strait between

Vancouver Island and the Canadian mainland.

Recent integrated models for slab thermal structure,

mineralogy, slab deformation and seismic activity at intermediate depths have given useful insight into the patterns of earthquake hypocenter distribution and the fine

structure of seismic wavespeeds in slabs and forearcs [e.g.,

Kirby et al., 1996; Peacock and Wang, 1999; Kirby, 2000].

In particular, these models recognize that earthquakes

deeper than about 35 km are unlikely to occur by ordinary brittle fracture and frictional sliding. It is widely

believed that deeper earthquakes occur by a process of

dehydration embrittlement where internal pore fluid

pressure due to water liberated from hydrous minerals

reduces the effective normal stresses along faults and

permits fracture and frictional sliding to occur provided

that the shear stresses are sufficiently high [Raleigh, 1967].

More recent investigations of the roles of fluids released from slabs by dehydration suggest that faults created in the ocean basins at mid-ocean ridges and trenches

may be reactivated in descending slabs by dehydration

embrittlement [Kirby et al., 1996; Kirby, 2000].

Not only do metamorphic reactions in mafic oceanic

crust liberate water but they also densify the crust by 12

to 15%, ultimately producing eclogite. If the crust does

not detach from the underlying stable slab mantle, such

large volume changes should cause large internal slab

stresses, extensional in the crust and compressional in

the underlying mantle [Kirby et al., 1996]. Thus, metamorphism of the oceanic crust should produce both an

internal source of slab stresses and a mechanism for stress

release by faulting associated with dehydration. The depths

at which such metamorphic changes occur depend largely

on slab thermal structure. In “warm” slabs (involving young

lithosphere or older lithosphere subducting at slow rates)

the mafic crust (hydrous and anhydrous) probably

densifies to eclogite facies at shallow depths (<100 km)

(Figure 1), whereas slab dehydration and eclogite formation in “cold” slabs probably persists to far greater depths
6 | Kirby and Wang: Global systems approach to Cascadia slab processes and associated earthquake hazards

(>>100 km) [Kirby et al., 1996; Peacock and Wang,

1999]. The seismic deformation associated with such

densification is therefore expected to be largest at shallowest depths in warm-slab environments. Its specific

seismic expression depends on how it interacts with other

sources of stress, such as the forces at triple junctions

[ten Brink, this volume], slab flexure, interplate coupling

(and its variation in the plate-boundary earthquake cycle)

[Taylor et al., 1996], membrane deformations accompanying subduction with reverse trench curvature [Chiao

and Creager, this volume], and slab interactions with the

underlying asthenosphere [Wang et al., this volume].

These mineralogical models for warm-slab subduction [Hacker et al., Peacock et al. and McKenna and

Blackwell, all this volume] are broadly consistent with

the observed characteristics of warm subduction (shallow intraslab earthquakes, feeble or absent arc volcanism,

localized “adakitic” arc lava compositions) compared to

cold subduction (earthquakes to much greater depths

and “normal” arc volcanism). Moreover, the models are

largely consistent with the fine wavespeed structures and

focal mechanisms of earthquakes in slabs in relation to

the fine structure of the Wadati–Benioff zone. Cascadia

is a classic end-member warm-slab environment and displays many of these characteristic features (see papers in

the section on other warm-slab subduction systems).

These preliminary models for warm-slab evolution

with descent are promising, and hence it follows that a

global “top-to-bottom” systems approach to providing

constraints on key factors important to seismic hazard

appraisal is called for. Our goals guiding this systems approach include improving the constraints on seismicity

rates and depth distribution (and how these vary along

strike), better estimates of the maximum credible earthquake size, and improvements in the factors that govern

model estimates of strong ground motions in the Cascadia

subduction system, including the source characteristics

of intraslab events and the fine wavespeed and attenuation structure of the slab and forearc. The components

of such an approach should therefore include:

1. The state of stress and the thermal, hydrologic, fault

and mineralogical structures of the offshore oceanic

plates entering the “trench” of the Cascadia system,

as deduced from marine geology and geophysics

[Goldfinger et al., this volume]. The subducted slab

travels northward in the mantle, and this has a firstFIGURE 1: Mineralogical structure during warm subduction (after Kirby [2000] and based on the models of Kirby et al. [1996], Peacock

and Wang [1999], and Peacock and Hyndman [1999]). The crust in the subducting plate is as much as 300 to 500°C warmer than for

cold subduction. This hotter thermal structure restricts the region of interplate thrust seismic slip on the plate contact to shallower

depths (defined by points A to B or A to C) but generally over a larger area of slip than for cold slabs, producing the potential for very

large offshore interplate earthquakes. The release of water stored in crustal minerals at shallow depths of 40 to 90 km is expected to

promote shallow metamorphism of normal crustal rocks (blue) to eclogite (purple), and facilitate shallow intraslab earthquakes (red

dots) within slabs. Shallow dewatering also is expected to reduce partial melting of the asthenosphere and reduce volcanic vigor.

Higher temperatures may also induce slab crustal melting (M?) and produce anomalous magma chemistries for such volcanoes (open

red triangle). The contour interval is 100oC, with the surface held at 0oC
The Cascadia Subduction Zone and Related Subduction Systems | 7

order effect on the stress pattern of the offshore plate

[Wang et al., 1997] and potentially on intraslab stresses.

Knowledge of the heat flow and hydrogeologic

regime of the Juan de Fuca (JDF) plate prior to subduction [Davis et al., 1999] is a prerequisite for any

understanding of the thermal and metamorphic processes of the plate after subduction. The isochron structure of the exposed and subducted Explorer, Juan de

Fuca and Gorda plates [Wilson, this volume] is an important constraint on the thermal structure of these

plates and slabs.

2. High-resolution redetermination of intraslab earthquake hypocenters in light of improved models of the

seismic wavespeed and attenuation structure of the

Cascadia slabs and forearc. This work obviously needs

to build upon recent seismic tomography based on

seismic station data from the Pacific Northwest and

GSC networks [Villasenor et al., this volume], and on

the observations made during SHIPS [Trehu et al. and

Crosson et al., this volume]. This element should also

include comparisons with seismic structures in other

warm slab environments, such as those in Latin

America [Singh et al. and Protti, this volume] and

Southwest Japan [Cummins et al., this volume].

3. We also need to exploit the global instrumental record

for large intraslab earthquakes in other warm-slab environments, including appraisal of the magnitudes,

focal mechanisms, and hypocenter distributions of

large intraslab shocks in the pre-digital era. This is necessary in order to overcome the problem of generally

low slab seismicity rates of large intraslab earthquakes

in Cascadia and to understand the effect of subduction

parameters on the variability of intraslab earthquakes

in these settings [Kirby et al., this volume].

4. Improvements in our knowledge of the heat transfer,

rheology, phase equilibria, and kinetic properties of

the oceanic crust and mantle, especially at estimated

slab P–T conditions. Deformation experiments on

hydrous mafic and ultramafic rocks at the P–T conditions where the crust in the Cascadia slab is dehydrating under stress are needed in order to test further

the dehydration–embrittlement mechanism and to

search for other potential faulting mechanisms.

5. Using the above constraints, construct two-dimensional

thermo-mechanical models of warm-slab evolution

with descent, including theoretical estimates of slab

stresses and deformation rates. Extend these models

to three dimensions by including knowledge of the

reconstructed isochron structure of the subducted

parts of the Explorer, Juan de Fuca and Gorda plates,

and the along-strike variations of convergence rates

and slab dip. These models should give insight into

the marked along-strike variations in intraslab

seismicity and earthquake hazards in Cascadia.

6. Use the results of these well-constrained numerical

models to reduce the uncertainties in the factors noted

above that influence the hazard appraisal for intraslab

earthquakes.

References

Chiao, L.-Y., and K.C. Creager, Geometry and membrane

deformation rate of the subducting Cascadia slab, in

The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 47–54, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Choy, G.L., J.L. Boatwright, and S.H. Kirby, The radiated

seismic energy and apparent stress of interplate and

intraslab earthquakes at subduction zone environments: implications for seismic hazard estimation, in

The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 107–114, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Crosson, R.S., N.P. Symons, K.C. Creager, L.A. Preston, T.

Van Wagoner, T.M. Brocher, M.A. Fisher and the

SHIPS98 Working Group , Three-dimensional structure of the Cascadia forearc region from SHIPS active

experiment and earthquake observations: tomographic

inversion provides a high-resolution view into the core

of the Cascadia forearc complex, in The Cascadia Subduction Zone and Related Subduction Systems, edited

by S.H. Kirby, K. Wang, and S.G. Dunlop, p. 33, U.S.

Geological Survey Open-File Report 02–328, Geological Survey of Canada Open File 4350, 2002.

Cummins, P.R., Y. Kaneda, and K.C. Hori, Large, intermediate-depth earthquakes in Southwest Japan, in

The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 99–102, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Davis, E.E., D.S. Chapman, K. Wang, H. Villinger, A.T.

Fisher, S.W. Robinson, J. Grigel, D. Pribnow, J. Stein,

and K.Becker, Regional heat-flow variations across the

sedimented Juan de Fuca ridge eastern flank: constraints on lithospheric cooling and lateral hydrothermal heat transport, J. Geophys. Res., 104, 17,675–

17,688, 1999.

Goldfinger, C., R.P. Dziak, and C. Fox, Offshore structure

of the Juan de Fuca plate from marine seismic and

sonar studies, in The Cascadia Subduction Zone and

Related Subduction Systems, edited by S.H. Kirby, K.

Wang, and S.G. Dunlop, pp. 13–16, U.S. Geological

Survey Open-File Report 02–328, Geological Survey

of Canada Open File 4350, 2002.

Hacker, B.R., G.A. Abers, and S.M. Peacock, Theoretical

mineralogy, density, seismic wavespeeds and H2

O
8 | Kirby and Wang: Global systems approach to Cascadia slab processes and associated earthquake hazards

content of the Cascadia subduction zone, with implications for intermediate-depth seismicity and earthquake hazard, in The Cascadia Subduction Zone and

Related Subduction Systems, edited by S.H. Kirby, K.

Wang, and S.G. Dunlop, pp. 133–138, U.S. Geological Survey Open-File Report 02–328, Geological Survey of Canada Open File 4350, 2002.

Kirby, S.H., Taking the temperature of slabs, Nature, 403,

31–34, 2000.

Kirby, S.H., E.R. Engdahl, and R. Denlinger, Intermediate-depth intraslab earthquakes and arc volcanism

as physical expressions of crustal and uppermost

mantle metamorphism in subducting slabs, in Subduction: Top to Bottom, Geophysical Monograph 96,

edited by G. Bebout, D. Scholl, S. Kirby, and J. Platt,

pp. 195–214, American Geophysical Union, Washington, D.C., 1996.

Kirby, S.H., E.R. Engdahl and A. Villaseñor, Warm-slab

subduction as a global process, in The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G. Dunlop, p. 75,

U.S. Geological Survey Open-File Report 02–328,

Geological Survey of Canada Open File 4350, 2002.

McKenna, J.R., and D.D. Blackwell, Heat sources in the

subduction zones: implications for slab seismicity and

arc volcanism in The Cascadia Subduction Zone and

Related Subduction Systems, edited by S.H. Kirby, K.

Wang, and S.G. Dunlop, pp. 127–132, U.S. Geological Survey Open-File Report 02–328, Geological Survey of Canada Open File 4350, 2002.

Peacock, S.M., and R. Hyndman, Hydrous minerals in

the mantle wedge and the maximum depth of subduction thrust earthquakes, Geophys. Res. Lett., 26,

2517–2520, 1999.

Peacock, S.M., and K. Wang, Seismic consequences of

warm versus cool subduction metamorphism: examples from Southwest and Northeast Japan, Science,

286, 937–939, 1999.

Peacock, S., K. Wang, and A.M. McMahon, Thermal structure and metamorphism of subducting oceanic crust:

insight into Cascadia intraslab earthquakes in The

Cascadia Subduction Zone and Related Subduction

Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 123–126, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Protti, M., Seismicity and seismotectonics of Central

America with emphasis on intraslab seismic activity

in The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and

S.G. Dunlop, pp. 93–98, U.S. Geological Survey

Open-File Report 02–328, Geological Survey of

Canada Open File 4350, 2002.

Raleigh, C.B., Tectonic implications of serpentinite weakening, Geophys. J. Royal Astro. Soc., 14, 113–118, 1967.

Rogers, G.C., and R.S. Crosson, Intraslab earthquakes

beneath Georgia Strait/Puget Sounds, in The Cascadia

Subduction Zone and Related Subduction Systems,

edited by S.H. Kirby, K. Wang, and S.G. Dunlop, p.

68, U.S. Geological Survey Open-File Report 02–328,

Geological Survey of Canada Open File 4350, 2002.

Singh, S.K., V. Kostoglodov, and J.F. Pacheco, Intraslab

earthquakes in the subducting oceanic plates below

Mexico, in The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K.

Wang, and S.G. Dunlop, pp. 87–92, U.S. Geological

Survey Open-File Report 02–328, Geological Survey

of Canada Open File 4350, 2002.

Taylor, M.A.J., G. Zheng, and J.R. Rice, Cyclic stressing

and seismicity at strongly coupled subduction zones,

J. Geophys. Res., 101, 8363–8381, 1996.

ten Brink, U.S., Slab deformation at the edge of subduction zones: an example from Cascadia, in The Cascadia

Subduction Zone and Related Subduction Systems,

edited by S.H. Kirby, K. Wang, and S.G. Dunlop, pp.

139–142, U.S. Geological Survey Open-File Report

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4350, 2002.

Tréhu, A.M., T.M. Brocher, K.C. Creager, M.A. Fisher, L.A.

Preston, G. Spence and the SHIPS98 Working Group,

Geometry of the subducting Juan de Fuca plate: new

constraints from SHIPS98, in The Cascadia Subduction Zone and Related Subduction Systems, edited

by S.H. Kirby, K. Wang and S.G. Dunlop, pp. 25–32,

U.S. Geological Survey Open-File Report 02–328,

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Villaseñor, A., D. Stanley, and H. Benz, Velocity structure of western Washington from local earthquake

tomography, in The Cascadia Subduction Zone and

Related Subduction Systems, edited by S.H. Kirby, K.

Wang and S.G. Dunlop, pp. 35–36, U.S. Geological

Survey Open-File Report 02–328, Geological Survey

of Canada Open File 4350, 2002.

Wang, K., J. He, and E.E. Davis, Transform push, oblique

subduction resistance, and intraplate stress of the Juan

de Fuca plate, J. Geophys. Res., 102, 661–674, 1997.

Wang, K., J. He, E.E. Davis, and C. Goldfinger, Stresses in

the Juan de Fuca Plate and the role of mantle resistance to horizontal slab motion, in The Cascadia Subduction Zone and Related Subduction Systems, edited

by S.H. Kirby, K. Wang, and S.G. Dunlop, p. 143,

U.S. Geological Survey Open-File Report 02–328,

Geological Survey of Canada Open File 4350, 2002.

Wilson, D.S. The Juan de Fuca plate and slab: isochron

structure and Cenozoic plate motions, in The Cascadia

Subduction Zone and Related Subduction Systems,

edited by S.H. Kirby, K. Wang and S.G. Dunlop, pp.

9–12, U.S. Geological Survey Open-File Report 02–

328, Geological Survey of Canada Open File 4350,

2002.
The Cascadia Subduction Zone and Related Subduction Systems | 9

The Juan de Fuca plate and slab:

Isochron structure and Cenozoic plate motions

Douglas S. Wilson

Department of Geological Sciences, University of California at Santa Barbara, Santa Barbara, California, 93106, USA

wilson@rapa.geol.ucsb.edu

The Juan de Fuca ridge has a long history of complex

tectonics, with propagating ridge offsets having been

more common than stable transform faults for most of

the time since 25 Ma. The spreading rate has been intermediate at about 50–65 mm/yr since about 20 Ma. Several incremental clockwise shifts in the spreading

direction during this time have led to reorganization of

the propagating rift pattern. The most prominent shift

occurred during a Pacific-wide reorganization at 5.9 Ma,

when Juan de Fuca–Pacific relative motion changed by

10-15 ̊and the first stable transform fault segment of the

Blanco fracture zone formed. Subsequent propagation

has shortened pre-existing ridge segments that became

intra-transform spreading centers when their bounding

offsets became additional strands of the transform fault

system, eventually developing into the four major strands

of the modern Blanco transform. The Juan de Fuca plate

has not remained rigid as its motion relative to the Pacific plate has evolved.

Since at least 10 Ma, the strike of the Mendocino

transform fault has not reoriented to remain parallel to

the increasingly clockwise direction of Juan de Fuca–Pacific relative motion. Since at least 3 Ma, the resulting

space problem has been accommodated in a diffuse

boundary zone east of the southern Gorda ridge commonly (but erroneously) known as the Gorda plate. Prior

accommodation may have included a minor component

of thrusting along the Mendocino transform. At roughly

4 Ma, the Explorer micro-plate separated from the northern Juan de Fuca plate and the motion of the Explorer

plate relative to North America is slow enough that it

could be considered to have been captured by the North

American plate. The Sovanco transform zone between

the Explorer and Pacific plates is either extremely broadly

distributed or has migrated north and south numerous

times (probably both), and remains poorly understood.

As is the case with other plate pairs in the Pacific,

minor changes in plate motion since 3 Ma indicate that

0–0.78 Ma averages may be more indicative of modern

motions than 0–3 Ma averages. The recent averages give

spreading rates of 55 mm/yr for Endeavour segment and

51 mm/yr for northern Gorda ridge [Wilson, 1993].

In a reference frame fixed to North America, convergence of the Juan de Fuca plate has been roughly constant since 20 Ma at the slow to intermediate rate of

about 40 mm/yr in a northeasterly direction. Current rates

increase northward from about 30 mm/yr at 42 ̊N to

about 45 mm/yr at 49 ̊N, but rate gradients have varied

substantially through time. Recent revisions in the determination of current Pacific–North America relative motion, as well as corrections for possible slight biases in

converting magnetic anomaly width to spreading rate,

both indicate a convergence direction several degrees

counterclockwise from previous determinations. Northward motion of the continental margin relative to rigid

North America reduces the obliquity of the convergence

direction [McCaffrey et al., 2000].

The age structure of the slab can be extrapolated from

offshore isochrons with fairly high confidence. A narrow,

young and barely subducting slab attached to the Explorer plate lies under northwestern Vancouver Island.

From southern British Columbia to central Oregon the

inferred age pattern is very simple, with age increasing

eastward, about 12 Ma below the coastline. In this area,

structural complexities inherited from shear deformation

of young seafloor near the ridge axis are restricted to

narrow, northeast-striking zones. Below southern Oregon

and northern California, the age pattern is much more

complex. A broad area was sheared prior to formation

of the Blanco transform at 5.9 Ma and the south edge of

the slab has been pervasively sheared as it passed through

the Gorda deformation zone over the past several million years. The location of active seismicity at the south

end of the slab is fairly well predicted by simple models

that treat the corner of the Pacific plate at the Mendocino

triple junction as a rigid indentor into the much weaker

Juan de Fuca plate, with little or no influence from inherited structures in the slab.

References

Botros, M., and H.P. Johnson, Tectonic evolution of the

Explorer-northern Juan de Fuca region from 8 Ma to

the present, J. Geophys. Res., 93, 10,421–10,437,

1988.
10 | Wilson: Juan de Fuca plate and slab: isochron structure and Cenozoic plate motions

FIGURE 1: History of estimates of convergence velocity between Juan de Fuca and stable North America, calculated for 46°N, 125°W.

All estimates are based on vector subtraction of Pacific–North America and Pacific–Juan de Fuca rates. Changes in Pacific–North

America rates reflect significant changes in input data, while changes for Juan de Fuca are dominated by time-scale revisions. The

current estimate for Juan de Fuca–Pacific includes a small correction based on the assumption that the axial magnetic anomaly is one

kilometer wider than the plate motion since the magnetic polarity reversal. Resulting rotation vector for North America–Juan de Fuca

is 33.7°N, -115.1°E, 1.257°/Myr. Actual convergence rates on the subduction zone will need a correction for motion of the continental

margin relative to stable North America, up to 6–9 mm/yr northward [McCaffrey et al., 2000].
The Cascadia Subduction Zone and Related Subduction Systems | 11

FIGURE 2: Seafloor age map, modified from Wilson [1988; 1993], and prediction of slab age, modified from Wilson [1988] to include

nonrigid plate behavior. No correction for slab dip has been applied. Areas in the slab with north–south isochron strikes should be

undeformed; other areas experienced significant shear prior to subduction.
12 | Wilson: Juan de Fuca plate and slab: isochron structure and Cenozoic plate motions

Cowan, D.S., M. Botros, and H.P. Johnson, Bookshelf tectonics: rotated crustal blocks within the Sovanco fracture zone, Geophys. Res. Lett., 13, 995–998, 1986.

Denlinger, R.P., A model for large-scale plastic yield of

the Gorda deformation zone, J. Geophys. Res., 97,

15,415–15,423, 1992.

DeMets, C., and T.H. Dixon, New kinematic models for

Pacific–North America motion from 3 Ma to present,

I: evidence for steady motions and biases in the NUVEL1A model, Geophys. Res. Lett., 26, 1921–1924.

DeMets, C., R.G. Gordon, D.F. Argus, and S. Stein, Current plate motions, Geophys. J. Int, 101, 425–478,

1990.

Embley, R.W., and D.S. Wilson, Morphology of the Blanco

transform, northeast Pacific: implications for its tectonic evolution, Mar. Geophys. Res., 14, 25–45, 1992.

Hey, R.N., A new class of “pseudofaults” and their bearing on plate tectonics: a propagating rift model, Earth

Planet. Sci. Lett., 37, 321–325, 1977.

Hey, R.N., and D.S. Wilson, Propagating rift explanation

for the tectonic evolution of the northeast Pacific:

the pseudomovie, Earth Planet. Sci. Lett., 58, 167–

188, 1982.

Hyndman, R.D., R.P. Riddihough, and R. Herzer, The

Nootka fault zone: a new plate boundary off western Canada, Geophys. J. R. Astron. Soc., 58, 667–

683, 1979.

Jachens, R.C., and A. Griscom, Three-dimensional geometry of the Gorda plate beneath northern California, J. Geophys. Res., 88, 9375–9392, 1983.

Johnson, H.P., J.L. Karsten, J.R. Delaney, E.E. Davis, R.G.

Currie, and R.L. Chase, A detailed study of the Cobb

offset of the Juan de Fuca ridge: evolution of a propagating rift, J. Geophys. Res., 88, 2297–2315, 1983.

Kleinrock, M.C., and R.N. Hey, Migrating transform and

lithospheric transfer at the Galapagos 95.5°W propagator, J. Geophys. Res., 94, 13,859–13,878, 1989.

McCaffrey, R., and 6 others, Rotation and plate locking

at the southern Cascadia subduction zone, Geophys.

Res. Lett., 27, 3117–3120, 2000.

McKenzie, D., The geometry of propagating rifts, Earth

Planet. Sci. Lett., 77, 176–186, 1986.

Riddihough, R.P., A model for recent plate interactions

off Canada’s west coast, Can. J. Earth Sci., 14, 384–

396, 1977.

Riddihough, R., Recent plate motions of the Juan de Fuca

plate system, J. Geophys. Res., 89, 6980-6994, 1984.

Silver, E.A., Tectonics of the Mendocino triple junction,

Geol. Soc. Am. Bull., 82, 2965–2978, 1971.

Vine, F.J., Magnetic anomalies associated with mid-ocean

ridges, in The History of the Earth’s Crust, edited by

R.A. Phinney, pp. 73–89, Princeton University Press,

Princeton, N. J., 1968.

Wang, K., J. He, and E.E. Davis, Transform push, oblique

subduction resistance and intraplate stress of the Juan

de Fuca Plate, J. Geophys. Res., 102, 661–674, 1997.

Wilson, D.S., A kinematic model for the Gorda deformation zone as a diffuse southern boundary of the Juan

de Fuca plate, J. Geophys. Res., 91, 10,259–10,269,

1986.

Wilson, D.S., Tectonic hisotry of the Juan de Fuca ridge

over the last 40 million years, J. Geophys. Res., 93,

11,863–11,876, 1988.

Wilson, D.S., Deformation of the so-called Gorda plate,

J. Geophys. Res., 94, 3065–3075, 1989.

Wilson, D.S., Kinematics of overlapping rift propagation

with cyclic rift failure, Earth Planet. Sci. Lett., 96, 384–

392, 1990.

Wilson, D.S., Confidence intervals for motion and deformation of the Juan de Fuca plate, J. Geophys. Res.,

98, 16,053–16,071, 1993.

Wilson, D.S., R.N. Hey, and C. Nishimura, Propagation

as a mechanism of reorientation of the Juan de Fuca

Ridge, J. Geophys. Res., 89, 9215–9225, 1984.
The Cascadia Subduction Zone and Related Subduction Systems | 13

Offshore structure of the Juan de Fuca plate

from marine seismic and sonar studies

Chris Goldfinger1, Robert P. Dziak2 and Chris Fox2

1 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97331, USA

2 National Oceanic and Atmospheric Administration, Pacific Marine Environmental Laboratory

Hatfield Marine Science Center, Newport, Oregon, 97365, USA

gold@oce.orst.edu, dziak@pmel.noaa.gov, fox@pmel.noaa.gov

Juan de Fuca deformation

Using sidescan sonar, seismic reflection profiles and swath

bathymetric data, we have mapped a set of west–northwest trending left-lateral strike-slip faults that deform the

Oregon and Washington submarine forearc. Evidence for

left-lateral separation includes: offset of accretionary

wedge folds, channels and other surficial features; sigmoidal left bending of accretionary wedge folds; and

offset of abyssal plain basement and sedimentary units.

Five of these faults cross the plate boundary, extending

5–21 km into the Juan de Fuca plate. Using offset of subsurface piercing points and offset of approximately dated

submarine channels, we calculate slip rates for these five

faults of 5.5 to 8.5 mm/yr. Little or no offset of these faults

by the basal thrust of the accretionary wedge is observed.

Holocene offset of submarine channels and unconsolidated sediments is observed in sidescan records and

directly by submersible [Goldfinger et al., 1997].

We also observe compressional and extensional regions near the ends of the rotating blocks that are consistent with the expected space problems inherent in any

block rotation kinematic model. The presence of underlying differential slip on the subducting slab has left a

characteristic pattern of accretionary wedge development. The slip of the underlying slab faults results in local apparent variations in plate convergence on the

décollement, resulting in variations in rates of growth of

overlying accretionary wedge anticlines.

The strike-slip faults are most likely driven by dextral

shearing of the subducting slab and propagate upward

through the overlying accretionary wedge. Plate coupling

forces are insufficient to account for rupture of the slab,

and thus independent slab rupture is the best explanation for the presence of these structures [Goldfinger et al.,

1997]. Tangential hydrodynamic drag caused by oblique

insertion of the slab into the mantle is a possible driving

mechanism. Four sinistral faults observed in only the upper plate may be remnant traces of previous basementdriven deformation. Such traces are expected as the slab

faults more to the northeast with the Juan de Fuca plate.

A model of overall right-lateral simple shear of the submarine forearc is consistent with the observed surface

faults, which may be R’ or antithetic shears to the overall

right-shear couple. The major strike-slip faults define elongate blocks that, because of their orientation and sinistral

slip direction, must rotate clockwise. We infer that the

deformation of the submarine forearc (here defined to

include the lower plate) is highly strain-partitioned into

arc-normal shortening and arc-parallel strike-slip and translation. The high slip rates of the strike-slip faults, coupled

with the lack of offset of these faults as they cross the

plate boundary, imply that the seaward accretionary wedge

is not moving at the expected convergence rate relative

to the subducting plate. We conclude that the accretionary wedge is rotating and translating northward, driven

by the tangential component of Juan de Fuca–North

American plate convergence. This conclusion is also supported by the lack of offset across the deformation front

between large landslide scarps and associated debris packages on the abyssal plain. These slides range in age from

14 ka to 1000 ka, yet the debris and the scarps appear to

be largely in their original relative positions [Goldfinger et

al., 2000].

Scholz and Campos[1995] proposed a dynamic model

of interplate coupling and decoupling at subduction zones

that incorporates the hydrodynamic resistance of the

motion of the slab through the viscous mantle. In their

model, the mantle is considered stationary in a hot spot

reference frame. This “sea anchor” force, in a mantle system that is fixed to the hot spot frame, should have trenchparallel and trench-normal components in the case of

oblique subduction [Scholz and Campos, 1995] (Figure

1). The trench–parallel component of hydrodynamic mantle resistance to subduction is an unbalanced force that

sets up a shear couple in the plane of the slab. For

Cascadia, the shear couple would be dextral, and antithetic shears should be left-lateral and east–west to northwest trending. We propose that this shear couple may be

responsible for the observed transverse rupture of the slab
14 | Goldfinger, Dziak and Fox: Offshore structure of the Juan de Fuca plate from marine seismic and sonar studies

we have observed in the central Cascadia forearc. Limited slab seismicity is consistent with a model of transverse shearing in the subducting slab. The largest

instrumentally recorded earthquake in Cascadia (1949,

Mb=7.1; 47.13°N, 122.95°W [Baker and Langston, 1987])

occurred in the Puget Sound region in Washington. The

focal depth of 54 km places this event in the upper part

of the subducting plate [Baker and Langston, 1987; Ludwin

et al., 1991]. The focal mechanism is strike-slip with a

preferred fault plane striking east–west ± 15° (azimuth

255–285°) and has nearly pure left-lateral slip [Baker and

Langston, 1987]. While this strike-slip mechanism indicates shearing within the Juan de Fuca plate, the fault

planes of the 1949 event are not parallel to the convergence direction, negating a possible tear fault origin

[Weaver and Baker, 1988]. A similar left-lateral strike-slip

solution was determined for a small 1981 earthquake

(magnitude 3.3) in the upper Juan de Fuca plate on the

southern Washington coast (Figure 2) [Taber and Smith,

1985; Weaver and Baker, 1988]. The Olympia event, and

the Washington coast strike-slip event, lie at the eastern

end of an east–west linear trend of earthquakes that lie

roughly on the projection of one of the basement sinistral

faults, the South Nitinat fault, and are near a major eastwest reverse fault mapped on the shelf. We suggest that

these slab events may have occurred on the downdip part

of this known slab structure.

Could these basement-involved faults be reactivated

structures in the oceanic basalt? Comparison of the strikes

of the oblique faults to the seafloor magnetic anomaly

map of Wilson [1993] shows that the three Washington

faults, with strikes of about 283°, could be reactivated

small fracture zones perpendicular to the magnetic

anomalies. The Oregon faults strike 292° to 325°, deviating from an orthogonal to the magnetic lineations by 12°

to 45°, and thus are unlikely to be related to pre-existing

basement structure. We have also considered conjugate

shortening of the forearc of the type described by Lewis

et al. [1988]; however, extensive mapping of margin structures [Wagner et al., 1986; Clarke, 1990; this paper] has

not identified structures conjugate to the west–northwest

faults. We have mapped two major northeast trending

structures, one on the slope off northern Oregon and

one off central Washington, but both of these are clearly

left-lateral tear faults of the accretionary fold-thrust belt.

Implications for plate coupling

In much of the Oregon and northern California forearcs,

two domains of fold orientation are separated by a break

in topographic slope. The dominant north–northwest to

west–northwest trending structural grain of the upper

slope and shelf is nearly orthogonal the plate convergence

direction, suggesting strong plate coupling. The lower

slope, in contrast, consists of margin parallel folds that

are more open, and have mixed or seaward vergence

directions. The sharp division of structural domains in

the mid-continental slope may correspond to the seismic front of Byrne et al. [1988]. We speculate that seaward of this boundary, basal shear stress is less than that

needed to form structures normal to the maximum principal stress (σ1

), and that a “backstop” effect at or near

the slope break (SB) dominates the margin-parallel orientation of the youngest structures [Goldfinger et al., 1992;

1996]. Landward of this boundary, folds that are oblique

to the continental margin but close to normal to the inferred plate convergence direction, dominate structural

orientations, suggesting strong coupling eastward to the

coastal region. We note a longitudinal pattern of deformation along several of the transverse strike-slip faults

that we infer is related to this sharp stress transition. We

observe strong expression of these faults on the abyssal

plain, poor expression on the lower slope, and strong

expression on the upper slope and outer shelf. This pattern could result from reduced fault slip across the overpressured and poorly coupled décollement beneath the

lower slope. The decrease of deformation from the plain

to lower slope is also consistent with a lower plate origin.

Rejuvenation of the faults on the upper slope is consistent with progressive dewatering of the wedge and resulting stronger interplate coupling as would be expected

for the rearward part of the wedge. The landward widening of individual fault zones is also consistent with fault

slip transmitted upward through and eastward thickening accretionary wedge [Goldfinger et al., 1996].

Gorda deformation

Several models have been proposed to explain the curvature and fanning of magnetic lineations. These include

a rigid plate rotation model proposed by Riddihough

[1980] that incorporates three subplates corresponding

to the three segments of the Gorda Ridge, driven by faster

spreading at the northern end of the ridge (Model A).

Another incorporates three subplates and accommodates

motion of the faster spreading northern segment with leftlateral slip along northeast trending faults inherited from

the Gorda Ridge (Model B) [Knapp, 1982]. A third incorporates a series of right-lateral faults to accommodate the

faster moving northern block (Model C) [Bolt et al., 1968].

A fourth model uses flexural-slip buckling to account for

the curvature of the magnetic lineations and assumes that

deformation is widely distributed, mostly along pre-existing weaknesses inherited from the ridge (Model D) [Silver, 1971; Stoddard, 1987]. This model also utilizes

principally left-lateral slip on northeasterly striking faults,

and includes a broad northeast-striking shear zone to

obviate the need for subduction or obduction at the

Mendocino transform [Wilson, 1986; 1989].

All of these models can explain, to a first-order, the

pattern of the magnetic lineations. However, each also

has potential problems, many of which were not appar-
The Cascadia Subduction Zone and Related Subduction Systems | 15

ent when they were developed. For example, the first

three do not produce the smoothly curved magnetic

lineations that are observed, although prior to the new

SeaBeam data it was not known whether or not the

smoothness was real or an artifact of the relatively lowresolution magnetic data. This is an important consideration and is further discussed below. Model C is partially

supported by GLORIA sidescan sonar data that suggest

two right-lateral northwest-trending shear zones in the

west-central part of the plate. However, this model is

not in good agreement with seismicity data which indicate that the northeast trending nodal planes are preferred [Silver, 1971; Velasco et al., 1994]. Model D fits

the seismicity data and limited fault mapping, but requires the plate to buckle in a full sinusoid, as for the

case of an end-loaded beam, an unlikely elastic state for

the hot, thin Gorda lithosphere. The absence of the full

waveform is explained by delamination of the lithosphere

and accretion of the missing upper crust to the Mendocino

Ridge [Stoddard, 1987]. Model E fits the seismicity data,

the shape of the magnetic lineations, and the evidence

for a shear zone in the central part of the Gorda plate,

but requires the northern part of the plate to be undeformed and moving entirely with the Juan de Fuca plate.

Although instrumental seismicity suggests that the northern Gorda segment is less seismically active at present,

evidence of active faulting (discussed below) indicates

that this part of the plate is deforming. This model also

does not allow convergence at the Mendocino Ridge.

Recent submersible observations and dated samples from

the Mendocino Ridge, however, suggest that the ridge is

constructed principally of Gorda plate material and that

the ridge has been raised above sea level during possible obduction events. This is indicated by wave cut platforms with rounded cobbles now on the ridge flanks at

~1800 m depth [Duncan et al., 1994; Krause et al.,

1997]. None of the models incorporate dip-slip faulting

that is occurring along both the fabric-parallel faults and

the pseudofaults, nor are most of them consistent with

the faulting we observe in the northern Gorda (except

Stoddard, [1987]).

New constraints from the 1997 SeaBeam survey

The new SeaBeam bathymetric dataset now available

for the Gorda plate provides new constraints on the kinematics of Gorda deformation. Prior to its collection, the

deformation was analyzed using magnetic lineations,

augmented in some cases by analysis of GLORIA regional

sidescan sonar data. A primary observation we can make

from the new data is that the pervasive basement ridges

present in the Gorda plate, and used in some of the

models to constrain kinematics, are smoothly curved and

mostly unbroken by strike-slip faulting crossing these

lineations (with two exceptions). This observation alone

places a strong constraint on models that require significant faulting at angles to them on major structures. Shearing along a multitude of faults, each with slip smaller

than the 100 m resolution of the data, is an unlikely but

possible alternative that is not resolved with SeaBeam

data. In some areas, the basement fabric, while remaining smoothly curved, is also “kinked.” Each of these kinks

is associated with a basement fold pair oriented parallel

to the kink axis. Additionally, basement fabric ridges appear to be drag folded adjacent to pseudofaults that may

have been activated as tectonic structures.

On a finer scale, we observe clear evidence for strikeslip separation along these linear curved ridges. The ridges

vary considerably in height along strike, and the exposed

highs are ubiquitously split and offset in the long direction. Though space permits showing only a few examples, this faulting style is pervasive and can be subdivided

into two domains, sinistral separation in the south, and

dextral separation in the north, along the same linear

features. This surprising result strongly supports a model

that is similar to the flexural-slip model first proposed by

Stoddard in 1987. In effect, the entire Gorda plate appears to be deforming as a huge flexural-slip fold [Yeats,

1986] in the horizontal plane. This model well explains

the left-lateral slip on the southern “limb” and lesser rightlateral slip on the northern “limb” of the asymmetric

syncline. As with a flexural-slip syncline, separation is

reduced near the axis and opposing limbs have opposite

senses of motion. Younger (in this case more westerly)

units have lesser amounts of slip, indicating that this slip

has occurred during the post-accretion history of the plate.

We can calculate approximate minimum slip-rates for

these faults using offsets from the SeaBeam data, where it

can be estimated, and the approximate age of the underlying basement as a maximum age of fault motion. This

method yields slip rates of 0.5–2 mm/yr for the offsets we

have measured thus far. The age of first motion and the

slip history of these faults is unknown, adding further

though unquantifiable uncertainty to these rates.

The flexural-slip model of Stoddard cannot be correct in a dynamic sense. His model required an elastic

flexure and thus must have a complete wavelength. On

this basis, he predicted that since a complete wave is

not observed, the missing part must have been obducted

or subducted along the Mendocino fault. To test this idea,

we did a simple line length balancing of the flexural-slip

buckle. To a first order, we find that line lengths are equal

along basement lineations between the Blanco and

Mendocino faults, indicating little or no area of the Gorda

plate has been differentially obducted or subducted. In

terms of the flexure model, the “missing” material is not

missing since dynamic flexure is not required. Obduction may have occurred but line length balancing suggests relatively consistent obduction and or subduction

along the ridge.
16 | Goldfinger, Dziak and Fox: Offshore structure of the Juan de Fuca plate from marine seismic and sonar studies

References

Baker, G.E., and C.A. Langston, Source parameters of

the 1949 magnitude 7.1 south Puget Sound, Washington, earthquake as determined from long-period

body waves and strong ground motions, J. Geophys.

Res., 77, 1530–1557, 1987.

Bolt, B.A., C. Lomnitz, and T.V. McEvilly, Seismological evidence on the tectonics of central and northern California and the Mendocino Escarpment, Bull. Seismol.

Soc. Am., 58, 1725–1767, 1968.

Byrne, D.E., D.M. Davis, and L.R. Sykes, Loci and maximum size of thrust earthquakes and the mechanics

of the shallow region of subduction zones, Tectonics,

7, 833–857, 1988.

Duncan, R.A., M.R. Fisk, A.G. Carey, D. Lund, L. Douglas, D.S. Wilson, D.C. Krause, R. Stewart, and C.G.

Fox, Origin and emergence of Mendocino Ridge, EOS

Trans. AGU, 75, p. 175, 1994.

Goldfinger, C., L.D. Kulm, R.S. Yeats, B. Appelgate, M.

MacKay, and G.F. Moore, Transverse structural trends

along the Oregon convergent margin: implications

for Cascadia earthquake potential, Geology, 20, 141–

144, 1992.

Goldfinger, C., L.D. Kulm, R.S. Yeats, C. Hummon, G.J.

Huftile, A.R. Niem, and L.C. McNeill, Oblique strikeslip faulting of the Cascadia submarine forearc: the

Daisy Bank fault zone off central Oregon, in Subduction: Top to Bottom, Geophys. Monogr. Ser., Vol. 96,

edited by G.E. Bebout et al., pp. 65–74, American

Geophysical Union, Washington, D.C., 1996.

Goldfinger, C., L.D. Kulm, R.S. Yeats, L.C. McNeill, and

C. Hummon, Oblique strike-slip faulting of the central Cascadia submarine forearc, J. Geophys. Res.,

102, 8217–8243, 1997.

Goldfinger, C., L.D. Kulm, L.C. McNeill, and P. Watts,

Super-scale failure of the southern Oregon Cascadia

margin, Pure and Applied Geophysics, 157, 1189–

1226, 2000.

Knapp, J.S., Seismicity, crustal structure and tectonics near

the northern termination of the San Andreas fault, Ph.D

Thesis, 343 pp., University of Washington, Seattle,

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Christofferson, Mendocino Ridge: vertical tectonics

through Gorda plate transpression, submitted abstract

for the Oceanography Society meeting, May, 1997.

Ludwin, R. S., C.S. Weaver, and R.S. Crosson, Seismicity

of Washington and Oregon, in Neotectonics of North

America, DNAG CSMV-1, edited by D.B. Slemmons,

E.R. Engdahl, D. Blackwell, and D. Schwartz, pp. 77–

98, Geological Society of America, Boulder, Colorado, 1991.

Riddihough, R.P., Gorda plate motions from magnetic

anomaly analysis, Earth Planet. Sci. Lett., 51, 163–

170, 1980.

Scholz, C., and J. Campos, On the mechanism of seismic

decoupling and back arc spreading at subduction

zones, J. Geophys. Res, 100, 22,103–22,115, 1995.

Silver, E.A., Tectonics of the Mendocino triple junction,

Geol. Soc. Am. Bull., 82, 2965–2978, 1971.

Stoddard, P.R., A kinematic model for the evolution of

the Gorda plate, J. Geophys. Res., 92, 11,524–11,532,

1987.

Taber, J. J., Jr., and S.W. Smith, Seismicity and focal mechanisms associated with the subduction of the Juan de

Fuca plate beneath the Olympic Peninsula, Washington, Bull. Seismol. Soc. Am., 75, 237–249, 1985.

Velasco, A.A., C.J. Ammon, and T. Lay, Recent large earthquakes near Cape Mendocino and in the Gorda plate:

Broadband source time functions, fault orientations,

and rupture complexities, J. Geophys. Res., 99, 711–

728, 1994.

Weaver, C.S., and G.E. Baker, Geometry of the Juan de

Fuca plate beneath Washington and northern Oregon

from seismicity, Bull. Seismol. Soc. Am., 78, 264–275,

1988.

Wilson, D.S., A kinematic model for the Gorda deformation zone as a diffuse southern boundary of the Juan

de Fuca plate, J. Geophys. Res., 91, 10,259–10,269,

1986.

Wilson, D.S., Deformation of the So-Called Gorda Plate,

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The Cascadia Subduction Zone and Related Subduction Systems | 17

Cascadia microplate models and within-slab earthquakes

Ray E. Wells1, Richard J. Blakely1 and Craig S. Weaver2

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

2 United States Geological Survey at University of Washington, Box 351650, Seattle, Washington, 98195, USA

rwells@usgs.gov, blakely@usgs.gov, craig@geophys.washington.edu

Introduction

Intermediate-depth earthquakes (30–90 km) in young,

warm slabs like the subducting Juan de Fuca plate may

in part be controlled by dehydration reactions and phase

changes, which in turn are a function of the age of the

plate and convergence rate [Kirby et al., 1996]. Because

the Cascadia convergent margin is caught in the dextral

shear couple between the much larger Pacific and North

American plates, the forearc is rotating clockwise in a

broad deformation zone along the plate boundary. Such

motion of the forearc can affect the convergence rate

and we note that within-slab earthquakes are rare beneath the rotating forearc block. We use a recent kinematic model for Cascadia forearc motion [Wells et al.,

1998; Wells and Simpson, 2001] to calculate long-term

convergence rates with respect to the forearc and discuss possible controls on the distribution of earthquakes

in the subducting plate.

Block model and implications for convergence

Wells et al. [1998] calculated a pole of rotation and

present-day velocity field for the Cascadia forearc by

assuming that paleomagnetically-determined Cenozoic

tectonic rotation of the forearc [e.g., Magill et al., 1982]

continues today and is linked to contemporary northward

migration of the Sierra Nevada block [e.g., Pezzopane and

Weldon, 1993; Walcott, 1993]. In the model, the Cascadia

forearc is migrating northward along the coast and breaking up into large, clockwise rotating blocks (Figure 1).

Deformation occurs mostly around the margins of a large,

relatively aseismic Oregon coastal block (OC), in part

composed of thick, accreted oceanic seamount crust

(compare Figure 1 with Figure 2b). This 400 km-long

block is rotating clockwise with respect to North America

about a rotation pole located in the backarc (OC–NA),

and thus it rotates over the trench at Cape Blanco and

creates an extensional volcanic arc on its trailing edge.

Northward motion of the block breaks western Washington into smaller, seismically active blocks and compresses them against the Canadian Coast Mountains

restraining bend. Arc-parallel transport of forearc blocks

is calculated to be up to 8 mm/yr with respect to stable

North America. Shortening appears to be concentrated

in Washington, where active east–west-trending thrust

faults and folds (e.g. Seattle fault) accommodate north–

south shortening.

Predicted long-term block motions are consistent with

global positioning system (GPS) velocities observed in the

Oregon forearc when elastic deformation due to subduction is taken into account. The GPS velocity field in the

forearc has largely been explained as the sum of two processes [Kahzaradze et al., 1999; Savage et al., 2000;

McCaffrey et al., 2000; Wells and Simpson, 2001]: 1) elastic

deformation above the locked Cascadia subduction zone

(also constrained independently by heat flow, seismic profiling and vertical deformation, Hyndman and Wang,

[1995], Flück et al. [1997]); and 2) the paleomagnetically

and geologically constrained long term clockwise rotation

of an Oregon forearc microplate in response to oblique

subduction, basin–range extension and Pacific–North

America dextral shear [Wells et al., 1998].

Rotation of the Oregon forearc significantly affects

Juan de Fuca plate (JDF) convergence rates (Figure 3,

Table 1). In a North America-fixed reference frame, convergence velocities are highly oblique and decrease from

about 40 mm/yr off Vancouver Island to 28 mm/yr near

the Mendocino triple junction (JDF-NA pole of DeMets

and Dixon, [1999]). However, adding clockwise rotation

of the Oregon forearc increases the convergence rate

south of Cape Blanco (up to 35 mm/yr) and decreases it

northward (down to 29 mm/yr off the mouth of the Columbia River). The northern limit of the “convergence

low” is constrained by the presumed northern limit of

the rigid OC block. Northwest Washington and Vancouver Island are not part of the rotating block, so faster

convergence is calculated there because we assume it

to be kinematically more like North America than coastal

Oregon.

Although there are large uncertainties in the OC–NA

pole, and convergence is quite sensitive to the pole location (compare results of Wells et al., [1998]), a similar

result is obtained from the well-constrained GPS-derived

pole of McCaffrey et al. [2000]). Given these caveats, convergence rates off central and northern Oregon are about

16% lower on average, and more than 25% lower than
18 | Wells, Blakely and Weaver: Cascadia microplate models and within-slab earthquakes

the maximum convergence along the northern and southern Cascadia subduction zone. These variations in convergence could affect earthquake magnitude-recurrence

intervals along the subduction zone and within the slab.

Seismicity and volcanism

Great subduction zone earthquakes have occurred repeatedly along the plate boundary in the recent geologic past [e.g., Atwater and Hemphill-Haley, 1997], but

the subduction zone is presently seismically quiet. Paleoseismic evidence indicates that the entire plate boundary has ruptured in the past, although geodetic evidence

suggests that the width of the plate boundary locked zone

may be much narrower off Oregon. The thick mafic core

of the rotating Oregon block [Tréhu et al., 1994] has the

lowest uplift rate along the coast [Mitchell et al., 1994;

Hyndman and Wang, 1995], and the inferred narrow

locked zone is entirely offshore, apparently following the

western limit of the accreted Siletzia mafic terrane. Outboard of the Siletz terrane, sediments of the Eocene to

Quaternary accretionary wedge and marginal basin complex comprise the geodetically and thermally defined

locked zone along the plate boundary megathrust

[Hyndman and Wang, 1995]. McNeill et al. [1998; 2000]

and Wells et al. [2000] have suggested that the accretionary-marginal basin complex is subdivided by a variety of oblique, upper plate folds and faults that may in

part be related to large-scale seismic segmentation of

the subduction zone. The Oregon segment, which is

characterized by a narrow locked zone and our calculated slower convergence rates, coincides with the region of low crustal and slab seismicity and may indicate

lower strains inboard of the narrow locked zone. Slower

inferred convergence rates (~30 mm/yr) in central

Cascadia are still consistent with 500 year great earthquake recurrence intervals, assuming 15 m average slip

per event is typical, as was calculated from the 1700 AD

tsunami waveforms [Satake and Wang, 2000].

Contemporary seismicity in Cascadia is concentrated

in western Washington, Vancouver Island and northern

California (Figure 2). The crustal earthquakes and active

faults outline a western Oregon region that has very low

rates of seismicity and internal deformation. Although

internal deformation rates are low, paleomagnetic rotation rates are high, consistent with its rotation as a large,

semi-rigid block, as discussed above. Focal mechanisms

in the forearc indicate north–south compression and are

consistent with north–south shortening against the Canadian Coast Mountains re-entrant.

On the trailing edge of the rotating Oregon forearc

block, the extensional Cascade arc accommodates some

of the westward motion through magmatism and normal faulting (Figure 2). The extensional arc ends near

the latitude of the calculated pole of rotation and the

sparse arc volcanoes to the north rest on a folded and

uplifted basement (Figure 2c). Compared to some other

circum-Pacific arcs, Cascade volcanic production rates

are modest [Sherrod and Smith, 1990]. The arc is producing two major end member primitive basalts which

reflect variable input of slab volatile component and

mantle sources: 1) Low potassium olivine tholeiite

(LKOT), which indicates hot dry melting of depleted

upper mantle; and 2) more fluid-rich calc-alkaline basalt (CAB; Bacon et al., [1997]; Conrey et al., [1997]).

LKOT dry melts are associated with rifting and are common in the basin and range and in the extensional arc as

far north as Mount St. Helens [Conrey et al., 1997]. These

magma types and modest volumes are consistent with

subduction of a young, warm slab [Kirby et al., 1996],

but it is worth noting that the largest slab earthquakes

occur north of Mount St. Helens and the LKOT-bearing

OC-SN

pole

FIGURE 1: Velocity field for Oregon forearc microplate calculated

from OC–NA pole (from Wells and Simpson, [2001], modified from

Wells et al., [1998]). Oregon block (OC) rotating at Neogene

paleomagnetic rate is linked to Sierra Nevada (SN) block moving

at vlbi rate by Euler pole (OC–SN) in Klamath Mountains. Extensional arc forms along trailing edge of Oregon forearc block which

absorbs Sierra Nevada displacement by rotating over trench.

North end of Oregon block deforms Washington forearc against

Canadian re-entrant in the margin, causing north–south compression, uplift, thrust faulting and forearc earthquakes. Rates from

very long baseline interferometry (vlbi), paleoseismology (ps), and

magmatic spreading (m).
The Cascadia Subduction Zone and Related Subduction Systems | 19

extensional arc segment. On the other hand, the California rifted arc segment is associated with abundant

Wadati–Benioff zone seismicity.

The distribution of in-slab earthquakes generally follows that of upper plate earthquakes in the forearc; they

are concentrated in the Puget Lowland, are rare beneath

Oregon and become most abundant in the Gorda “plate”

near the Mendocino triple junction (MTJ) (Figure 2). Kirby

et al. [1996] and Rogers [this volume] argue that the

co-location of upper and lower plate earthquakes is the

result of fluid flow into the upper plate from dehydration reactions in the subducting plate. Some of the earthquakes in the lower plate beneath Washington appear

to line up along gravity gradients presumed to be caused

by structures in the upper plate (Figure 4). Although slab

earthquakes and those in the upper plate have similar

distributions, the stress regimes for the two populations

are quite different [Ma et al., 1996]. Earthquakes in western Washington and in the MTJ area have T-axes that lie

within the slab and are consistent with a slab-pull force,

in addition to local stresses due to bending of the slab

and transformation of basalt crust to eclogite. Ma et al.

[1996] show that T-axes follow the arch in the plate beneath Puget Sound, with focal mechanisms that are consistent with bending of the arch. In general, within-slab

earthquakes appear to be concentrated in areas where

strain is concentrated in the bending slab and near the

Mendocino and JDF/Explorer/NA triple junctions.

Why are there few in-slab earthquakes beneath Oregon?

The paucity of intermediate-depth earthquakes beneath

southwest Washington and western Oregon is puzzling

(Figure 4). If dehydration reactions and phase changes

are major constituents in triggering intermediate-depth

earthquakes, one might expect earthquakes beneath

Oregon as well. However, our calculated 20% southward decrease in convergence between 47° and 46°N

coincides with the drop off in within-slab seismicity (Figure 4) but it is not clear if this modest decrease could

affect the seismic behavior of the slab. Perhaps it is warm

enough to promote greater aseismic deformation.

Alternatively, the lack of slab seismicity beneath Oregon may indicate a low stress environment in the

subducting plate. One important difference between

Washington and Oregon is the depth to which the subducting slab can be recognized beneath the continent.

Beneath Washington, the slab can be imaged to a depth

of about 500 km by teleseismic methods, but no deep

slab under Oregon is recognizable below about 150 km

[Rasmussen and Humphreys, 1989]. If no deep slab exists

beneath Oregon, then resulting slab-pull forces would

be low. Lower slab-pull, coupled with lower forces along

the narrow locked zone offshore, may produce lower

stresses in the Juan de Fuca plate beneath the Oregon

FIGURE 2: Cascadia earthquakes, faults, volcanoes and Neogene forearc rotation. A) lower plate seismicity; B) upper plate seismicity,

recent focal mechanisms (Mw>5), and late Cenozoic faults; C) Quaternary arc volcanism (white), major volcanoes (open triangles),

post-five Ma volcanic vents (filled triangles) and post-15 Ma forearc rotations with uncertainties (arrows).
20 | Wells, Blakely and Weaver: Cascadia microplate models and within-slab earthquakes

FIGURE 3: Cascadia convergence velocities with respect to the forearc. North of 47°N, calculated from the JDF–NA pole of DeMets and

Dixon [1999]; south of 47°N, from the JDF–OC pole of Wells and Simpson [2001]. Note that convergence rate is low offshore of Oregon.

Site JDF-NA (DD99) JDF-OC (WS00)

(Lat ° N/Lon ° W) Az ° E rate mm/yr Az ° E rate mm/yr

40/125 72.7 35.0

41/125 45.6 28.7 72.1 34.0

42/125 47.6 29.8 71.6 33.0

43/125 49.4 30.9 71.0 32.0

44/125 51.1 32 70.3 31.0

45/125 52.7 33.1 69.6 29.9

46/125.5 52.1 34.9 69.2 28.8

47/126 53.5 36 68.7 27.6

48/126 54.8 37.2 67.9 26.6

49/127 54.2 38.9 67.7 25.2

50/127 53.5 40.5 66.7 24.2

TABLE 1: Convergence velocities of the Juan de Fuca plate (JDF) with respect to North America (NA) and Oregon Coastal block (OC).

Note: DD 99 is pole of DeMets and Dixon (1999); WS00 is pole of Wells and Simpson, [2001];

bold rates are used in this paper.
The Cascadia Subduction Zone and Related Subduction Systems | 21

FIGURE 4: Earthquakes within the Juan de Fuca plate (M ≥ 2; Yelin et al., [1994], Weaver and Meagher, unpub compilation) plotted on

isostatic residual gravity [Blakely and Jachens, 1990], along with downdip limit of locked zone and transition zone (450°C) from

Oleskevich et al. [1999]. Sixty km depth contour on Wadati–Benioff zone seismicity modified from Ludwin (http://

www.geophys.washington.edu/SEIS/PNSN/INFO_GENERAL/platecontours.html) and Shedlock and Weaver [1991], and including the

most recent M=2.6 slab earthquake 63 km beneath Salem, Oregon (http://www.geophys.washington.edu/SEIS/EQ_Special/

WEBDIR_00123001530o/welcome.html).

forearc and thus few within-slab earthquakes. If the deep

slab is missing beneath Oregon, it implies a condition

that may have existed for a considerable period of time.

Long-term variations in slab-pull and within-slab seismicity

along strike may have implications for deformation history in the forearc that could be explored and tested. If

subsidence of the Puget–Willamette lowland is related to

phase changes and geometry of the slab at depth, then

variations in subsidence history along strike could be used

to infer the long-term distribution of slab earthquakes

[Rogers, this volume; Engebretson and Kirby, 1995]. The

number of within-slab earthquakes and thickness of postmid Miocene Puget–Willamette basin fill both decrease

southward from Puget Sound [Yeats et al., 1996] and terminate at Eugene. This is approximately at the same latitude as the northward termination of the basin and range

and suggests a possible long-term relationship between

back arc extension, the rotating forearc and the lack of

slab seismicity beneath Oregon.

Summary

Within-slab earthquakes in Cascadia are generally consistent with tectonic stresses expected from slab-pull,

basalt–eclogite phase changes, triple junction deforma-
22 | Wells, Blakely and Weaver: Cascadia microplate models and within-slab earthquakes

tion and slab bending beneath the Olympic arch. The

lack of within-slab earthquakes beneath Oregon coincides with the rotating Oregon forearc, which is inferred

to reduce the subduction rate offshore by 15–25%. This

may advance dehydration rea ctions and reduce slab

seismicity, although the lack of slab-pull beneath Oregon

is also a possible cause. Dehydration reactions and phase

changes may be important constituents in causing intermediate-depth earthquakes, but the overall pattern of

within-slab earthquakes in Cascadia suggests that tectonic stresses may also be a necessary condition of their

occurrence. Additional work is needed to addresss the

possible relationship between lowland physiography,

within-slab earthquakes, rotation of western Oregon and

the inferred lack of a deep slab beneath Oregon.

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and co-seismic slip in great subduction zone earthquakes, Supplement to EOS, Transactions of the American Geophysical Union, 81, p. 907, 2000.

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Washington and Oregon earthquake history and hazards, 11 pp., U. S. Geological Survey Open-File Report 94–226B, 1994.
24 |
The Cascadia Subduction Zone and Related Subduction Systems | 25

Geometry of the subducting Juan de Fuca plate:

New constraints from SHIPS98

Anne M. Tréhu1, Tom M. Brocher2, Kenneth C. Creager3, Mike A. Fisher2, Leiph A. Preston3, George Spence4

 and the SHIPS98 Working Group

1 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97331–5503, USA

2 United States Geological Survey, 345 Middlefield Road, Menlo Park, Caifornia, 94025, USA

3 Earth and Space Sciences, Box 351650, University of Washington, Seattle, Washington, 98195, USA

4 Earth and Ocean Sciences, University of Victoria, PO Box 3055, Victoria, British Columbia, V8W 3P6, Canada

trehu@oce.orst.edu, brocher@usgs.gov, kcc@geophys.washington.edu, mfisher@usgs.gov,

prestola@u.washington.edu, gspence@uvic.ca

ABSTRACT

We have processed seismic reflection profiles acquired in

the Strait of Juan de Fuca, inverted travel times of first

arrivals from onshore recordings of the offshore shots to

determine the velocity of the upper crust in this region

and inverted the travel times of first and secondary arrivals that undershoot the central core of the Olympic Mountains. The seismic reflection profiles show a pattern of

crustal reflectivity similar to that recorded beneath Vancouver Island. Inversion of first arrivals, interpreted to be

diving waves through the upper and mid-crust indicate a

5–7 km-deep linear, northwest-trending basin beneath

the southwestern shore of the Strait and uplift of the basement Crescent terrane rocks beneath the northwestern

shore. Velocities beneath the central Olympic Peninsula

at 15–25 km depth show strong lateral variation, with

higher velocities underlying the surface exposures of the

Crescent terrane and lower velocities at these depths beneath the Olympic core rocks. Inversion of secondary seismic arrivals interpreted to be reflections from the base of

the crust of the subducted Juan de Fuca plate indicate

that the Moho is at a depth of about 34 km beneath the

western Strait of Juan de Fuca and dips about 7o to the

east, reaching a depth of about 46 km beneath the eastern boundary of Olympic National Park. Correlating these

results with results of previous experiments on Vancouver Island and in southwestern Washington confirm earlier interpretations of an arch in the subducted slab

beneath the Olympic peninsula but indicate that the arch

is asymmetric and less pronounced than previously

thought. We attribute both the northwest-trending folding of the Crescent terrane and the asymmetry in the shape

of the subducting slab to resistance to the northward

motion of the Paleocene-age Cascadia forearc terrane by

the thick lithosphere of the pre-Tertiary terranes of British

Columbia.

Introduction

In March 1998, as part of the USGS-sponsored SHIPS

(Seismic Hazards Investigations of Puget Sound) project

[Fisher et al., 1999; Brocher et al., 1999], the R/V

Thompson towed a 2.4 km-long, 96 channel digital seismic streamer and a 13 gun, 79.3 liter (4838 in3

) airgun

array through the waterways of the Puget Basin, Strait of

Juan de Fuca and Strait of Georgia (Figure 1). During the

early part of the survey (lines 1–4), before deploying the

streamer, it was able to fire a 16 airgun, 110.3 liter (6730

in3) array at a 40 second interval. The airgun shots were

also recorded on 273 REFTEK seismometers (Figure 1)

and on 15 ocean-bottom seismometers. Because they

were far from sources of cultural noise, many stations on

the Olympic Peninsula recorded clear direct arrivals (Pg)

and secondary arrivals interpreted to be reflections from

the base of the crust of the Juan de Fuca plate (PmP)

from shots in the Strait of Juan de Fuca, Puget Sound

and the Hood Canal (Figure 2). Using these data, we

have developed a preliminary model for the shape of

the subducting plate beneath the Olympic Peninsula and

for the overlying crust beneath the Strait of Juan de Fuca

and Olympic core.

Multi-channel seismic data

The multi-channel seismic data from line 7 (labeled JDF1

and JDF2 in Figure 1) were processed through brute stack

stage (i.e. the data were edited to remove bad traces,

sorted, band pass filtered and stacked using a simple

velocity function). The primary objective of this analysis

was to determine the pattern of lower crustal reflectivity.

Because of the relatively low frequency of arrivals from

the lower crust and the insensitivity of these deep arrivals to velocity for the limited aperture of the streamer,

this simple analysis is adequate to meet our primary objective of imaging lower crustal reflectivity. More detailed
26 | Tréhu et al.: Geometry of the subducting Juan de Fuca plate

velocity and static analysis is needed to determine shallow crustal structure. The results of the brute stack for

JDF1 are shown in Figure 3a. On both profiles, a subhorizontal, two-second wide band of high reflectivity is

observed at a two-way travel time of 7–9 seconds beneath the Strait west of ~123.6oW. Adopting the

LITHOPROBE terminology, these reflections are labeled

“E” in Figure 3. Beneath this band of reflections, a second reflection is intermittently observed, corresponding

to the LITHOPROBE “F” reflection (Figure 3b). East of

123.6 ̊W, the “E” reflections dip east. The wide-angle

analysis discussed below indicates that the “F” reflection

comes from a boundary 3–6 km above the Moho of the

subducted Juan de Fuca plate crust. This boundary may

be the top of the subducted oceanic crust, as previously

interpreted for the LITHOPROBE data [Hyndman et al.,

1990]. The geologic nature of the “E” reflection zone

remains controversial [Calvert, 1996].

Onshore/offshore REFTEK/airgun data

To image the geometry of the subducted plate beneath

the core of the Olympic Mountains, we first correct the

travel time observations for effects of upper crustal structure beneath Puget Sound, the Hood Canal and the Strait

of Juan de Fuca by replacing the upper ten km beneath

each shot by material with a velocity of 5.5 km/s. We

then invert the corrected data for the three-dimensional

velocity structure and Moho position beneath the Olympic Peninsula. This two-step approach permits us to use a

small grid spacing to model the large effects on observed

travel time of shallow basin structure and a relatively sparse

grid to model the velocity and Moho structure beneath

the Olympic Peninsula. The corrections in step 1 are based

on a three-dimensional model for Puget Sound from

Brocher et al. [2001] and our new model, presented below, for the Strait of Juan de Fuca.

Inversion of first arrivals for velocity structure

beneath the Strait of Juan de Fuca

First arrivals from shots on line 4 (coincident with line 7,

but shot with a larger volume airgun) recorded on stations deployed along the northern Olympic Peninsula

were inverted to obtain velocities in a volume with one

kilometer grid spacing using code developed by John Hole

[Hole, 1992]. Several horizontal slices through the resulting model are shown in Figure 4. Only model cells constrained by data are shown. The starting model was a

one-dimensional model with velocity increasing with

FIGURE 1: Topographic map of the Olympic Peninsula and Puget Basin showing the locations of mulit-channel seismic profiles (lines),

temporary REFTEK seismometers (squares) and University of Washington network stations (triangles). Lines and REFTEKs used to date in

our analysis are shown in black. The boxes show the regions included in the first arrival and wide-angle tomography presented here. These

regions overlap and complement analyses done by other participants in SHIPS. Stations for which data are shown in Figure 2 are labeled.
The Cascadia Subduction Zone and Related Subduction Systems | 27

FIGURE 2: A) Data from JDF2 recorded at station 9022. Note wide-angle reflection interpreted as PmP, shallow reflections at nearer

offset (PiP) and a 0.5 second travel time delay due to the Clallum basin. B) Examples of record sections recorded on the northern

Olympic Peninsula from shots in Puget Sound and the Hood Canal. Note the strong amplitude of secondary arrivals. Comparison of the

travel time and offset of these arrivals with the arrival interpreted as PmP in Figure A suggests that these are also PmP.

FIGURE 3: Multi-channel seismic line JDF1 and LITHOPROBE line 2. "E" and "F" reflections on JDF1 are interpreted to be the same as

the similarly named events in the LITHOPROBE data. The location of LITHOPROBE line 2 is shown in Figure 7 (LP2). Figure is from the

on-line LITHOPROBE atlas (http://www.litho.ucalgary.ca/atlas/atlas.html). Velocity modeling suggests that 6.5 is a reasonable velocity

for converting two-way travel time to depth. Correlation with the Moho depths shown in Figure 6 suggests that the "F" reflection is ~6

kilometers above Moho, suggesting that it is the top of the crust of the subducting plate and probably the plate boundary.
28 | Tréhu et al.: Geometry of the subducting Juan de Fuca plate

depth and a root-mean-squared (rms) misfit between

observed and calculated travel times of 0.56 seconds.

The model converged smoothly in 11 iterations to a model

with an rms misfit of 0.08 seconds, similar to the picking

uncertainty. Tests with several significantly different initial models indicated that the final model is not sensitive

to the starting model. Because of the high velocities at

relatively shallow depth, few rays penetrate beneath 15-

km depth, even with offsets as great as 110 km.

East of km 40, the model of Brocher et al. [2001] has

been “patched” into our model because it includes data

from shots along short profiles in the eastern Strait of

Juan de Fuca (Figure 1). The good match between the

two models where they overlap validates the SHIPS

working group approach of having various groups workFIGURE 4: A) Three-dimensional inversion of first arrivals from shots in the Strait of Juan de Fuca recorded on stations along the

northern Olympic Peninsula. Rms travel time misfit of the initial one-dimensional velocity model was 0.49 seconds; misfit was decreased to 0.08 seconds after 11 iterations. B) Velocity in the upper ten km along each line of shots used in the three-dimensional

Olympic Peninsula lower crust and PmP inversions. For the Strait of Juan de Fuca, the model of 4a was used. For the Hood Canal and

Puget Sound, the model of Brocher et al. [2001] was used. The vertical travel time beneath each shot to a depth of ten km is also

shown. C) Velocity scale for 4a and 4b.

ing on different subsets of the data to address specific

scientific problems. Eventually, all data will be merged

to generate a regional “super-model”. Figure 4b shows

the vertical velocity in the upper ten km beneath each

shot for the Strait of Juan de Fuca (line 4), Hood Canal

(line 3) and Puget Sound (line 9), and the corresponding

vertical travel time through the model. These times were

used to correct the observed travel times to replace the

upper ten km beneath the shots with material with a

constant velocity of 5.5 km/s.

Low velocities in the upper crust along JDF2 west of

km 100 indicate the presence of a linear, northwesttrending basin that is not present beneath the northern

line. This basin is known as the Clallam basin where it

extends on land on the Olympic Peninsula. We interpret
The Cascadia Subduction Zone and Related Subduction Systems | 29

this basin to be the result of warping of the Siletz/Crescent terrane along a west–northwest-trending axis, perhaps in response to compression generated by northward

migration of the Tertiary forearc relative to the pre-Tertiary terranes of southwestern Canada. The basin beneath

the eastern Strait of Juan de Fuca, known as the Port

Townsend basin, is discussed by Brocher et al. [2001].

Inversion of first arrivals for mid-crustal

velocity structure beneath the Olympics

Because the first arrivals recorded from shots in the Hood

Canal and Puget Sound, corrected for the model of Figure 4, show clear evidence for additional lateral variation in velocity structure, we inverted these arrivals for

the velocity structure of the mid-crust beneath the Olympic Mountains using a five kilometer grid within the

box outlined in Figure 1 for the PmP inversion. The results are shown in Figure 5. The 1 km/s lateral variation

in velocity at mid-crustal depth is similar to that obtained

by Crosson et al. [this volume] at this depth in a more

comprehensive regional model based on earthquake and

SHIPS data (but not yet including the source/receiver

pairs used here). Because it is based on data from the

same shot/receiver pairs for which PmP arrivals that undershoot the Olympics were observed, the model of Figure 5 provides a useful starting model for the Moho

interface inversion while waiting until the final “master

model” incorporating all SHIPS is complete. The effect

of ignoring this mid-crustal lateral velocity variation is

discussed below.

Inversion of PmP wide-angle reflection times

for Moho depth beneath the Olympics

Figure 6 shows our working three-dimensional model

for the Moho of the subducting Juan de Fuca plate, determined using the interface inversion technique of Zelt

et al. [1996] as provided by J. Hole. Reflection times

were calculated by ray tracing through the model of Figure 5 after application of the correction for shallow structure discussed above.

The model indicates that the Moho is at 34–36 km

depth beneath the western Strait of Juan de Fuca, dipping ~7o to the east–southeast. Because the “F” reflection on lines JDF1 and JDF2 falls at a depth of ~30 km,

this result supports the LITHOPROBE interpretation that

the “F” reflection is not the Moho and may be the top of

the subducted ocean crust (and probably the plate

boundary). The apparent east–southeast dip of the Moho

is not inconsistent with the apparently sub-horizontal “E”

and “F” reflections beneath the western Strait of Juan de

Fuca since the portion of the subducted plate imaged by

the common-midpoint data and that imaged by the PmP

wide-angle reflections are not coincident. Moreover, the

LITHOPROBE data indicate that the “F” reflection generally dips somewhat more steeply than the better-deFIGURE 5: Inversion of first arrivals from shots in the Hood Canal

and Puget Sound (small black dots) recorded at sites along the

northern Olympic Peninsula (black triangles). Initial model had

velocity 5.5 km/s from -5 to 10 km depth, 6.5 km/s at 15 km, and

7.0 km/s for depth >20 km. The model converged after nine iterations. Rms travel time misfit of the initial model was 1.04 seconds, decreasing to 0.16 seconds for the model shown here.

Contours are shown at 0.1 km/s intervals. Velocities in each layer

are only constrained within the thick gray line. Note the pronounced lateral contrast in velocity at 15 and 20 km depth, where

velocities beneath the Crescent terrane are 1 km/s greater than

those beneath the Olympic core. This is consistent with the

anomaly at this depth obtained by Crosson et al. [2002] using

earthquake data and a different subset of the SHIPS data. With

the smoothing parameters used for this inversion, the anomaly

extends to 30 km, beneath which the velocity is constant at 7.0

km/s.
30 | Tréhu et al.: Geometry of the subducting Juan de Fuca plate

fined “E” reflection. Our model for the depth and dip of

Moho is also generally consistent with the model of

Creager et al. [this volume] derived using a different

modeling approach and a different subset of the data.

The shape of the Moho obtained via inversion of PmP

travel times is quite sensitive to the starting model.

Ignoring the lateral velocity variation shown in Figure 5,

and assuming a mid-crustal velocity that increases only

with depth, leads to a solution that has considerable

curvature and an rms misfit of 0.36 seconds.

The inversion results are also quite sensitive to which

secondary arrivals are included in the solution. Initially,

we included all secondary arrivals observed from shots

in the Hood Canal into stations on the northern Olympic Peninsula. Including these arrivals, however, results

in a best fit solution that also has considerable curvature

and a large rms misfit of 0.56 seconds. Re-examination

of the data indicated that some arrival time picks from

shots in the Hood Canal to stations near Port Angeles

were incompatible with picks from these shots to stations further west. The origin of these incompatible arrivals remains enigmatic. Unlike other SHIPS observations

reported by Creager [2002], the travel times of these particular arrivals do not appear to be consistent with a simple planar interface parallel to and shallower than the

Moho, suggesting additional, more complex lateral heterogeneity in the lower crust beneath the Olympics.

Comparison to other data

Figure 7 shows the results of the Moho model, converted

to depth to the plate boundary by assuming a constant

thickness of the Juan de Fuca plate crust of six kilometers

and assuming that the plate boundary is at the top of the

Juan de Fuca plate crust. Normalization to either the plate

boundary or the base of the crust of the subducting plate

is necessary for comparing reflection data (LITHOPROBE

data from Vancouver Island), which are more sensitive

to the former, to wide-aperture data [Parsons et al., 1998;

this study], which are more sensitive to the latter. Extension of the contours south of the Parsons et al. [1998]

profile is constrained by the results of Trehu et al. [1994]

and Gerdom et al. [2000]. This comparison confirms that

the dip of the subducting plate is relatively shallow beneath the Olympic Peninsula. The dip of 7–8o beneath

the Olympics is similar to the dip of the subducting plate

beneath the continental shelf throughout the Cascadia,

but significantly less than the dip of 11–15o determined

beneath the Coast Ranges to the south from northern

California to southwestern Washington or the dip of ~20o

determined beneath Vancouver Island [e.g. Hyndman

et al., 1990; Trehu et al., 1994; 1995; Parsons et al.,

1998; Clowes, this volume; Gerdom et al., 2000. These

results support earlier inferences of an arch beneath the

Olympics but suggest that the arch is less pronounced

than previously thought and that it is asymmetric. We

speculate that both the steeper dip of the plate beneath

Vancouver Island and the northwest-trending folding of

the Crescent terrane beneath the Strait of Juan de Fuca

may be due to the interaction between the subducting

plate and the thicker, stronger lithosphere beneath the

pre-Tertiary terranes of British Columbia.

Figure 7 also shows the epicenters of large, damaging lower plate earthquakes in this region, including the

January, 2001 Nisqually event. All of these events appear to have hypocenters 6–8 km below the inferred

plate boundary, or 0–2 km below the Moho of the subducting oceanic plate. Given uncertainties in hypocenteral parameters and in the velocity model, we cannot

determine conclusively whether these events initiated

in the uppermost mantle or at the base of the crust of

the subducted plate. We expect that the answer to this

question, which is important for understanding the physical processes causing these events, will be obtained

FIGURE 6: Inversion of PmP arrival times to determine the depth

to the Moho of the subducted Juan de Fuca plate. Travel times,

corrected as described in Figure 4, were calculated through the

velocity model of Figure 5. Upper figure shows the depth to Moho

after five iterations. Rms travel time misfit is 0.30 seconds, compared to an rms travel time misfit of 0.86 seconds for the initial

model, which was a horizontal surface at 45 kilometers depth.

Lower panel shows the observed/calculated travel times. Positive and negative residuals are distributed over the entire surface.
The Cascadia Subduction Zone and Related Subduction Systems | 31

through more detailed analysis of the Nisqually event.

Acknowledgements

We thank the many individuals who helped survey station sites and deploy instruments. We also thank the many

landowners that graciously offered us access to their land.

This work was funded by the United States Geological

Survey through grants 1434HQ98GR00041 and

99HQGR0071 to Oregon State University.

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FIGURE 7: Contours of depth to the top of the subducting plate, assuming that the plate boundary is six kilometers above Moho.

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southwestern Washington. The results confirm earlier reports of an arch in the subducting plate beneath the Olympic Peninsula, but

suggest that the arch is asymetric. Also shown as red squares are the epicenters of historic, damaging lower plate earthquakes,

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the core of the Cascadia forearc complex, in The

Cascadia Subduction Zone and Related Subduction

Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, p. 33, U.S. Geological Survey Open-File

Report 02–328, Geological Survey of Canada Open

File 4350, 2002.

Fisher, M.A., T.M. Brocher, R.D. Hyndman, A. M. Tréhu,

C.S. Weaver, K.C. Creager, R.S. Crosson, T. Parsons,

A.K. Cooper, D. Mosher, G. Spence, B.C. Zelt, P.T.

Hammer, U. ten Brink, T.L. Pratt, K.C. Miller, J.R.

Childs, G.R. Cochrane, S. Chopra, and R. Walia, Seismic survey probes urban earthquake hazards in Pacific Northwest, EOS, Trans. Am. Geophys. Un., 80,

13–17, 1999.

Gerdom, M., A.M. Tréhu, E.R. Flueh, and D. Klaeschen,

The continental margin off Oregon from seismic investigations, Tectonophysics, 329, 79–97, 2000.

Hole, J.S., Nonlinear, high-resolution, three-dimensional

seismic travel time tomography, J. Geophys. Res., 97,

6553–6562, 1992.

Hyndman, R.D., C.J. Yorath, R.M. Clowes, and E.E. Davis,

The northern Cascadia subduction zone at Vancouver Island: seismic structure and tectonic history, Can.

J. Earth Sci., 27, 313–329, 1990.

Parsons, T., A.M. Tréhu, J.H. Luetgert, F. Kilbride, K. Miller,

R.E. Wells, M.A. Fisher, N.I. Christensen, U. ten Brink,

E. Flueh, and W.D. Mooney, A new view into the

Cascadia subduction zone and volcanic arc: implications for earthquake hazards along the Washington margin, Geology, 26, 199–202, 1998.

Tréhu, A.M., I. Asudeh, T.M. Brocher, J. Luetgert, W.D.

Mooney, J.L. Nabelek, and Y. Nakamura, Crustal architecture of the Cascadia forearc, Science, 265, 237–

243, 1994.

Tréhu, A.M., and the Mendocino Working Group, Pulling the rug out from under California: seismic images of the Mendocino triple junction region, EOS,

Trans. Am. Geophys. Un, 76, 369–381, 1995.

Zelt, B.C., R.M. Ellis, R.M. Clowes, and J.A. Hole, Inversion of three-dimensional wide-angle seismic data

from the southwestern Canadian Cordillera, J.

Geophys. Res., 101, 8503–8530, 1996.
The Cascadia Subduction Zone and Related Subduction Systems | 33

Three-dimensional structure of the Cascadia forearc region

from SHIPS active experiment and earthquake observations:

Tomographic inversion provides a high-resolution view

into the core of the Cascadia forearc complex

Robert S. Crosson1, Neill P. Symons2, Kenneth C. Creager1, Leiph A. Preston1, T. Van Wagoner1,

Thomas M. Brocher3, Michael A. Fisher3 and the SHIPS98 Working Group

1 Department of Earth and Space Sciences, University of Washington, Seattle, Washington, 98195–1650, USA

2 Sandia National Laboratories, Box 58000, Albuquerque, New Mexico, 87111, USA

3 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

crosson@u.washington.edu, npsymon@sandia.gov, kcc@geophys.washington.edu, prestola@u.washington.edu,

tvan@geophys.washington.edu, brocher@andreas.wr.usgs.gov, mfisher@octopus.wr.usgs.gov

Phase I of SHIPS (Seismic Hazards in Puget Sound) was a

large-scale airgun deployment in the inland waterways of

western Washington State and western British Columbia.

The coincidence of SHIPS travel time data in the Puget

Sound region with earthquake arrival time data from the

Pacific Northwest Seismograph Network has allowed us

to carry out high-resolution joint active source and earthquake travel time inversion (tomography) for P wave

velocity structure to depths of approximately 60 km over

a portion of the Cascadia forearc margin. A total of about

70,000 P wave arrival observations were used in our

initial effort; approximately 1000 earthquakes comprise

45% of the total number of P wave observations. Excellent constraint of the shallow structure by SHIPS data allows greatly improved resolution of the deep structure

through earthquake observations, compared to inversion

using earthquake data alone. The three-dimensional

velocity model fits both airgun and earthquake picks to

approximately 0.1 second root-mean-squared (rms) overall. Simple variance reduction from a best fitting onedimensional model to our preferred three-dimensional

model exceeds 90%.

After inversion, earthquake hypocenters migrate in

general to slightly shallower depths relative to depths

computed with a best fitting one-dimensional starting

model. Increasing the average velocity of the one-dimensional model could bring the one-dimensional and threedimensional depths into closer agreement; however,

because of very significant three-dimensional effects, such

a model may still not provide a satisfactory one-dimensional average earth model. Earthquakes in the Wadati–

Benioff zone are moved an average of about three

kilometers shallower but generally maintain the average

dip of the zone (about 14°). The scatter of Wadati–Benioff

zone earthquakes about a plane is also reduced by a

factor of about two in standard deviation in going from

one-dimensional to three-dimensional models.

Mafic Crescent/Siletz volcanic rocks form the basement of the Puget basin region with P wave velocity

exceeding 7 km/s at 20 km depth. In the central Puget

basin, Crescent/Siletz is in turn underlain at a depth of

about 30 km by lenses of lower-velocity, subducted

accretionary prism rocks which are gradational with rocks

of the Olympic core further to the west. This suggests

that the Olympic core rocks are mechanically emplaced

by subduction beneath the Crescent/Siletz eastward of

the Olympic Mountains, producing a local low velocity

zone in the deep crust beneath the Puget basin (although

in the conventional sense “crust” may be ill-defined in

this region). The P wave velocity reversal produced by

this emplacement process is typically about 0.5 km/s.

The relationship between Olympic core rocks and the

Crescent/Siletz terrane is consistent with models of

“unroofing” of the Olympic core [e.g., Brandon et al.,

1998].

References

Brandon, M.T., M.K. Roden-Tice, and J.I. Garver, Late

Cenozoic exhumation of the Cascadia accretionary

wedge in the Olympic Mountains, northwest Washington State, GAS Bulletin, 110, 985–1009, 1998.
34 |
The Cascadia Subduction Zone and Related Subduction Systems | 35

Velocity structure of western Washington

from local earthquake tomography

Antonio Villaseñor, D. Stanley and Harley Benz

United States Geological Survey, MS 966, Denver Federal Center, Denver, Colorado, 80225, USA

antonio@gldvisitor.cr.usgs.gov, benz@usgs.gov

Accurate assessment of seismic hazards in Washington,

especially the heavily populated Puget Sound region,

requires insights into the complex crustal and upper

mantle structure and associated seismicity. Numerous geophysical studies have used seismic refraction and reflection data [Tréhu et al., 1995; Johnson et al., 1996], seismic

tomography [Michaelson and Weaver, 1986; Rasmussen

and Humphreys, 1988; Bostock and VanDecar, 1995],

magnetotelluric soundings [Stanley et al., 1992], and gravity and magnetic modeling [Finn, 1990] to understand

crust and upper mantle development in the Pacific Northwest due to subduction of the Juan de Fuca plate beneath the western margin of North America. The geometry

of the subducting Juan de Fuca plate has been demonstrated to be complex [Crosson and Owen, 1987] and

better definition of this geometry is a prerequisite for understanding tectonic interactions between the subducting plate and the overlying continental plate and intraslab

earthquakes. The combination of an accurate three-dimensional velocity model for the crust and upper mantle

beneath Washington combined with other geophysical

data such as gravity, magnetic, and electromagnetic sounding data, provides new constraints on the lithology and

distribution of key geologic units that control seismic energy release. In this talk, we describe new three-dimensional P- and S-wave velocity models of the crust and upper

mantle beneath Washington and use the results to interpret the structural framework under which earthquakes

occur in western Washington.

The three-dimensional P- and S-wave velocity models interpreted in this study were determined by inversion of local earthquake arrival time data from the Pacific

Northwest Seismic Network (PNSN) earthquake catalog

using a technique developed by Benz et al. [1996]. Travel

times are calculated by extrapolation of the Eikonal equation in a finite difference grid. The three-dimensional

velocity structure is determined by solving for perturbation to the velocity model using a linear inverse formulation that is solved with a large, sparse matrix solver (LSQR)

with smoothing constraints. The study region is a 280 km

east–west by 344 km north–south area centered on Puget

Sound, which is subdivided into 8 x 8 x 2 km (2-kmvertical) constant velocity cells. In addition, the model

extends from four kilometers above sea level to 80 km

depth. A finer 1 x 1 x 1 km grid is used to calculate the

travel times with errors less than ± 0.05 seconds. Velocities in the finer grid are determined by interpolation from

the coarser inversion grid. From 36,565 earthquakes in

the PNSN catalog in the time interval from 1970–1996,

3431 well-recorded events were selected for the inversion process. This resulted in 61,906 source-station pairs

to 87 seismic stations in western Washington. These events

also provided 16,972 S-wave arrival times that were inverted to obtain a three-dimensional S-wave velocity

model using the same grid as the P-wave inversion.

Results show a complicated arch structure under the

Olympic Peninsula and Puget Sound region, and highvelocity mafic complexes in the middle to lower crust in

western Washington that appear to control the distribution of crustal seismicity. The main mafic rock unit imaged

in the crust is the Eocene-age Crescent Formation, a sequence of basalt flows that forms the basement beneath

much of western Washington. In addition, a deeper highvelocity section of the mid-crust under Puget Sound is

interpreted as either imbricated, pre-Eocene oceanic

crust, or an advancing mantle wedge that generated the

Crescent Formation. The Juan de Fuca plate may be partially coupled to the overlying, rigid mafic units. Thick,

low-velocity accretionary sedimentary rocks in the Olympic Peninsula contrast with the high-velocity Crescent

Formation units, which form the backstop for tectonic

thickening of the sedimentary rocks. Northwest-trending,

strike-slip zones in the Cascades volcanic belt and stepovers between these slip zones appear to also be controlled by the geometry of the Crescent Formation. In

general, relatively slow lower-crustal velocities occur beneath the Cascades, possibly due to thermal effects and/

or a depleted crust.

High-resolution velocity models of the crust and

subducting Juan de Fuca plate in western Washington

provide detailed images of the subducting plate and several important crustal features. A previously interpreted

arch in the Juan de Fuca plate has been mapped in more

detail. Mafic blocks composed of Crescent Formation
36 | Villaseñor, Harvey and Benz: Velocity structure of western Washington from local earthquake tomography

and possibly an underlying section of pre-Eocene mafic

rocks, stand out in the velocity model, especially in contrast to thick low-velocity sedimentary rocks in the

Olympic Peninsula and in the Seattle basin. The occurrence of a thick mafic wedge above the subducting Juan

de Fuca plate may control seismicity patterns in the Puget

Sound region due to enhanced stress transfer through

the rigid mafic rocks. Regions of high stress within the

mafic rocks may be detectable with Vp/Vs ratios, although

further analysis and detailed modeling is required. We

interpret crustal blocks as being forced northeast around

the arch in the Juan de Fuca plate. Northeast shear stresses

from the subduction direction are expressed in crustal

focal mechanisms in southwest Washington. The development of the velocity model has allowed us to accurately relocate most of the earthquake database for

western Washington, with the result that seismicity patterns can be more completely understood.

References

Benz, H.M., B.A. Chouet, P.B. Dawson, J.C. Lahr, R.A.

Page, and J.A. Hole, Three-dimensional P and S wave

velocity structure of Redoubt Volcano, Alaska, J.

Geophys. Res., 101, 8111–8128, 1996.

Bostock, M.G., and J.C. VanDecar, Upper mantle structure of the northern Cascadia subduction zone, Can.

J. Earth Sci., 32, 1–12, 1994.

Crosson, R.S., and T.J. Owens, Slab geometry of the

Cascadia subduction zone beneath Washington from

earthquake hypocenters and teleseismic converted

waves, Geophys. Res. Lett., 14, 824–827, 1987.

Finn, C., Geophysical constraints on Washington convergent margin structure, J. Geophys. Res., 95, B12,

19,533–19,546, 1990.

Johnson, S.Y., C.J. Potter, J.M. Armentrout, J.J. Miller, C.A.

Finn, and C.S. Weaver, The southern Whidbey Island

fault and active structure in the Puget lowland, Washington, Geol. Soc. Am. Bull., 108, 334–354, 1996.

Michaelson, C.A., and C.S. Weaver, Upper mantle-structure from teleseismic P arrivals in Washington and

northern Oregon, J. Geophys. Res., 91, 2077–2094,

1986.

Rasmussen, J., and G. Humphreys, Tomographic image

of the Juan de Fuca plate beneath Washington and

western Oregon using teleseismic P-wave travel times,

Geophys. Res. Lett., 15, 1417–1420, 1988.

Stanley, W.D., W.J. William, G. Latham, and K. Westhusing,

The southern Washington Cascades conductor—a previously unrecognized thick sedimentary sequence? Am.

Assoc. Petrol. Geol. Bull., 76, 1569–1585, 1992.

Tréhu, A.M., G. Lin, Maxwell, and C. Goldfinger, A seismic reflection profile across the Cascadia subduction

zone offshore central Oregon: new constraints on

methane distribution and crustal structure, J. Geophys.

Res., 100, 15,101–15,116, 1995.
The Cascadia Subduction Zone and Related Subduction Systems | 37

Three-dimensional reflection image

of the subducting Juan de Fuca plate

Kenneth C. Creager1, Leiph A. Preston1, Robert S. Crosson1, T. Van Wagoner1,

Anne M. Tréhu2 and the SHIPS98 Working Group

1 Department of Earth and Space Sciences, University of Washington, Seattle, Washington, 98195–1650, USA

2 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97330, USA

kcc@ess.washington.edu, prestola@u.washington.edu, bob@geophys.washington.edu,

tvan@geophys.washington.edu, trehu@oce.orst.edu

The dynamics and environment under which intraslab

earthquakes occur have been debated during the past

several decades. Earthquake nucleation dynamics proposed for the shallow crust, involving brittle fracture when

shear stresses overcome normal stresses, are difficult to

extend to the high lithostatic pressures encountered in

subduction zones. A process known as transformational

faulting has been invoked to explain catastrophic faulting of ice at high confining pressure as ice undergoes a

phase transition to ice II [Kirby, 1987]. The olivine to spinel

phase transition exhibits similar behavior in the laboratory, offering a reasonable explanation for deep earthquakes [e.g., Green and Burnley, 1989; Houston and

Green, 1995; Kirby, 1995; Kirby et al., 1991]. An important phase transformation in the shallowest 100 km of

subduction zones is the transition of the basalt in the

subducting oceanic crust to eclogite. Dehydration embrittlement, faulting associated with the release of fluids as a

result of this phase transition, provides an especially attractive explanation for these earthquakes [Kirby, 1995].

The depth range over which this reaction proceeds is

strongly temperature dependent because the equilibrium

mineral assemblages depend strongly on the temperature [Peacock and Wang, 1999] and because the transition may be thermally inhibited in a cold, fast moving

slab [Kirby, 1995]. For example, the oceanic crust of the

Pacific plate subducting under northeast Japan is very old

and fast moving. Earthquakes apparently in the subducted

crust occur to depths as great as 100 km, as well as in a

more active zone clearly within the mantle part of the

subducted plate that extends deeper. In contrast, the

Philippine Sea plate subducting below southwest Japan

is young and slow, and contains shallow Wadati–Benioff

seismicity down to only 60 km. Geotherms calculated by

thermal modeling of these two subduction zones can

explain the first order differences in observed seismicity

in terms of basalt to eclogite phase transitions in equilibrium [Peacock and Wang, 1999].

An obvious test of the hypothesis that earthquakes

are associated with the basalt to eclogite phase transition is to determine whether the earthquakes occur within

the subducted crust or the subducted mantle. For example, do the earthquakes occur above or below seismic

reflectors associated with the subducted oceanic crustmantle boundary, or with the interplate interface? P–to–

S and S–to–P conversions from local earthquakes to

seismometers in northern Honshu, Japan show convincing evidence for a strong discontinuity at the upper edge

of the seismicity, probably associated with the interplate

interface [e.g., Hasegawa et al., 1978; Matsuzawa et al.,

1990; Zhao et al., 1997]. They infer that earthquakes

occur within both the crustal and mantle portions of the

old subducted Pacific plate.

In contrast, the width of the Cascadia seismogenic zone

is much narrower, perhaps as narrow as five kilometers.

Our goal is to determine the relative position of intraslab

earthquakes to reflector surfaces estimated using data

from the 1998 active-source Wet SHIPS experiment

[Fisher et al., 1999]. We observe strong secondary arrivals

(Figure 1) with mid-points under a broad area of the

northern Olympic Peninsula (Figure 2). The amplitudes

and polarities of reflected arrivals vary systematically over

spatial scales as short as one to two kilometers (Figure 1).

Beyond offsets of 80 km, the reflected arrivals are typically larger than direct arrivals. We picked approximately

500 travel times from stacks of reflected arrivals after

applying a low-pass filter at 14 Hz. Using three-dimensional ray tracing through a three-dimensional tomography model obtained from SHIPS and earthquake first

arrival data [Crosson et al., 1999], we determine the

locus of all point scatterers which would correspond to

an observed travel time for a given source and receiver.

This defines an approximately ellipsoidal surface from

which a ray could reflect and produce the observed travel

time. For coherent arrivals, a reflector must be locally

tangent to this ellipsoid. We invert for a smooth reflective surface that optimally is tangent to each of the data

ellipsoids.
38 | Creager et al.: Three-dimensional reflection image of the subducting Juan de Fuca plate

The optimal reflector is approximately 30 km deep

under the central Olympics, dipping about 13° to the

east, reaching a depth of about 60 km under Puget Sound

(Figures 2 and 3). Earthquakes that have been relocated

as part of the three-dimensional tomographic inversion

[Van Wagoner et al., 2000] define a planar surface that is

parallel to and consistently above this reflector (Figures

2 and 3). Note that earthquakes in this area move consistently up by a few kilometers, relative to the standard

Pacific Northwest Seismic Network (PNSN) catalog hypocenters, as a result of correcting for the inferred threedimensional velocities of this area. We interpret the

reflector as the Moho of the subducting slab. Thus, these

earthquakes occur completely within the subducting

oceanic crust.

We also image a weaker reflector, more geographically localized, which lies parallel to and just above the

seismicity, about seven kilometers above the deep reflector (Figure 4). Our preliminary interpretation is that

this reflector is associated with the interplate boundary.

The two major sources of error that affect uncertainties in the geometry of these reflectors are picking errors

and errors associated with the velocity model. Picking

errors are estimated to be less than 0.1 second for most

of the data; errors of this size will not effect our primary

conclusions. However, velocity model errors could be

quite large where sampling is poor. The shallowest structures under the Strait of Juan de Fuca are poorly sampled in our current three-dimensional model. Perturbing

the model in the upper ten kilometers of this region by

up to ± 30% and reinverting, the lower reflector surface

moved up by as much as ten kilometers in the westernmost quarter of our study region. However, the remaining eastern three-quarters of our area was unaffected

and remained beneath the seismicity. Below a depth of

20 km, the estimated model variance increases. We perturb this region uniformly by as much as ± 3%. Reinverting, we observed the surface to move vertically by

FIGURE 1: Record sections from station 6010 (triangle in Figure 2 at 123 ̊W, 47.7oN) from shots (red pluses in Figure 2) reduced at 9

km/s showing examples of the phases that are inferred to reflect off the subducted slab. The first arrivals are off the left of the

sections. Raw seismograms are low-pass filtered with a corner at 14 Hz. Each trace in this figure represents a phase-weighted stack

[Schimmel and Paulssen, 1997] of five adjacent, non-overlapping seismograms. Midpoints for each stack shift by 250 m. Note the rapid

spatial variation in the amplitude, apparent velocity, first motion and general character of the reflected arrivals.
The Cascadia Subduction Zone and Related Subduction Systems | 39

FIGURE 2: Map view of the inferred reflector geometry (color shading), stations with observed reflectors (triangles), shots (plusses,

red for shots recorded in Figure 1), relocated earthquakes (circles), and three-dimensional migrated reflection points (squares). Depths

of the reflector surface, earthquakes and migrated reflection points are given by the color bar.

FIGURE 3: Cross section of volume within the black box of Figure 2 showing standard Pacific Northwest Seismic Network (PNSN)

earthquake locations (small yellow circles), hypocenters relocated using the three-dimensional velocity model (large green circles)

and the migrated reflection points (squares). Larger squares represent higher quality picks. Cross sections of the data ellipsoids (see

text) are also shown as dotted blue lines. The reflector surface (thick blue line) optimally fits both the reflection points and the slope of

the data ellipsoids at the migrated reflection points.
40 | Creager et al.: Three-dimensional reflection image of the subducting Juan de Fuca plate

up to five kilometers with slight changes in slope. Again,

the lower surface remained beneath the seismicity. We

conclude that the lower reflector is indeed beneath the

seismicity despite uncertainties in the velocity model.

Our observations indicate that the intraplate seismicity

under the northern part of the Olympic Peninsula is completely confined to the crust of the subducting Juan de

Fuca slab, consistent with the dehydration embrittlement

hypothesis [Kirby, 1995].

Bounce point locations associated with rapid spatial

amplitude variations, like those seen in Figure 1 shots

2200-2600, correlate well with each other indicating that

these “bright spots” are associated with slab structure

and are perhaps a consequence of the partial transformation of basalt to eclogite. We have not yet been able

to rigorously determine polarities for the two reflector

surfaces, but the data appear to be of sufficient quality

to provide this information. Preliminary analysis of data

with clear polarities demonstrate that, like the amplitudes, polarities can also coherently and clearly change

over short spatial scales, suggesting the Moho reflector is

not a simple interface. The focal mechanisms of Wadati–

Benioff zone earthquakes in this region are quite variable but are most consistent with in-plane tension. This

is consistent with the stresses in the slab associated with

the 15% decrease in volume from the basalt to eclogite

phase transition. As the crustal part of the subducting

plate shrinks it imposes an in-plane compressive stress

on the subducting mantle, which in turn resists shrinking

thus putting the subducted crust in in-plane tension.

References

Crosson, R.S., K.C. Creager, N.P. Symons, T. Van Wagoner, Y. Xu, L.A. Preston, T.M. Brocher, T. Parsons,

M.A. Fisher, T.L. Pratt, C. Weaver, U.S. ten Brink, K.

Miller, A. Trehu, R. Hyndman, and G. Spence, Highresolution 3-D regional P wave velocity tomography

of the Puget basin region from SHIPS first-arrival data,

EOS, 80 (46), F763, 1999.

Fisher, M.A., T.M. Brocher, R.D. Hyndman, A.M. Trehu,

C.S. Weaver, K.C. Creager, R.S. Crosson, T. Parsons,

A.K. Cooper, D. Mosher, G. Spence, B.C. Zelt, P.T.

Hammer, U. ten Brink, T.L. Pratt, K.C. Miller, J.R.

Childs, G.R. Cochrane, S. Chopra, and R. Walia, Seismic survey probes urban earthquake hazards in Pacific Northwest, EOS, Trans. AGU, 80 (2), 13–17, 1999.

Green, H.W., II, and P.C. Burnley, A new self-organizing

mechanism for deep-focus earthquakes, Nature, 341,

733–737, 1989.

Hasegawa, A., N. Umino, and A. Takagi, Double planed

deep seismic zone and upper mantle structure in the

northeastern Japan arc, Geophys. J. R. Astron. Soc.,

54, 281–296, 1978.

Houston, H., and H.W. Green, II, The mechanics of deep

earthquakes, Ann. Rev. Earth Planet. Sci., 23, 169–

213, 1995.

FIGURE 4: Cross sectional view containing all data collapsed along the north–south direction showing both the upper and lower

reflector surfaces. See caption for Figure 3 for an explanation of symbols.
The Cascadia Subduction Zone and Related Subduction Systems | 41

Kirby, S.H., Localized polymorphic phase transformations

in high-pressure faults and applications to the physical mechanism of deep earthquakes, J. Geophys. Res.,

92, 13,789–13,800, 1987.

Kirby, S.H., Intraslab earthquakes and phase changes in

subducting lithosphere, Rev. Geophys., Suppl., U.S.

National Report to IUGG 1991–1994, 287–297, 1995.

Kirby, S.H., W.B. Durham, and L.A. Stern, Mantle phase

changes and deep-earthquake faulting in subducting

lithosphere, Science, 252, 216–225, 1991.

Matsuzawa, T., T. Kono, A. Hasegawa, and A. Takagi,

Subducting plate boundary beneath the northeastern Japan arc estimated from SP converted waves,

Tectonophysics, 181, 123–133, 1990.

Peacock, S.M., and K. Wang, Seismic consequences of

warm versus cool subduction metamorphism: examples from southwest and northeast Japan, Science,

286, 937–939, 1999.

Schimmel, M., and H. Paulssen, Noise reduction and detection of weak, coherent signals through phaseweighted stacks, Geophys. J. Int., 130, 497–505, 1997.

Van Wagoner, T.M., R.S. Crosson, K.C. Creager, L.A. Preston, N.P. Symons, T.M. Brocher, T. Parsons, T.L. Pratt,

C. Weaver, A. Trehu, and M.A. Fisher, 3-D structure

of the Cascadia forearc region from tomographic inversion of SHIPS active experiment and earthquake

observations, EOS, 81 (22), WP125, 2000.

Zhao, D., T. Matsuzawa, and A. Hasegawa, Morphology

of the subducting slab boundary in the northeastern

Japan arc, Phys. Earth Planet. Int., 102, 89–104, 1997.
42 |
The Cascadia Subduction Zone and Related Subduction Systems | 43

Modeling post-critical Moho reflections from intraslab

earthquakes in the Cascadia subduction zone

Jeff Shragge1, Michael Bostock1 and Garry C. Rogers2

1 Earth and Ocean Sciences, The University of British Columbia, Vancouver, British Columbia, V6T 1Z4, Canada

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

shragge@geop.ubc.ca, bostock@geop.ubc.ca, rogers@pgc.nrcan.gc.ca

Ten years ago, Cohee et al. [1991] undertook a study to

predict ground response of the Seattle Basin and Puget

Sound lowlands to a megathrust earthquake in Cascadia.

Their modeling suggested that the response at distances

beyond 75 km from the rupture could be dominated by

waves post-critically reflected from the oceanic Moho.

Wavefield radiation from intraslab earthquakes may be

similarly affected by post-critical Moho reflections. It is

therefore important to investigate this potential contribution to intraslab earthquake hazard for major population centres in the Pacific Northwest.

To examine this phenomenon and its importance in

southwestern British Columbia, we present two-dimensional pseudo-spectral synthetic seismogram computations for a model of the Cascadia subduction zone.

Large-scale velocity and density structure are obtained

from LITHOPROBE seismic studies on southern Vancouver Island [Drew and Clowes, 1990], gravity anomaly

modeling [Riddihough, 1979], and Wadati–Benioff

seismicity compiled by the GSC (see Figure 1). The

modeling includes both P–SV and SH propagation, and

the effect of source directivity for intraslab earthquakes

within and below the oceanic crust. This permits us to

examine the characteristics of post-critical reflections

from the oceanic Moho for a variety of source locations

and mechanisms.

To illustrate the potential hazard of post-critical

reflections from the dipping oceanic Moho, three displacement seismogram sections for an intraslab event

located beneath the accretionary sedimentary prism

complex offshore of Vancouver Island are presented in

Figures 2–4 (Uz, Ux, and Uy components respectively).

Post-critical reflections are highly visible at distances corresponding to the British Columbia lower mainland

between 160 and 250 km in epicentral distance (Figure

1) for both P–SV (Figures 2 and 3) and SH (Figure 4)

propagation. Note that in the case of SH propagation the

amplitude of the post-critical reflection at 190 km in epicentral distance is comparable to that of the direct S-wave

arrival at 50 km. Similar behavior is observed for P–SV

although results are generally dependent on the chosen

source mechanism.

References

Cohee, B.P., P.G. Somerville, and N.A. Abrahamson,

Simulated ground motions for hypothesized Mw=8

subduction earthquakes in Washington and Oregon,

Seism. Soc. Am. Bull., 81, 28–56, 1991.

Drew, J.J., and R.M. Clowes, A re-interpretation of the

seismic structure across the active subduction zone

of western Canada, in Studies of Laterally Heterogeneous Structures Using Seismic Refraction and Reflection Data, edited by A. G. Green, Geol. Surv. Can.

Paper 89–13, 115–132, 1990.

Riddihough, R.P., Gravity and structure of an active margin—British Columbia and Washington, Can. J. Earth

Sci., 16, 350–363, 1979.
44 | Shragge, Bostock and Rogers: Modeling post-critical Moho reflections from intraslab earthquakes

FIGURE 1: Top: model of Cascadia subduction zone beneath southwestern British Columbia after Drew and Clowes [1990].

Bottom: location of intraslab earthquake source.

FIGURE 2: Vertical component of displacement (Uz) for model configuration in Figure 1. The source consists of a summation of point

double couples to yield an extended double couple line source oriented along the dip of the slab.
The Cascadia Subduction Zone and Related Subduction Systems | 45

FIGURE 3: Radial component of displacement (Ux) for model configuration in Figure 1. The source consists of a summation of point

double couples to yield an extended double couple line source oriented along the dip of the slab.

FIGURE 4 : Transverse component of displacement (Uy) for model configuration in Figure 1. The source consists of a summation of

point rotations to yield an extended couple oriented parallel to the strike of the slab.
46 |
The Cascadia Subduction Zone and Related Subduction Systems | 47

Geometry and membrane deformation rate

of the subducting Cascadia slab

Ling-Yun Chiao1 and Kenneth C. Creager2

1 Institute of Oceanography, National Taiwan University, Taipei, Taiwan, R.O.C.

2 Department of Earth and Space Sciences, University of Washington, Seattle, Washington, 98195–1650, USA

chiao@ccms.ntu.edu.tw, kcc@ess.washington.edu

ABSTRACT

The Juan de Fuca plate subducts along the west coast of

North America along a plate boundary that extends some

1000 km from northern California to Vancouver Island.

In map view, the plate boundary is nearly straight along

its northern and southern regions, but exhibits a 35 ̊

change in orientation offshore from northern Washington State. Landward from this bend is an arch in the shape

of the subducting plate, a concentration in Wadati–Benioff

seismicity and an anomalously thick Miocene accretionary prism that makes up the two kilometer-high Olympic

Mountains. We speculate that all three are related,

directly or indirectly, to the fact that the Cascadia slab is

subducting into a concave, oceanward corner which

creates a geometric space problem analogous to pleats

in a tablecloth hanging off the corner of a table. To test

this hypothesis, we develop a theory of steady state, incompressible three-dimensional fluid flow that is asymptotically valid for a thin slab with high linear (or non-linear)

viscosity relative to the surrounding fluid. Flow in this simple model is constrained by kinematic boundary conditions rather than buoyancy forces. Fixing the slab geometry

to the same trench-normal profile along the entire subduction zone, we invert for the flow field that minimizes

the dissipation power of the system and find that strain

rates, characterized by along-strike compression, are concentrated along the west coast of Washington, reaching

values of 2x10-16 s-1. In a second experiment, we constrain the same 20 ̊ dip along the northern and southern

edges of the subducting plate and find that the optimal

slab geometry contains an arch in a location similar to

the seismologically observed arch, and that the root-mean

square of the strain rates are reduced by a factor of five.

Thus, if these strain rates are accommodated by intraplate

earthquakes, this arching model provides a geometry that

allows a five-fold reduction in seismic moment release

rates. The predicted strain rates are consistent with the

observed moment release rate associated with one Mw=7

and several Mw>6 earthquakes per century in the Olympic Mountain/Puget Sound region, and a much lower rate

to the north and south. Predicted strain rates associated

with slab bending and unbending are generally smaller

than the predicted membrane strain rates. The anomalously thick accretionary prism associated with the

Olympic Mountains can be explained by the critical

taper theory as being caused by the shallow dip of the

slab along the arch.

Introduction

The trend of the deformation front demarking the boundary between the North American and Juan de Fuca plates

bends by 35 ̊ from a north–south trajectory off Oregon

and Washington to N 35°W off Vancouver Island (Figure

1). Landward of this bend, the subducting oceanic plate

has a shallow 10 ̊ dip to a depth of about 50 km [e.g.,

Crosson and Owens, 1987; Weaver and Baker, 1988;

Creager et al. and Tréhu et al., this volume], whereas the

dip under Vancouver Island to the north is steeper (~18 ̊)

[Davis and Hyndman, 1989; Cassidy and Ellis, 1993]. The

slab dip to the south underneath Oregon is less well resolved. However, a few small earthquakes that are inferred to be slab events in Oregon suggest a steeper dip

in this region as well. The shallow dipping region represents an arch-like structure with an east–west trending

axis under the Olympic Mountains and Puget Sound (Figure 1). Support for this arching structure comes from the

analysis of receiver functions (P-to-S-wave conversions

off the subducting plate) [Crosson and Owens, 1987;

Owens et al., 1988; Cassidy and Ellis, 1993]. Tomographic

images [Rasmussen and Humphreys, 1988; VanDecar,

1991; Bostock and VanDecar, 1995] also indicate that

after passing beneath the Puget Sound region, the slab

goes to a steeper ~55 ̊ dip.

The arch structure has been attributed to a space

problem [e.g., Rogers, 1984; Crosson and Owens, 1987],

which can be visualized by considering the analogy between the subducting plate and a tablecloth hanging over

the corner of a table. The tablecloth has folds as it hangs

over the corner. The extra material can be accommodated either by along-arc compression, or by geometric

buckling or arching. The primary purpose of this paper
48 | Chiao and Creager: Geometry and membrane deformation rate of the subducting Cascadia slab

is to analyze the possible relationships among this bend

in the trench, the shape and internal deformation of the

subducted plate, the spatial distribution of deep seismicity, and the location of the Olympic Mountains.

The well-located microseismicity inferred to be in the

down-going plate, recorded by the Pacific Northwest Seismic Network (PNSN) and the Pacific Geoscience Centre

(PGC) during the past three decades, as well as large earthquakes during the twentieth century (Table 1), are concentrated in the vicinity of the Puget Sound basin (Figure

1). There are markedly fewer events south of Olympia,

Washington and even fewer south of Portland, Oregon.

Perhaps even more striking is the variation in strain rate

inferred from this seismicity. From 47.1 ̊ to 48.8 ̊N (Olympia to Victoria, British Columbia), several magnitude

5.5 to 7 earthquakes are well documented whereas to

the north and south the largest known deep earthquake

is a magnitude 3.7. Using standard magnitude to seismic

moment scaling [Hanks and Kanamori, 1979], it appears

that the seismic moment release rate (which is proportional to strain rate) within these latitudes is five orders of

magnitude greater than that outside.

Nearly all other subduction zones, which are characterized by an oceanward concave bend in the axis of the

trench, also exhibit an arching structure with anomalously

shallow dip landward of the bend. Examples include the

Pacific plate subducting under Alaska [Page et al., 1989]

and the Kamchatka, Kuril, Japan, Izu–Bonin trench system which, taken as a whole, has an oceanward-concave curvature. The Japan slab, at the center, exhibits a

shallow 30 ̊ dip, while the dips under Izu–Bonin to the

south and Kamchatka to the north are 90 ̊ and 50 ̊

respectively. On a smaller scale, arc–arc junctions such

as the Hokkaido corner, the corner between Honshu and

Izu–Bonin, and the New Britain–Solomon corner all exhibit pronounced arches [Yamaoka et al., 1986; Yamaoka

FIGURE 1: Lambert projection of earthquakes (circles) inferred to be within the subducting Cascadia plate, slab depth contours (thick

lines) showing the arch structure inferred by Crosson and Owens [1987], and active volcanoes (triangles). Earthquake magnitudes and

depths are indicated by the scale. Earthquakes with magnitudes exceeding 2 are from the Pacific Northwest Seismic Network and the

Pacific Geoscience Centre Network. Older, large earthquakes are also shown.
The Cascadia Subduction Zone and Related Subduction Systems | 49

and Fukao, 1987; Chiao, 1991] in the upper few hundred kilometers. A counter example is the Bolivia orocline

which currently has an anomalously steep dip landward

of a bend in the trench [Creager et al., 1995].

The Olympic Mountains are part of an Eocene to

Miocene highly deformed sedimentary and meta-sedimentary accretionary complex which is bounded to the

north and east by the horseshoe shaped Eocene Crescent basalts. This highly anomalous, subaerial accretionary wedge is confined to less than 100 km along the arc,

but extends nearly 200 km from the deformation front.

This unusual geometry has been attributed to the shallower slab dip beneath the Olympic Peninsula [Davis and

Hyndman, 1989; Brandon and Calderwood, 1990;

Brandon et al., 1998] as described in more detail below.

In order to understand these relations among slab

geometry and internal deformation, membrane strains

(or strain rates) associated with slab flow have been

analyzed using physical models [Yamaoka et al., 1986;

Yamaoka and Fukao, 1987], analytic solutions to very simple geometries [Bevis, 1986; Bevis, 1988] and numerical

analyses [Burbach and Frohlich, 1986; Chiao, 1991;

Creager and Boyd, 1991; Creager et al., 1995]. Deformation in plate interiors, as observed by seismic activity,

is generally very low. However, the Gorda plate is a counter example in which high levels of seismicity occur

throughout the interior of the plate. These large in-plane

strain rates are forced by geometric constraints imposed

by the much older and stronger Pacific plate [Denlinger,

1992]. Similarly, large in-plane or membrane strain rates

are required by the geometry of subduction. In this

paper, we extend the methods developed by Creager

and Boyd [1991] and Chiao [1991] to analyze the membrane and bending strain rates in the Cascadia subduction zone. We explore, not only the orientation and

amount of the strain rate required of a given slab geometry, but also determine the geometry that is most compatible with a given trench configuration and other

geometric boundary conditions.

Kinematic modeling of in-plane deformation rate

We postulate that the arching structure, the concentration of intraslab earthquakes along the arch axes, and

the anomalously thick accretionary prism located above

the arch are all linked directly or indirectly to the bend

in the trench. The essence of our approach is to consider solutions to the three-dimensional fluid flow problem that are asymptotically valid in the vicinity of a

smooth surface defining the slabs cold thermal core. We

assume that: 1) the effective viscosity of the fluid is high

near this surface and decreases rapidly away from it; 2)

no flow crosses this surface; 3) the flow is steady state;

and 4) incompressible. First, we provide arguments to

justify and motivate these assumptions and then show

how they can be used to reduce the essence of a fully

three-dimensional flow problem to a calculation in twodimensions on the ‘slab surface’ which is valid both on

the surface and in the vicinity of the surface. See Chiao

[1991] for details.

The rheology of subducting slabs and the surrounding

mantle is not well constrained. However, most mechanisms associated with steady state flow are observed in

the laboratory to depend exponentially on homologous

temperature (absolute temperature divided by the solidus)

[e.g., Carter, 1976]. Creager and Boyd [1991] estimate

the homologous temperatures along the Aleutian subduction zone by dividing thermal models by the pressure

dependent solidus of a dry peridotite [Takahashi and

Kushiro, 1983]. They estimate very low values of homologous temperature (<0.4) in the seismogenic parts of the

slab, comparable to the homologous temperatures predicted for the seismogenic parts of the oceanic lithosphere

prior to subduction. This suggests that the effective viscosity in the cold core of slabs is orders of magnitude

greater than in the surrounding asthenosphere (see also

Davies [1999]). The thickness of the stiff part of the

subducted plate is thus expected to be comparable to the

mechanical thickness of the oceanic lithosphere, which

for young oceanic plates is 10–20 km [McNutt, 1984]. It

appears reasonable to treat the slab as a thin sheet given

that its lateral dimensions are in the order of 1000 km. A

consequence of a large contrast in effective viscosity is

that large-scale shear strain rates parallel to the slab surface will occur preferentially in the weaker fluid. We assume that within the cold core of the slab, there is no

shear-strain rate parallel to the slab. This is analogous to

the Kirchhoff’s Hypothesis in studying the deformation of

elastic shells [e.g., Calladine, 1983].

Consider a three-dimensional, incompressible, steadystate flow field v(x). The symmetric part of the gradient of

the particle velocities is the deformation rate tensor field

D(x). Consider a smooth, curved surface that describes

the slab geometry and assume that the slab flow does not

cross this surface. A mathematically convenient way to

impose Kirchhoff’s Hypothesis is to construct the local

projection operator (P=I – n⊗n) [e.g. Parlett, 1980] which

projects vectors onto the slab surface if n is the unit vector locally normal the surface, n⊗n is a dyadic (outer)

product, and I is the identity tensor. Operating on D by P

from both sides, Dpp= P D PT annihilates shear parallel to

the slab as well as compression or extension in the slab

normal direction [Chiao, 1991], while not modifying the

in-plane part of the deformation rate. To comply with

the assumption that the fluid is incompressible, while

maintaining Kirchhoff’s Hypothesis, we define

Din= Dpp – tr(Dpp) n⊗n

where tr is the trace (sum of diagonals). Din is then the

strain-rate tensor describing the membrane deformation

rate and the thickening (or thinning) of the thin viscous
50 | Chiao and Creager: Geometry and membrane deformation rate of the subducting Cascadia slab

slab. It represents the deformation of a flow field that

satisfies Kirchhoff’s Hypothesis and incompressibility. Furthermore, these two approximations allow us to uniquely

evaluate all nine components of the deformation rate

tensor within the high viscosity part of the fluid from

knowledge of the flow field in the plane of the slab. In

other words, within the high-viscosity region of the slab,

D asymptotically approaches Din as the viscosity contrast

increases and the slab thickness decreases.

The effective strain rate (γ) within the slab is defined by:

γ2 = 1/2 Dij Dij ≈ 1/2 Din

ij Dinij (1)

= 1/2 [Dpp

ij Dppij + tr(Dpp)

2

 ]

Suppose the slab has a power-low viscous rheology [eg.,

Ranalli, 1987] in which D is proportional to the nth power

of the deviatoric stress: Dij = B-n (1/2 τkl τkl)

n-1 τij, where τkl

is the deviatoric stress tensor. B depends on temperature

and is much larger in the cold slab than outside. The

dissipation power of the entire flow within the region of

high viscosity equals the thickness, H, of this region times

the surface integral over the entire slab of half the stress

deviator contracted with the strain rate tensor [Chiao,

1991]:

I = H ∫ 1/2 τij Dij dA = H B ∫ γ1+1/n dA (2)

In summary, given a slab geometry and a continuous

flow field following that surface, we calculate the deformation rate tensor and the dissipation power (2) associated with the flow field in the high viscosity region near

the slab. The goal is to use this forward formulation to

solve the inverse problem of determining the slab geometry and the flow field that minimizes the dissipation

power. In this simplified formulation, the flow is driven

by boundary conditions rather than buoyancy forces. We

specify the flow along a spherical shell corresponding to

the pure rotation of the oceanic lithosphere prior to subduction and force the slab to remain at the earth’s surface until it reaches a predefined trench location. We

then specify part or all of the slab geometry and invert

for the flow field that minimizes the dissipation power.

According to equation (2), a Newtonian rheology (n=1)

corresponds to minimizing the L2 norm of the effective

strain rate, while a power-law rheology with n=∞ corresponds to minimizing the L1 norm. We consider solutions in these two end-member rheologies.

Results

Constant dip geometry, Newtonian rheology

In our first experiment, we fix the geometry to have a

dip of 20 ̊ along the entire trench. To go from a dip of 0

at the trench to 20 ̊, we determine the geometry that

minimizes the integral of the change of curvature [Chiao,

1991], which minimizes the total bending strain rates.

The boundary condition is that, prior to entering the

trench, the flow field satisfies the relative plate motions

on a sphere described by the Euler pole at (29.11 ̊N,

-112.72 ̊E) and an angular velocity of 1.05 ̊ per million

FIGURE 2: Flow field inversion of the constant-dip experiment with a Newtonian rheology. a) Depth contours, at 30 km intervals,

showing the fixed geometry of the subducting plate, earthquakes, and three of the inferred particle paths. b) Principal axes of calculated membrane strain rates (bold red bars show compression and thin blue line segments show extension). Length is proportional to

the value of the strain rate as shown by the scale. c) Spatial distribution of effective strain rate. Root mean square of the effective

strain rate over the entire area of the subducted plate is 1.3x10-16 s-1 .
The Cascadia Subduction Zone and Related Subduction Systems | 51

years [Nishimura et al., 1984]. Using a finite element

formulation, we invert for the flow field v(x) on the slab

surface that minimizes the dissipation power (2) and satisfies the given slab geometry and boundary condition

(Figure 2). From this flow field, we calculate the strainrate tensor and display its principal axes, their corresponding magnitudes, and the effective strain rate, γ

(Figure 2). The bend of the trench clearly imposes a

smoothly varying along-arc compression which reaches

values as high as 2x10-16 s-1. This integrates along particle

paths to strains of several percent. The root-mean square

of the effective strain rate is 1.3x10-16 s-1.

Arch geometry, Newtonian rheology

Next we relax the geometric boundary conditions such

that the geometry is imposed only along the trench (and

oceanward of the trench) and along two cross sections

at the northern and southern edges of the slab. These

cross sections are the same as in the constant dip model.

Now both the geometry of the slab surface and the flow

field within this surface are allowed to vary to find the

minimum deformation-rate solution. The root-mean

square of the effective strain rate is reduced by a factor

of five relative to the constant dip model.

Arch geometry, power-law rheology

In a third experiment, we use the same boundary conditions, but consider a power-law rheology with a very

large exponent (n=105

). The resulting best slab geometry is nearly identical to the solution for the linear

rheology so we show only the result for the power-law

rheology (Figure 3). The main difference is that the powerlaw rheology causes a concentration of the region of high

effective strain rate and a higher contrast relative to background values. The resulting slab geometry exhibits a

pronounced arch (Figure 3a), much like that previously

estimated (Figure 1) from seismicity and the analysis of

receiver functions [Crosson and Owens, 1987].

Furthermore, the 10 ̊ and 12 ̊ slab-dip isolines (Figure 3c) outline the geographic location of the Olympic

Mountains. These mountains are highly anomalous in that

they represent an extremely thick subaerial accretionary

prism, with an along-strike extent of less than 100 km

and a trench normal extent of nearly 200 km. One explanation for these mountains is that the geometric bend

in the trench produces a geometric space problem that

can be accommodated by large compressional strain rates

within the plane of the slab (Figure 2) or geometric arching (Figure 3). The observed slab geometry (Figure 1) suggests that the minimum global energy solution is to allow

arching. An important consequence of the arching is the

anomalously shallow slab dip (<10 ̊ for 200 km landward of the deformation front). The critical taper theory,

invoked to explain the deformation and mechanics of

accretionary prisms [e.g. Davis et al., 1983; Davis and

Hyndman, 1989], suggests that for a wide range of rheological parameters, a slab dip of less than 11 ̊ produces a

positive slope to the top of the wedge, such that the

topography of the earth’s surface will go up landward of

the trench. In contrast, once the slab dip increases past

FIGURE 3: The variable-dip experiment showing results of the full inversion involving both the flow field and the slab geometry (see

captions for Figure 2) for a power-law rheology (n=106

). The 10 ̊ and 12 ̊ iso-dip contours of the slab dip shown in panel c), come

onshore under the Olympic Mountains, providing a possible explanation for the existence and location of this thick accretionary prism

as explained in the text. The mean effective strain rate (MES) for power-law rheology is 2x10-17 s-1.
52 | Chiao and Creager: Geometry and membrane deformation rate of the subducting Cascadia slab

11 ̊, the wedge will decrease in elevation. Indeed, the

predicted location of the 11 ̊ dip isoline (Figure 3) shows

a clear triangular area that extends inland in the Olympic

Peninsula region, imitating the distribution of the Crescent basalt backstop and the locus of the thick accretionary wedge material.

Bevis [1988] derived a simple formula to relate the

seismic moment release and the average strain rate for

intraslab earthquakes:

Date Latitude Longitude Depth Magnitude Seismic moment

( ̊N) ( ̊E) (km) (Newton-m)

01/11/09 48.7 -122.8 75 6.0

11/13/39 47.4 -122.6 60 6.2

02/15/46 47.3 -122.9 60 6.2

04/13/49 47.1 -122.7 54 7.1 1.5x1019 [Baker and Langston, 1987]

04/29/65 47.4 -122.3 59 6.5 1.4x1019 [Langston and Blum, 1977]

dynamics and impose that geometry as a boundary condition. We also do not explicitly incorporate bending

deformation, though our two-dimensional calculations

suggest that maximum strain rates associated with bending are comparable in magnitude to the membrane

deformation rates. It is possible that most of the bending

deformation occurs just below the seismicity cutoff

through ductile creep mechanisms allowed in these

warmer portions of the slab. Also, the basalt–eclogite

phase transition causes a volume reduction of about 15%

in the subducted oceanic crust. This undoubtedly has

an important effect on slab deformation that has not been

incorporated in our calculations. We also assume a steady

state flow, even though it is clear that the geometry of

subduction has changed during the past 20 Ma as the

Oregon coast ranges have rotated clockwise to exaggerate the bend in the deformation front.

Even though our calculations do not provide a complete formulation of the dynamics of the Cascadia subduction zone, they provide a method to quantify the

space problem imposed by a subduction zone with a

concave oceanward trench. We show that an arch structure is an efficient way to accommodate the space problem and reduce the in-plane deformation by a factor of

five. Indeed, it is observed that nearly all subduction

zones with concave oceanward trenches exhibit similar

pronounced arches in slab geometry. The shallow slab

dip offers an explanation for the anomalously thick

subaerial accretionary prism associated with the Olympic Mountains. In addition, the space problem is most

severe landward of the bend in the deformation front,

providing a plausible explanation of the high level of

earthquake moment release rates in the Puget Sound

region relative to regions to the north and south. Finally,

the predicted membrane strain rates are the same size

as the strain rate inferred by summing the seismic moments of slab earthquakes.

TABLE 1: Large 20th century earthquakes inferred to be in the subducting Cascadia slab.

where ε is the average effective strain rate, μ is the rigidity, Σ M is the total seismic moment release summed

over a time period t, and W, L, and H are along-strike

width, downdip length and slab thickness. Assuming μ=

7x1010 Newtons/m2, W=200 km, L=200 km, H=10 km,

and t=100 years for the Cascadia slab, we calculate a

strain rate of 2x10-16 s-1 in excellent agreement with the

peak calculated membrane strain rates (Figure 2). Note

that over 99% of the summed moment release comes

from the five events in Table 1.

The bending strain rate, averaged across the thickness (H) of a plate is V (dκ/ds) H/4 [Chiao, 1991] where V

is the velocity of the subducting plate and dκ/ds is the

rate of change of the curvature of the plate with respect

to position along particle paths. In keeping with the philosophy of the membrane strain rate calculations, we

calculate slab cross sections that produce the minimum

integrated dissipation power associated with bending

strain rates and find that the bending strain rates must

reach values of at least 10-16s-1 to reach a slab dip of 20 ̊,

so the predicted bending strain rates are slightly smaller

than the membrane strain rates, but in some places may

be more important.

Discussion

Our formulation is kinematic. We have not included important dynamics such as gravitational buoyancy that

drives the subduction process. Instead, we observe the

slab geometry, which presumably is a result of the

.
The Cascadia Subduction Zone and Related Subduction Systems | 53

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western Oregon using teleseismic P-wave travel times,

Geophys. Res. Lett., 15, 1417–1420, 1988.

Ranalli, G., Rheology of the Earth: Deformation and Flow

Processes in Geophysics and Geodynamics, 366 pp.,

Allen and Unwin Inc., Winchester, Massachusetts,

1987.

Rogers, G.C., Seismotectonics of British Columbia, PhD

thesis, University of British Columbia, Vancouver,

1984.

Takahashi, E., and I. Kushiro, Melting of a dry peridotite

at high pressure and basalt magma genesis, Am. Mineralogist, 68, 859–879, 1983.

Tréhu, A.M., T.M. Brocher, K.C. Creager, M.A. Fisher, L.A.

Preston, G. Spence and the SHIPS98 Working Group,

Geometry of the subducting Juan de Fuca plate: new

constraints from SHIPS98, in The Cascadia Subduc-
54 | Chiao and Creager: Geometry and membrane deformation rate of the subducting Cascadia slab

tion Zone and Related Subduction Systems, edited

by S.H. Kirby, K. Wang and S.G. Dunlop, pp. 25–32,

U.S. Geological Survey Open-File Report 02–328,

Geological Survey of Canada Open File 4350, 2002.

VanDecar, J.C., Upper-mantle structure of the Cascadia

subduction zone from non-linear teleseismic inversion, PhD thesis, University of Washington, Seattle,

1991.

Weaver, C.S., and G.E. Baker, Geometry of the Juan de

Fuca plate beneath Washington and northern Oregon

from seismicity, Bull. Seismol. Soc. Amer., 78, 264–

275, 1988.

Yamaoka, K., and Y. Fukao, Why do island arcs form cusps

at their junctions?, Geology, 15, 34–36, 1987.

Yamaoka, K., Y. Fukao, and M. Kumazawa, Spherical shell

tectonics: effects of sphericity and inextensibility on

the geometry of the descending lithosphere, Rev.

Geophys., 24, 27–55, 1986.
The Cascadia Subduction Zone and Related Subduction Systems | 55

Crustal structure of the northern Juan de Fuca plate and

Cascadia subduction zone—new results, old data

Ron M. Clowes

Earth and Ocean Sciences, The University of British Columbia, 6339 Stores Road, Vancouver, V6T 1Z4, Canada

 clowes@lithoprobe.ubc.ca

Background

In 1980, the Vancouver Island Seismic Project (VISP 80;

Ellis et al. [1983]) recorded a series of marine and

onshore–offshore seismic refraction/wide-angle reflection

(R/WAR) profiles on the Juan de Fuca plate and Vancouver Island. Most of the results from this experiment have

been published. However, interpretation of crustal structure for the Juan de Fuca plate along profile VISP II (Figure 1), while completed for a decade, has never been

published. The results are highly unusual, indicating a

thickness of the oceanic plate, exclusive of sediments, in

excess of ten kilometers near ocean bottom seismograph

(OBS) 4 but just over seven kilometers near OBS 1.

In 1989, the Reflection-refraction Experiment–

Cascadia Accretionary Prism (RECAP’89) was a “piggyback” experiment in which The University of British

Columbia (UBC) deployed seismographs on Vancouver

Island to record airgun shots fired along a series of profiles during an offshore multi-channel reflection survey

acquired by the GSC in support of the Ocean Drilling

Program [Spence et al., 1991]. Wang [1997] has interpreted data from two lines, 89–03/02 and 89–09,

recorded at sites 2A–B and 4A–B, respectively (Figure

1); however, the results have never been published. The

crustal velocity structure models provide the only wellconstrained interpretation of seismic velocities for the

region underlying the continental shelf and slope, and

thus have significance for accurate determination of

earthquake hypocenters offshore.

Crustal velocity structure of

northern Juan de Fuca plate

The VISP II refraction profile was recorded on three OBSs

deployed at the ends and middle of a 110 km line oriented parallel to the margin on the ocean basin of the

Juan de Fuca plate (Figure 1). An airgun source at 0.5 km

spacing and explosive charges at ~2.5 km spacings were

fired into the OBSs. A 3.5 kHz sounding profile and a

single-channel reflection profile were acquired along the

line to provide water depths, structure of the sediments

and basement topography. Data from both the airgun

and explosive sources were rich in direct, converted and

multiple reflected refracted phases, as observed from

vertical and horizontal component geophones and the

hydrophone channel. From these phases, three onedimensional velocity models were interpreted; these were

interpolated to a two-dimensional starting model. Iterative forward modelling of travel times and amplitudes,

using a two-dimensional, ray-theory-based synthetic

seismogram program, RAYAMP, was used to interpret the

final velocity structure model for line VISP II (Figure 2).

Two significant features of the VISP II model (Figure 2)

are: 1) the large increase in oceanic crustal thickness, from

~7.5 km in the southern part of the model to ~10.5 km

in the northern part, this increase being due almost entirely to the high velocity wedge (hvw) and the underlying

low velocity layer (lvl) included in layer 3B; and 2) the

decrease in velocity in the uppermost mantle in the northern half of the model, this being required by travel times

of clear Pn phases recorded on the three OBSs.

Figure 3 is a map of the thickness of oceanic crust for

northern Juan de Fuca and Explorer plates compiled from

VISP II and other refraction interpretations and from

multi-channel seismic reflection travel times, converted

to depths using 6.5 km/s, for data on which oceanic Moho

could be readily identified. The pronounced increase of

>3 km in crustal thickness of the Juan de Fuca plate

near the Nootka transform fault zone is the most prominent feature. An enigma is the lack of a corresponding

anomaly in the free air gravity field. The position of the

thickened crust may have relevance for the hypothesis

that the Explorer plate is in the process of being dismembered and a new Pacific–North American plate boundary is being established [Rohr and Furlong, 1995], because

the southern end of the latter coincides approximately

with the thickened crust. An alternative or additional

explanation for the thickened crust may be asymmetric

magma accretion associated with the outside corner of

the Juan de Fuca Ridge–Sovanco transform junction, similar to results from recent studies of the Mid-Atlantic Ridge

[Allerton et al., 2000].
56 | Clowes: Crustal structure of the northern Juan de Fuca plate and Cascadia subduction zone

FIGURE 1: Location of seismic refraction/wide-angle reflection studies discusssed in this contribution on a seismicity map of Vancouver Island and the offshore region. New results are presented for the Vancouver Island Seismic Project (VISP) II (Figure 2) from ocean

bottom seismographs (OBS) 1, 2 and 4 on the northern Juan de Fuca plate; VISP I (OBSs 1, 3 and 5) has been published. Also, new

results are presented for the northern Cascadia subduction zone: airgun reflection lines 89–03/89–02 recorded at land stations 2A and

2B and line 89–09 recorded at land stations 4A and 4B (Figure 4). DF, deformation front; DK, Dellwood Knolls; ExR, Explorer Ridge;

JdFR, Juan de Fuca Ridge; NFZ, Nootka fault zone; SFZ, Sovanco fracture zone; TWK, Tuzo Wilson Knolls. Inset shows location of main

map area within the Juan de Fuca plate system.

FIGURE 2: VISP II velocity structure model with primary oceanic crustal layers identified. Velocities (km/s) at the top of layers and the

velocity gradients (km/s/km) in the layers (numbers in parentheses) are shown on the model. Velocity-depth profiles at the position of

each ocean bottom seismograph (OBS) are superimposed. Grey lined region is not constrained by the data.
The Cascadia Subduction Zone and Related Subduction Systems | 57

Crustal velocity structure of the

Cascadia subduction zone

For RECAP’89, four continuous recording seismographs

at eight land sites on the west coast of Vancouver Island

acquired good quality data from airgun sources fired at

50 m intervals along four offshore mulit-channel seismic

(MCS) reflection lines; two of these lines recorded at four

stations have been interpreted (Figure 1). Four principal

phases—refractions through the upper continental crust,

reflections from the top of the oceanic crust and the

Moho, and refractions through the oceanic mantle—were

identified. An iterative combination of two-dimensional

travel time inversion and amplitude forward modeling

was used to interpret crust and upper mantle P-wave

velocity structure.

Figure 4 shows the final velocity structure model for

line 89–09; that for 89–02/03 is similar in general, but

has differences in detail. Velocities for the accretionary

wedge and Pacific Rim terrane are well constrained; they

increase landward and with depth. The crust of the

subducting Juan de Fuca plate is defined well by the wideangle reflection data; along 89–09, the top of the plate is

~2 km shallower than that along 89–02/03 below the

ocean basin, and 3–4 km shallower below the continental slope. Whereas profile 89–02/03 required a substantial down plate increase in velocity, this feature was not

required for 89–09. For both profiles, average oceanic

upper mantle velocities of ~8.0 km/s were determined,

although some unusual lateral variations were interpreted;

these must be viewed with caution due to relatively poor

resolution in this part of the model. These results are the

first velocity models for the region below the continental

shelf and slope that are well constrained by turning and

reflected ray paths in that region. Given that most earthquakes in southwestern Canada occur offshore (Figure 1)

but are located primarily from stations located on land,

an improved velocity structure for the region between the

hypocenters and the stations should lead to more reliable

earthquake locations.

Acknowledgements

W.R. Hugh White of Toronto, Ontario, now retired and

formerly with Centennial College in Toronto, carried out

a detailed interpretation of the VISP II profile during and

following a one-year study leave period at the University of British Columbia in 1981–82. Carl Wang, now

working in the petroleum industry in Calgary, Alberta,

FIGURE 3: Contours (heavy black lines) and point determinations (numbers in white ellipses) of thicknesses of the oceanic plates in

kilometers from Vancouver Island Seismic Project (VISP) II and other refraction and reflection lines (thin grey lines) in the offshore

region superimposed on a free air/Bouguer gravity anomaly map. Thick and dashed grey lines outline tectonic elements, with same

abbreviations as Figure 1. Wiggly line in the Nootka fault zone (NFZ) indicates a possible vertical offset in crustal thickness.
58 | Clowes: Crustal structure of the northern Juan de Fuca plate and Cascadia subduction zone

carried out interpretation of profiles 89–02/03 and 89–

09 for his M.Sc. at the University of British Columbia.

References

Allerton, S., J. Escartin, and R.C. Searle, Extremely asymmetric magmatic accretion of oceanic crust at the

ends of slow-spreading ridge segments, Geology, 28,

179–182, 2000.

Ellis, R.M., G.D. Spence, R.M. Clowes, and 13 others,

The Vancouver Island seismic project: a COCRUST

onshore–offshore study at a convergent margin, Can.

J. Earth Sci., 20, 719–741, 1983.

Rohr, K.M.M., and K.P. Furlong, Ephemeral plate tectonics at the Queen Charlotte triple junction, Geology,

23, 1035–1038, 1995.

Spence, G.D., R.D. Hyndman, S. Lanton, C.J. Yorath, and

E.E. Davis, Multi-channel seismic reflection profiles

across the Vancouver Island continental shelf and slope,

Geological Survey of Canada, Open File 2391, 1991.

FIGURE 4: a) Velocity structure model for profile 89–09: airgun shots at 50 meter intervals between arrows at top of model were

recorded at stations 4A and 4B. Numbers within model represent P-wave velocities in km/s at those positions. Dashed lines are

boundaries for model layers or blocks; heavy solid lines are boundaries from which wide-angle reflections were observed. b) Velocity

structure model with 1:1 representation. SED, sedimentary layer; AW, accreted wedge; CT, Crescent terrane; PRT, Pacific Rim terrane;

WR, Wrangellia; JdF, Juan de Fuca plate.

Wang, C.X., Structure of the Cascadia subduction zone

off Vancouver Island: new evidence from seismic refraction data, M.Sc. thesis, University of British Columbia, Vancouver, 1997.
The Cascadia Subduction Zone and Related Subduction Systems | 59

Seismic reflection imaging of the

Cascadia plate boundary offshore Vancouver Island

Andrew Calvert

Department of Earth Sciences, Simon Fraser University, Burnaby, British Columbia, V5A 1S6, Canada

acalvert@sfu.ca

Two seismic reflection surveys have been shot over the

continental slope and shelf of the convergent margin offshore Vancouver Island, in 1985 and 1989 (Figure 1).

The 1989 survey, which is shown here and discussed

more fully by Calvert [1996], was shot using a tuned

airgun array with a total volume of 128 l (7820 in3

), and

recorded by a 144 channel hydrophone streamer with a

25 m group interval and 3600 m far offset. The 50 m

shot point separation provided 36 fold data. The recording parameters for the earlier 1985 survey are similar to

those used in the 1989 survey, with the exception that a

120-channel streamer was then employed. The processing sequences applied to the data from both surveys are

essentially similar, consisting of amplitude recovery,

designature, F–K filter, predictive deconvolution to suppress multiples, velocity analysis, normal moveout correction, and stack. A post-stack predictive deconvolution

was also applied in order to suppress water-bottom

reverberations. A constant velocity migration using a value

of 1500 ms-1 was applied to each line to focus waterborne coherent noise that contaminates a number of the

profiles; the noise appears in the unmigrated data as

broad hyperbolae at late record times. A dynamic amplitude equalization was then applied to reduce the

amplitude of the focussed coherent noise relative to the

rest of the profile and followed by a residual migration.

Both vintages of seismic data image structural boundaries in both the sedimentary prism and two accreted

terranes. The subduction décollement, which separates

the oceanic crust of the descending Juan de Fuca plate

from the overlying continental units, is observed intermittently on the seismic profiles from the area (Figure 2).

At the base of the continental slope, thrust faults within

the incoming sedimentary section sole out close to the

top of the igneous oceanic crust, suggesting that the subduction décollement occurs close to the top of the igneous crust even though it cannot be identified directly

[Davis and Hyndman, 1989]. The top of the oceanic crust

seaward of the continental slope is marked by a high

amplitude two-cycle reflection. However, beneath the

continental slope this reflection evolves into the broader

(in time) zone of reflectivity that characterizes the plate

boundary beneath the continental shelf [Calvert and

Clowes, 1990]. The amplitude of the seismic reflection

from the plate boundary appears to vary in both dip and

strike directions. The variation normal to the margin is

mainly attributable to the difficulty in obtaining a consistent image of the plate boundary beneath the strong

lateral variations in seismic velocity caused by the rugged topography at the top of the accretionary wedge.

However, amplitude variations parallel to the margin,

particularly beneath the volcanic Crescent terrane, where

near-surface conditions change slowly, may well be

attributable to changes in physical properties at the plate

boundary. Close to southern Vancouver Island, the plate

boundary reflection splits in two: the upper highamplitude reflection package continues into the broad

band of ‘E’ reflectivity identified beneath Vancouver

Island; and the lower weaker, short duration reflection

marks the top of the descending oceanic plate. This splitting of the plate boundary reflection is interpreted to

occur at the seaward edge of a slab of mafic material

underplated to southern Vancouver Island and is not

observed further north (Figure 3).

References

Calvert, A.J., and R.M. Clowes, Deep, high amplitude

reflections from a shear zone above the subducting

Juan de Fuca plate, Geology, 18, 1091–1094, 1990.

Calvert, A.J., Seismic reflection constraints on imbrication

and underplating of the northern Cascadia convergent

margin, Can. J. Earth Sci., 33, 1294–1307, 1996.

Davis, E.E., and R.D. Hyndman, Accretion and recent

deformation of sediments along the northern Cascadia

subduction zone, Bull. Geol. Soc. Am., 101, 1465–

1480, 1989.

Dehler, S.A., and R.M. Clowes, Integrated geophysical

modelling of terranes and other structural features

along the western Canadian margin, Can. J. Earth Sci.,

29, 1492–1508, 1992,

Hyndman, R.D., C.J. Yorath, R.M. Clowes, and E.E. Davis,

The northern Cascadia subduction zone at Vancouver Island: seismic structure and tectonic history, Can.

J. Earth Sci., 27, 313–329, 1990.
FIGURE 1: Location of the major terranes beneath the Tofino basin in the southern part of the Canadian convergent margin (after Hyndman e

and Dehler and Clowes [1992]). Offshore, the terrane boundaries are defined by magnetic data and the indicated seismic lines; onshore

mapping and potential field data are the primary constraints. The location of the multi-channel seismic lines which extend over the continen

shown. CMP numbers are annotated to lines shown here.
FIGURE 2: a) Line 89–09. b) Line 89–06. Both seismic profiles are displayed after migration with a velocity of 1500 ms-1 to focus water-borne 

undeformed stratigraphy above the igneous oceanic crust can be seen at the seaward end of both profiles. High amplitude reflections arising in 

just above the subducting igneous crust can be followed beneath the continental slope on line 89–09 where the incoming sediments are deform

into the accretionary wedge. In contrast, strong reflections from the top of the subducting crust do not extend seaward of the continental 

Arrows indicate the approximate position and orientation of some thrusts in the sedimentary prism. Both seismic lines are displayed with no 

order to permit an approximate comparison of the reflection amplitudes. The changing water depth, the presence of multiple reflections, 

variations arising from deformation of wedge and slope sediments prevent detailed interpretation of the seismic data beneath the continental 
62 | Calvert: Seismic reflection imaging of the Cascadia plate boundary offshore Vancouver Island

FIGURE 3: Cartoon showing a generalized interpretation of the main features observed in the offshore reflection profiles normal to the

margin, and a possible model for formation of the zone of E reflectivity. Vertical exaggeration is approximately 1:1. a) Following

accretion of the Crescent and Pacific Rim terranes, subduction continued for some time with the subduction décollement lying at the

top of the subducting igneous crust. As a result, underthrust sediments were sheared at the interplate décollement forming the E

reflections. Some time later, the subduction thrust stepped down into the igneous oceanic crust, possibly as a result of reactivation of

a pre-existing fault. b) A slab, originally from the downgoing plate, was underplated to the base of the overlying continent. Continuing

subduction resulted in some deformation and shortening of the overlying slab and parts of the E region. A number of earthquakes

presently occur within the downgoing slab beneath the west coast of Vancouver Island. They may be the result of deformation in the

subducting slab as it is thrust beneath the underplated region.
The Cascadia Subduction Zone and Related Subduction Systems | 63

An image of the Cascadia subduction zone

across central Oregon

Stephane Rondenay and Michael G. Bostock

Earth and Ocean Sciences, The University of British Columbia, 6339 Stores Road, Vancouver, V6T 1Z4, Canada

rondenay@emma.geo.brown.edu, bostock@geop.ubc.ca

Between May 1993 and April 1994, researchers from

Oregon State University deployed an array of broadband,

three-component seismometers across central Oregon [Li,

1996]. The array recorded 32 high-quality teleseisms from

a total of approximately 70 receiver locations spaced at

five kilometer intervals and occupied at various times during the recording period. Conventional, receiver-function

processing of the data set by Li [1996] and Li and Nabelek

[1999] identified free-surface multiples as the most prominent signal from the subducting slab. However, because

of the effect of moveout from dipping structure, only forward modeling of individual events was performed.

We have developed a method of formally inverting

the teleseismic P-wave coda recorded on dense arrays of

receivers for underlying, two-dimensional elastic structure. The method accommodates both forward scattering of the incident wavefield and the back-scattering

response created by free-surface multiples. Recordings

from multiple earthquakes can, moreover, be inverted

simultaneously. We have applied this method to the

Oregon data set. We image the two principal structural

features inferred by Li [1996], notably: 1) the appearance of a prominent continental Moho some 150 km

from the coast line near 40 km depth; and 2) the lowvelocity oceanic crust of the Juan de Fuca plate dipping

at 10 ̊ beneath North America. In addition, however, we

image a disruption in the signature of subducted oceanic

crust which commences approximately 50 km from the

coastline, and which appears to coincide with an increase

in dip of the down-going plate. This disruption coincides

with the appearance of enhanced mantle conductivity

imaged in an earlier magnetotelluric (MT) experiment

(EMSLAB, e.g. Wannamaker et al., [1989]) along the same

profile.

The coincidence of seismic and MT anomalies is consistent with an interpretation involving dehydration reactions in the subducted crust occurring near 30 km depth.

Following Hyndman [1988], we infer that low velocities

above the oceanic plate manifest the upward expulsion

of hydrous fluids and their entrapment by an impermeable interface. This interface may represent the locus of

rehydration reactions producing an effective seal or it may

be structurally controlled. Similar interpretations of geophysical data from the northern Cascadia subduction zone

and central Andes strongly suggest that hydration reactions in the oceanic crust may play an important role in

controlling the physical properties in the overlying wedge.

References

Hyndman, R.D., Dipping seismic reflectors, electrically

conductive zones, and trapped water in the crust over

a subducting plate, J. Geophys. Res., 93, 13,391–

13,405, 1988.

Li, X.-Q., Deconvolving orbital surface waves for the

source duration of large earthquakes and modeling

the receiver function for the earth structure beneath

a broadband seismometer array in the Cascadia subduction zone, Ph.D. Thesis, Oregon State University,

Corvallis, Oregon, 1996.

Li, X.-Q., and J.L. Nabelek, Deconvolution of teleseismic

body waves for enhancing structure beneath a

seismometer array, Bull. Seismol. Soc. Am., 89, 190–

201, 1999.

Wannamaker, P.E., J.R. Booker, A.G. Jones, A.D. Chave,

J.H. Filloux, H.S. Waff, and L.K. Law, Resistivity cross

section through the Juan de Fuca subduction system

and its tectonic implications, J. Geophys. Res., 94,

14,127–14,144, 1989.
64 |
The Cascadia Subduction Zone and Related Subduction Systems | 65

Intraslab earthquakes beneath Georgia Strait/Puget Sound

Garry C. Rogers1 and Robert S. Crosson2

1 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

2 Geophysics Program, University of Washington, PO Box 351650, Seattle, Washington, 98195–1650, USA

rogers@pgc.nrcan.gc.ca, crosson@u.washington.edu

Observations of earthquakes in the northern half of the

Cascadia subduction zone can be divided into three time

periods: pre-instrumental from about 1850 and 1899,

early instrumental from about 1900 to 1969, and modern instrumental from about 1970 to present. The microearthquake pattern revealed in the modern instrumental

period, shows most of the seismicity in the slab to be in

the 40 to 60 km depth range beneath southern Georgia

Strait and Puget Sound (Figure 1). Some very small events

extend to maximum depths of about 90 km. The seismicity

maps out a gentle arch in the subducted plate opposite

the change in orientation of the coastline. The main concentration is in a band paralleling the coastline and

decreasing in intensity to the north and to the south.

Focal mechanisms are variable amongst the small events

but most of the larger events have a downdip tension

axis in common.

Most of the larger events (Table 1; Figure 2) occurred

in the early instrumental period from 1900 to 1970.

Locations are teleseismic in the early part of that period

but improve with time as the number of seismographs in

the region increases and timing becomes more accurate.

Depth control is poor for many events in this period, but

the low intensities in the epicentral region or lack of

aftershocks suggest that some events belong to the deeper

inslab suite of events. The locations of the larger (M>5)

events occur in the region of the modern day concentration of micro-earthquakes beneath southern Georgia

Strait and Puget Sound. The two largest events in 1965

(Mw=6.7) and 1949 (Mw=6.8) occurred near the south

end of the modern micro-earthquake concentration.

In the pre-instrumental period from about 1850 to

1899, it appears there were no events greater than magnitude 6.5. The population was sparse during the early

part of this period, but for a few widely felt earthquakes,

the lack of aftershocks or low intensities in the epicentral region suggest that they were subcrustal in-slab events.

Recent reviews of various aspects of seismicity in the

region have been published in the Decade of North American Geology (DNAG) volumes [Ludwin et al., 1991; Rogers

and Horner, 1991] and in papers contributed to volumes

published by the USGS [Ma et al., 1996] and the GSC

[Rogers, 1998]. These papers include discussions on inslab seismicity and a complete listing of publications.

References

Ludwin, R.S., C.S. Weaver, and R.S. Crosson, Seismicity

of Washington and Oregon, in Neotectonics of North

America, Decade Map Volume 1, edited by D.B.

Slemmons, E.R. Engdahl, M.D. Zoback and D.

Blackwell, pp. 77–98, The Geological Society of

America, Boulder, Colorado, 1991.

Ma, L., R.S. Crosson, and R. Ludwin, Western Washington earthquake focal mechanisms and their relationship to regional tectonic stress, in Assessing earthquake

hazard and reducing risk in the Pacific Northwest,

United States Geological Survey Professional Paper

1560, v.1, edited by A.M. Rogers, T.J. Walsh, W.J.

Kockelman and G.R. Priest, pp. 257–283, 1996.

Rogers, G.C., Earthquakes and earthquake hazard in the

Vancouver area, in Geology and Natural Hazards of

the Fraser River Delta, British Columbia, Geological

Survey of Canada, Bulletin 525, edited by J.J. Clague,

J.L. Luternauer and D.C. Mosher, pp. 17–25, Geological Survey of Canada, 1998.

Rogers, G.C. and R.B. Horner, An overview of western

Canadian seismicity, in Neotectonics of North America,

Decade Map Volume 1, edited by D.B. Slemmons,

E.R. Engdahl, M.D. Zoback and D. Blackwell, pp. 127–

130, The Geological Society of America, Boulder,

Colorado, 1991.
66 | Rogers and Crosson: Intraslab earthquakes beneath Georgia Strait/Puget Sound

Date Time Latitude Longitude Depth1 Magnitude2 Source

(yr/mon./day) (hr/min./sec.) (km)

1909 01 11 2349 48.7 -122.8 DEEP 6.0 no aftershocks, low I

1920 01 24 0710 0.0 48.6 -123.0 DEEP 5.5 no aftershocks

1939 11 13 0745 53.0 47.4 -122.6 DEEP 6.2 no aftershocks, low I

1943 11 29 0143 48.4 -122.9 DEEP 4.8 no aftershocks

1946 02 15 0314 50.0 47.3 -122.9 DEEP 6.4 UW publication

1949 04 13 1955 41.0 47.13 -122.95 54 6.9 publications

1960 09 10 1506 33.6 47.68 -123.18 60 4.9 PGC

1965 04 29 1528 43.6 47.4 -122.3 63 6.7 publications

1965 10 23 1628 3.0 47.5 -122.4 DEEP 4.8 ISC 64 km

1969 02 14 0833 36.0 48.94 -123.07 52 4.3 PGC/ISC

1976 05 16 0835 15.1 48.80 -123.35 60 5.4 PGC

1976 09 08 0821 2.0 47.38 -123.08 48 4.6 NEIC

1978 08 19 0151 18.8 48.64 -123.58 52 4.3 PGC

1983 08 28 1247 48.0 47.93 -122.85 51 4.1 UW

1989 03 05 0642 0.6 47.81 -123.36 46 4.6 UW

1989 06 18 2038 37.3 47.41 -122.78 45 4.1 UW

1999 07 03 0143 29.0 47.10 -123.50 41 5.7 UW

1. If there is a number in the depth column, it has been calculated by the agency in the last column or has been the object of a published

study. The word DEEP means it has been assigned to the deep suite for the reason listed in the last column.

2. Magnitude values are Mw when known.

TABLE 1: Larger known or suspected in-slab earthquakes beneath Georgia Strait and Puget Sound.
The Cascadia Subduction Zone and Related Subduction Systems | 67

FIGURE 1: Fifteen years of seismicity

within the oceanic plate. At the northernmost end of active Cascadia subduction, a band of seismicity exists

beneath the west coast of Vancouver Island and another concentration

exists beneath southern Georgia

Strait and Puget Sound. Circle size is

related to earthquake magnitude and

ranges from magnitude about zero for

the smallest circles to an offshore

event of magnitude 6.3. Data from the

Geological Survey of Canada and the

University of Washington.

FIGURE 2: Larger earthquakes within the subducting

plate. Green stars are magnitude 4 and greater (complete from 1951). Blue stars are 5 and greater (complete from 1899). Red stars are magnitude 6 and

greater events, complete from about 1865.
68 |
The Cascadia Subduction Zone and Related Subduction Systems | 69

Precise relocations of slab seismicity in the northern

Cascadia subduction zone

John F. Cassidy1 and Felix Waldhauser2

1 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

2 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

cassidy@pgc.nrcan.gc.ca, felix@usgs.gov

ABSTRACT

The Gulf Islands region of southwest British Columbia is

an ideal location to study, in detail, slab seismicity of the

subducting Juan de Fuca plate. Here, there is abundant

slab seismicity, a relatively dense network of seismograph

stations to obtain accurate hypocenters and some recently

deployed three-component broadband stations to constrain the velocity structure. We have relocated slab seismicity near 50–60 km depth in the Gulf Island region

using a “double difference” earthquake relocation algorithm designed to provide high-resolution relative hypocenter locations over large distances. The precise relative

locations obtained for an initial test set of 27 deep earthquakes in a narrow corridor near Saturna Island, B.C.,

indicates that most of the slab seismicity here is in a three

to four kilometer-thick region dipping at 20 ̊ to the northeast. There is some suggestion of a deeper parallel band

of seismicity (7–9 km) below the top of the upper band

that may be located in the upper mantle of the Juan de

Fuca plate.

Introduction

The Cascadia subduction zone is a seismically active

region (Figure 1) with three distinct types of earthquakes

[Rogers, 1998]: those in the continental crust (black dots

in Figure 1); those in the subducting oceanic plate (red

dots in Figure 1); and megathrust earthquakes that occur on the megathrust fault between the oceanic and

continental plates. Of the three types, those that occur

within the oceanic plate (“in-slab earthquakes”, see Figure 2) pose the greatest hazard to the major cities in the

region. The processes that cause these earthquakes and

that limit their maximum size and spatial distribution are

not well-understood.

We applied a new, high-resolution “double-difference”

earthquake location algorithm (Figure 3) of Waldhauser

and Ellsworth [2000] to obtain precise relative hypocentral locations for deep earthquakes (Figure 2) over

distances of up to 40–60 km near Saturna Island, B.C.

This is an ideal location to examine deep seismicity as

there is a relatively dense seismograph network (Figure

2) that can be used to obtain accurate earthquake hypocenters and for utilizing waveform cross-correlation techniques, as well as a recently deployed three-component

broadband station that can be used for structural studies. This allows us to examine the relationship between

the structure and well-resolved seismicity of the subducting

slab. In this study, we report on the results obtained

using the double-difference technique with catalog data

only, for test sets of up to 27 deep earthquakes.

Data and analysis method

For these tests we selected deep earthquakes (>40 km)

that lie within a 20-km wide corridor about 70 km long,

centered near the Saturna Island seismic station (SNB,

Figure 4). The initial locations (Figure 4) are taken from

the Canadian Earthquake Epicentre File and were determined using standard location methods, with P- and Swave arrival times from regional stations.

In the double-difference earthquake relocation

scheme, P- and S-wave differential travel times for all

possible pairs of events are minimized. Solid and open

circles represent trial hypocenters that are linked to

neighboring events by cross-correlation (solid lines) or

catalog data (dashed lines). For two events, i and j, the

initial locations (open circles) and corresponding slowness vectors s with respect to two stations, k and l, are

shown. Ray paths from the sources to the stations are

shown. Thick arrows (∆x) indicate the relocation vector

for events i and j obtained from the full set of equations.

Dt is the travel time difference between the events i and

j observed at stations k and l, respectively. For details of

the method, see Waldhauser and Ellsworth [2000].

Results and summary

The initial earthquake locations (Figure 4) show a band

of seismicity as much as 10–15 km thick, dipping to the

northeast.

Applying the double-difference earthquake location

scheme to these deep earthquakes in the vicinity of

Saturna Island, we find that the band of seismicity
70 | Cassidy and Waldhauser: Precise relocations of slab seismicity in the northern Cascadia subduction zone

narrows considerably, to become less than about five

kilometers thick and dipping at 20 ̊ to the northeast (Figure 5). There is the suggestion of two bands of seismicity,

separated by approximately five kilometers. It is noteworthy that the location of the upper band corresponds

with the low-velocity zone interpreted as the oceanic

crust beneath the Pacific Geoscience Centre (PGC)

[Cassidy, 1995]. This would suggest that the lower band

may fall within the uppermost mantle of the subducting

Juan de Fuca plate. Tests using various subsets of the data,

for example, the subset of events that lie within the 20

km x 30 km rectangular region shown in Figure 5, show

these relative hypocentral locations to be robust.

Acknowledgements

We thank Taimi Mulder for her assistance in collecting

the catalog data and Richard Franklin for his assistance

with figure preparation. Many diagrams were generated

using Generic Mapping Tools (GMS) [Wessell and Smith,

1995].

References

Cassidy, J.F., A comparison of the receiver structure beneath stations of the Canadian National Seismograph

Network, Can. J. Earth Sci., 32, 938–951, 1995.

Rogers, G.C., Earthquakes and earthquake hazard in the

Vancouver area, in Geology and Natural Hazards of

FIGURE 1: Cascadia seismicity during the period 1985–1999. Red dots indicate earthquakes within the oceanic plate, black dots are

earthquakes within the continental plate. Earthquake magnitudes range from 0 to 5.5. Cross section shows seismicity along a 50 kmwide corridor centered on the line A–B.
The Cascadia Subduction Zone and Related Subduction Systems | 71

FIGURE 2: Location of seismograph stations (squares are three-component broadband, triangles represent short-period vertical) and

the location of deep (>40 km depth) slab seismicity. Magnitudes range from 0 to 4.5.

FIGURE 3: Illustration of the double-difference earthquake relocation technique (from Waldhauser and Ellsworth [2000]). Solid and

open circles represent trial hypocenters that are linked to neighboring events by cross-correlation (solid lines) or catalog data (dashed

lines). For two events, i and j, the initial locations (open circles) and corresponding slowness vectors s with respect to two stations, k

and l, are shown. Ray paths from hypocenters to the stations are shown. Thick arrows (x) indicate the relocation vector for events i

and j. Dt is the travel-time difference between the events i and j observed at stations k and l, respectively.
72 | Cassidy and Waldhauser: Precise relocations of slab seismicity in the northern Cascadia subduction zone

the Fraser River Delta, British Columbia, Geological

Survey of Canada Bulletin 525, edited by J.J. Clague,

J.L. Luternauer, and D.C. Mosher, pp. 17–25, Geological Survey of Canada, 1998.

Waldhauser, F., and W.L. Ellsworth, A double-difference

earthquake location algorithm: method and application in the Northern Hayward Fault, California, Bull.

Seism. Soc. Am., 90, 1353–1368, 2000.

Wessell, P., and W.H.F. Smith, New version of the Generic Mapping Tools released, EOS Trans., 76, 329,

1995.

FIGURE 4: Initial locations of earthquake hypocenters in the vicinity of Saturna Island (SNB) shown in both map view (top) and cross

section (bottom) along line shown.
The Cascadia Subduction Zone and Related Subduction Systems | 73

FIGURE 5: Relocated earthquake hypocenters near SNB (Saturna Island). Cross sections show results for all events (lower left) and a

subset of events (lower right) closest to SNB (only the earthquakes within box shown on the map). The low-velocity layer was

interpreted [Cassidy, 1995] as the oceanic crust beneath the seismic station PGC (Pacific Geoscience Centre).
74 |
The Cascadia Subduction Zone and Related Subduction Systems | 75

Large earthquakes under our feet and deep beneath the

Puget lowland and Georgia Strait: How often do they occur,

how big can they get and what are their likely effects?

Stephen H. Kirby1, Kenneth C. Creager2 and Leiph A. Preston2

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

1 Department of Earth and Space Sciences, University of Washington, Seattle, Washington, 98195–1650, USA

skirby@usgs.gov, kcc@geophys.washington.edu, prestola@u.washington.edu

Several large deep earthquakes have occurred in the last

60 years beneath the Puget lowland in the urban corridor between Olympia, Washington and Vancouver, B.C.,

and they represent the most destructive type of seismic

event in western Washington State in this century. They

occur inside the Juan de Fuca plate subducting beneath

the North America plate at depths of 40 to 80 km. While

smaller than the big, shallow plate-boundary earthquakes, deep shocks occur much more frequently, are

closer to population centers and tend to be more

destructive for a given magnitude than plate-boundary

events. This presentation describes some of the effects

of large deep earthquakes in the western hemisphere in

the 20th century and some recent ideas of how to estimate better their maximum magnitudes, how often they

occur, and what their effects might be.
76 |
The Cascadia Subduction Zone and Related Subduction Systems | 77

In-slab seismicity at the north end of Cascadia

Garry C. Rogers1,2, Alison L. Bird1, John F. Cassidy1,2, Taimi L. Mulder1, Maiclaire K. Bolton2 and John P. Ristau2

1 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

2 Earth and Ocean Sciences, University of Victoria, PO Box 3055, Victoria, British Columbia, V8W 3P6, Canada

rogers@pgc.nrcan.gc.ca, bird@pgc.nrcan.gc.ca, cassidy@pgc.nrcan.gc.ca, mulder@pgc.nrcan.gc.ca,

bolton@pgc.nrcan.gc.ca, ristau@pgc.nrcan.gc.ca

The in-slab seismicity in the Canadian portion of the

Cascadia subduction zone differs from that to the south

in northern Washington State in two principal ways. First,

there is a very active band of seismicity along the coast

in the 25 to 35 km depth range where the angle of the

subducting plate goes from horizontal to a dip of about

20 ̊. Second, the deeper seismicity in the 40 to 60 km

range that dominates below Puget Sound, gradually

tapers off as the age of the plate being subducted becomes younger with the change in orientation of the

coastline from north–south along Washington and Oregon to northwest–southeast along British Columbia. The

focal mechanisms of the deeper suite are a mixture of

strike-slip and normal faulting events with downdip tension events predominating. Focal mechanisms beneath

the coast have a very confusing pattern and contain strikeslip, normal and thrust events with a wide variation in

the orientation of both P and T axis. No low-angle thrust

events have been detected that would be consistent with

movement of the Cascasdia subduction fault. The largest events in both the coastal region and the deeper

events beneath southern Georgia Strait have been in the

magnitude 5 to 6 range compared to larger events at the

south end of Puget Sound.

FIGURE 1: Fifteen years of seismicity within the oceanic plate. At the northernmost end of active Cascadia subduction, a band of

seismicity exists beneath the west coast of Vancouver Island and another concentration exists beneath southern Georgia Strait and

Puget Sound. Circle size is related to earthquake magnitude and ranges from magnitude about zero for the smallest circles to an

offshore event of magnitude 6.3. Data from the Geological Survey of Canada and the University of Washington.
78 |
The Cascadia Subduction Zone and Related Subduction Systems | 79

Warm-slab subduction as a global process

Stephen H. Kirby1, E. Robert Engdahl2 and Antonio Villaseñor2

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

2 Department of Physics, University of Colorado at Boulder, Boulder, Colorado, 80309, USA

skirby@usgs.gov, engdahl@lemond.colorado.edu, antonio@gldvisitor.cr.usgs.gov

The Cascadia subduction zone is a prime example of

“warm-slab” subduction. As with many such subduction

zones, the intraslab seismicity rate is low even though sizable and damaging slab earthquakes can occur in them.

This argues for a global approach to estimating seismic

hazards in these settings, and for investigations of modern and historical slab events and how seismic moment

release rates per trench length vary with slab input

parameters. We have identified 28 subduction zone settings that have warm-slab characteristics (Table 1). These

common features [Kirby et al., 1996] include: a) hypocentral depths that are largely shallow for intraslab events

(<100 to 125 km depth); b) shallow dip of the Wadati–

Benioff zone; c) feeble or absent volcanic arc [Simkin and

Siebert, 1984]; and d) the occasional occurrence of unusual lava compositions thought to represent melting of

eclogitic slab crust [Defant and Drummond, 1990].

Most of the subduction zone settings in Table 1 involve slow convergence rates, young plate age entering

the trench or low slab dip or some combination of these

attributes. These factors reduce the temperature contrast

between the slab and the surrounding mantle. Slow convergence rates may occur where the overall relative plate

motions are slow (<20 mm/a, such as the Makran and

South Shetland margins), where plate motions are very

oblique over a relatively straight trench (e.g., northwestern New Guinea), or where trench normals become very

oblique to plate motions near the ends of subduction

zones (e.g., the northern and southern Lesser Antilles,

the South Scotia arcs, and the western Aleutians). Examples of warm-slab environments where young lithosphere

is subducting include subduction zones where midocean

spreading ridges are just offshore (e.g., the Rivera subduction zone, the southeast segment of the Cocos system in Costa Rica, and the northwest part of the Cocos

system in Mexico) or where a young fossil backarc spreading system is being subducted (e.g., the Nankai/Shikoku

system in Japan and the central Luzon segment of the

northern Philippine system). Most of the systems of Table

1 also have shallow average dips where they are known

(δ<15–20°), including the avolcanic segments of the

classic flat-slab subduction in Peru and northern Chile.

The plate tectonics settings of these slab segments

are consistent with the contrast between their minimum

slab temperatures and the surrounding mantle being low.

This contrast is roughly proportional to the slab thermal

parameter [Kirby et al., 1991], given by:

φ =AV• cosθ_sinδ=AVn

•sinδ=AVz

f: A measure of slab “coldness” and is proportional to

the depth, Z, that v-shaped isotherms are advected

downward inside slabs. It has units of length.

A: Age of the plate at the trench

V: Speed of relative plate motion

Vn: Plate convergence speed

Vz: Slab vertical descent rate

q: Obliquity angle of plate motion to the trench normal

d: Slab dip angle

Most of the “warm-slab” subduction zones in Table 1

have thermal parameters less than about 500 km whereas

most subduction zones have thermal parameters exceeding 1000 km [Kirby et al., 1996].

Cascadia being placed in this “warm-slab” category

is important for earthquake hazard appraisal for several

potential reasons that stem from inferences drawn from

recent models of “warm-slab” subduction [Kirby et al.,

1996; Peacock and Wang, 1999; Peacock et al., this volume; Kirby, 2000]. Dehydration of the subducting crust

and densification to eclogite are expected to take place

at shallow depths (40–90 km) under the North American plate forearc, an expectation consistent with the fine

seismic wavespeed structure of the Nankai/ Shikoku system [Hori et al., 1985]. This shallow metamorphism

means that the forearc has an abundant deep source of

slab fluids that may be important in the mechanics of

upper plate forearc earthquakes that occur in abundance

above the intraslab events. Second, shallow liberation of

fluids from the slab should permit slab seismogenic faulting by dehydration embrittlement should slab stresses

be adequate. Dehydration embrittlement is a mechanism of unstable faulting in which the pore pressures

generated by dewatering of hydrous phases reduced the

effective pressures or normal stresses that tend to suppress brittle fracture and frictional sliding [Raleigh, 1967].

This mechanism of high-pressure faulting is a leading candidate for seismogenic faulting at depths greater than a
80 | Kirby, Engdahl and Villaseñor: Warm-slab subduction as a global process

few tens of kilometers. Third, densification of the

subducted crust should be a major cause of slab deformation and hence, shallow slab seismicity. Thus warmslab subduction can focus seismic activity at shallow

depths beneath the physiographically low ground in the

urban corridor from the Georgia Straight and the

Willamette Valley. This places the cultural geography of

humans in intersection with the geography of earthquakes

in Cascadia.

References

Defant, M.J., and M.S. Drummond, Derivation of some

modern arc magmas by melting of young subducted

lithosphere, Nature, 347, 662–665, 1990.

Hori, S., H. Inoue, Y. Fukao, and M. Ukawa, Seismic detection of the untransformed “basaltic” oceanic crust

subducting into the mantle, Geophys. J. R. Astron. Soc.,

83, 169–197, 1985.

Kirby, S.H., Invited news and views article: taking the

temperature of slabs, Nature, 403, 31-34, 2000.

Kirby, S.H., W.B. Durham, and L.A. Stern, Mantle phase

changes and deep earthquake faulting in subducting

lithosphere, Science, 252, 216–225, 1991.

Kirby, S.H., E.R. Engdahl, and R. Denlinger, Intermediateintraslab earthquakes and arc volcanism as physical

expressions of crustal and uppermost mantle metamorphism in subducting slabs, in Subduction: Top to Bottom, Geophysical Monograph 96, Bebout, edited by

G., D. Scholl, S. Kirby and J. Platt, pp. 195–214, American Geophysical Union, Washington, D.C., 1996.

Peacock, S., and K. Wang, Seismic consequences of warm

versus cool subduction metamorphism: examples

from southwest and northeast Japan Science, 286,

937–939, 1999.

Peacock, S., K. Wang, and A.M. McMahon, Thermal

structure and metamorphism of subducting oceanic

crust: insight into Cascadia intraslab earthquakes in

The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 123–126, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Raleigh, C.B., Tectonic implications of serpentinite weakening, Geophys. J. R. Astron. Soc., 14, 113–118, 1967.

Simkin, T., and L. Siebert, Explosive eruptions in space

and time: durations, intervals, and a comparison of

the world’s active volcanic belts, in Explosive

Volcanism: Inception, Evolution, and Hazards, edited

by Panel on Explosive Volcanism, pp. 110–121, National Academy Press, Washington, D.C., 1984.

1. Nankai/Shokoku (Philippine#): SW Japan Y^, Sh^

2. Cascadia (Juan de Fuca): WA, OR, BC Y, Sh

3. Southern Mexico (Rivera) Y

4. Southern Mexico (Cocos) Y, Sh (Flat slab)

5. Southeast Costa Rica (Cocos) Y, Sh

6. Northern Panama (Caribbean) Sh, Sl^

7. Northern Colombia (Caribbean) Sl

8. Northwest Colombia (Nazca) Y

9. Peru (Nazca) Sh (Flat slab)

10. Northern Chile & Argentina (Nazca) Sh (Flat slab)

11. Southern Chile (Nazca) Y

12. Tierra del Fuego (Antarctica) Y, Sh, Sl

13. South Shetlands (Antarctica) Y, Sh, Sl

14. S. Scotia, N & S Sectors (South America) Sl

15. New Zealand, Fiordland (Australia) Y, Sl

16. Western New Guinea (Caroline) Sl

17. Yap and Palau (Caroline) Sl

18. South Marianas (Pacific) Sl

19. Luzon (Eurasia, Borneo Microplate) Y

20. Western Aleutians (Pacific) Sl

21. Gulf of Alaska (Pacific) Sh

22. Wrangell (Pacific) Sh, Sl

23. Andaman/Nicobar (Indian) Y, Sl

24. Makran (Arabian) Sh, Sl

25. Cyprus/South Turkey (African) Sl

26. Hellenic (African) Sl

27. Calabrian (African) Sl

28. Lesser Antilles, N & S ends (N & S American) Sl

TABLE 1: Examples of “warm-slabs”*.

*Collision zones not included.

# Slab provenance plate indicated in parentheses.

Y^= young plate subducting (<20 Ma).

Sh^= shallow dip of the seismic zone (d <15 to 20°).

Sl^= slow convergence (<20 mm/a) either due to slow plate

motions or very oblique subduction.
The Cascadia Subduction Zone and Related Subduction Systems | 81

The structure and intraslab seismicity of the

Philippine Sea plate in Southwest Japan

Yoshiyuki Kaneda, Phil R. Cummins and Kouichi Uhira

Frontier Research Program for Subduction Dynamics, JAMSTEC

2–15 Natsushima-cho, Yokosuka,237–0061, Japan

kaneday@jamstec.go.jp, cummins@jamstec.go.jp

Tectonic setting of the Nankai Trough

The Nankai Trough in Southwest Japan is a convergent

plate boundary where the Philippine Sea plate (PSP) is

subducting beneath Southwest Japan at rate of about

4.5 cm/yr (Figure 1). Great megathrust earthquakes occur along this subduction zone with a recurrence interval of 100–200 years [Ando, 1975]. Hence, it is an area

of great concern in terms of seismic risk.

The age of the subducting PSP varies considerably

along strike in the Nankai Trough, because of the subduction of a fossil spreading ridge which began 15 Ma. As a

result the age of the PSP currently entering the Nankai

Trough subduction zone is thought to vary from about 15

Ma near the axis of the fossil spreading ridge to about 30

Ma in the Suruga Trough and off northern Kyushu.

A well-developed accretionary complex is present in

the Nankai Trough, composed of off-scraped and underplated trough fill turbidites and hemipelagic sediments.

This accretionary prism is over 100 km wide and 7 km

deep [Kodaira et al., 1999] and its landward extremity

abuts another, older accretionary complex known as the

Shimanto belt.

Seismic surveys and constraints on slab structure

After the 1995 Kobe earthquake, a number of research

initiatives aimed at earthquake disaster mitigation were

established in Japan. One of these was the Frontier

Research Program for Subduction Dynamics (Frontier)

established at the Japan Marine Science and Technology

Center (JAMSTEC), with the goal of studying subductionzone earthquakes. With this goal in mind, Frontier

began a vigorous program of marine seismic surveys, part

of which concentrated on the Nankai Trough.

Several of the ocean bottom seismometer (OBS) survey lines shot in the Nankai Trough are illustrated in Figure 1. These were run in combination with multi-channel

seismic surveys to provide constraints on structure down

to about 30 km depth. The results of these surveys are

illustrated in Figure 2, where cross sections from the easternmost survey to the westernmost are arranged from

top to bottom. As can be seen from the figure, there is

considerable along-strike variation in structure. The

accretionary prism thickens considerably in the central

part of the Nankai Trough (KR9704), while contact between the island arc crust and the subducting plate

covers a wider downdip segment of the plate boundary

in the western (KR9806) and eastern (KR9810) cross

sections.

We have used the results of our surveys as well as

those of other research groups in Japan to construct a

model of the plate boundary in the Nankai Trough, as

illustrated in Figure 3. All the survey lines used to construct the model are illustrated. At shallow depth the

model was constrained to agree with these survey lines,

while at deeper depth the shape of the plate boundary

was constrained to follow that of the intraslab seismicity

[Nakamura et al., 1997]. The contours of the resulting

plate boundary model (Figure 3) show substantial contortion of the plate. Dip is shallow in the central part of

the Nankai Trough but is steeper to the east and west.

Intraslab seismicity of the Philippine Sea plate

We have used seismic source mechanism determinations

of the Japanese Meteorological Agency (JMA) to study

the intraslab seismicity of the PSP. These data include

earthquakes occurring throughout Japan during the years

1926–1997. For those earthquakes in Southwest Japan,

we used the plate boundary model described above to

separate those events which occurred in the upper plate

from those which occurred in the subducting PSP. Known

interplate thrust events were omitted, as were some

events which occurred so close to the plate boundary

that they could not be unambiguously assigned to either

the upper or the lower plate.

Earthquakes occurring in the island arc forearc of

Southwest Japan are shown in the upper part of Figure

4. These are predominantly strike-slip events with P axes

oriented east–west to northwest–southeast. These are

thought to reflect a forearc stress field that is close to

margin parallel [e.g., Seno, 1999], and Wang and He

[1999] have shown that this suggests a very weak
82 | Kaneda, Cummins and Uhira: Structure and intraslab seismicity of the Phillipine Sea plate in Southwest Japan

megathrust in the Nankai Trough.

The intraslab events, shown in the bottom part of

Figure 4, span a range of depths whose along-strike maximum increases with the age of the subducting plate:

beneath northern Kyushu seismicity extends to 160 km

depth, and beneath the Tokai region to 90–100 km

depth, while beneath Shikoku the seismicity only extends

to 60 km maximum depth. These depths correlate

roughly with the maximum depth at which hydrous

phases in the subducting plate remain stable [Peacock

and Wang, 1999], and are thus consistent with the dehydration embrittlement model for intraslab earthquakes

[Kirby et. al, 1996].

The pattern of focal mechanisms of PSP intraslab

earthquakes is complicated but in general follows the

contortions of the subducting plate, as suggested by

Imagawa et al. [1985].

Conclusions

Recent seismic survey results in the Nankai Trough, in

combination with seismicity data, have revealed detailed

structure of the highly contorted Philippine Sea plate

subducting beneath Southwest Japan. This has facilitated

the discrimination of intraslab from upper plate earthquakes in order to study the intraslab seismicity of the

PSP. This maximum depth of seismicity has a pattern

consistent with that expected from dehydration embrittlement of the PSP and the along-strike age variation of the

subducting PSP.

Considerations of seismic risk in the Nankai Trough

generally focus on the great megathrust earthquakes

which repeatedly occur there. All of the recent intraslab

seismic activity is associated with relatively small earthquakes which have caused little damage. Early in this

century, however, a few large intraslab earthquakes occurred which caused considerable damage and some

FIGURE 1: Tectonics and seismic survey lines in the Nankai Trough. The survey lines of three JAMSTEC ocean bottom seismometer

experiments are indicated, along with plate convergence rate and direction, and the age variation of the Philippine Sea plate.
The Cascadia Subduction Zone and Related Subduction Systems | 83

FIGURE 2: Three cross sectional models of seismic velocity structure, corresponding to each of the survey lines in Figure 1. Note the

changes in plate boundary dip, and area of contact between the island arc crust and the subducting oceanic plate.
84 | Kaneda, Cummins and Uhira: Structure and intraslab seismicity of the Phillipine Sea plate in Southwest Japan

loss of life. What little information is available concerning these earthquakes will be presented in this talk and

in a companion poster session.

References

Ando, M., Source mechanisms and tectonic significance

of historical earthquakes along the Nankai Trough,

Japan, Tectonophyics, 27, 119–140, 1975.

Kirby, S., E.R. Engdahl, and R. Denlinger, Intermediatedepth intraslab earthquakes and arc volcanism as physical expressions of crustal and uppermost mantle

metamorphism in subducting slabs, in Subduction: Top

to Bottom, edited by G.E. Bebout et al., pp. 195–214,

American Geophysical Union, Washington, D.C., 1996.

Kodaira, S., N. Takahashi, J. Park, K. Mochizuki, M.

Shinohara, and S. Kimura, The western Nankai Trough

seismogenic zone: results from wide-angle ocean-bottom seismographic survey, J. Geophys. Res., 105,

5887–5905, 1999.

Imagawa, K., K. Hirahara, and T. Mikumo, Source mechanisms of subcrustal and upper mantle earthquakes

around the northeastern Kyushu Region, southwestern Japan, and their tectonic implications, J. Phys.

Earth, 33, 257–277, 1985.

Nakamura, M., H. Watanabe, T. Konomi, S. Kimura, and

K. Miura, Characteristics activities of subcrustal earthquakes along the outer zone of southwestern Japan,

Ann. Disas. Prev. Res. Inst., Kyoto U., 40, 1–21, 1997.

Peacock, S., and K. Wang, Seismic consequences of warm

versus cool subduction metamorphism: examples

from Southwest and Northeast Japan, Science, 286,

937–939, 1999.

Seno, T., Synthesis of regional stress fields of the Japanese islands, The Island Arc, 8, 66–79, 1999.

Wang, K., and J. He, Mechanics of low-stress forearcs:

Nankai and Cascadia, J. Geophys. Res., 104, 15,191–

15,205, 1999.

FIGURE 3: A model for the depth to the top of the subducting Philippine Sea plate based on seismic survey results (JAMSTEC and nonJAMSTEC) and seismicity data.
The Cascadia Subduction Zone and Related Subduction Systems | 85

FIGURE 4: Japanese Meteorological Agency earthquake source mechanism solutions. The model from Figure 3 has been used to

separate this into a) earthquakes in the island arc crust and b) intraslab earthquakes.
86 |
The Cascadia Subduction Zone and Related Subduction Systems | 87

Intraslab earthquakes in

the subducting oceanic plates below Mexico

Shri K. Singh, Vladimir Kostoglodov and J. F. Pacheco

Instituto de Geofísica, UNAM, CU, 04510 Mexico, DF, MEXICO

krishna@ollin.igeofcu.unam.mx, vladimir@ollin.igeofcu.unam.mx

Introduction

Moderate and large intraslab earthquakes in the

subducted Cocos plate below Mexico are relatively frequent. For example, four of the seven large (Mw≥ 6.9)

earthquakes which occurred in Mexico between 1995

and 1999 were intraslab earthquakes. Two of these four

earthquakes also caused damage and deaths in towns

near the epicentral region. Historically, some of the

intraslab earthquakes have devastated cities and towns

located in the Mexican altiplano. Recent intraslab earthquakes have been well recorded by local and regional

seismographs and accelerographs. The analysis of these

data, along with previous results, has improved our

knowledge of the Benioff zone, source process of intraslab earthquakes, and expected ground motions from

such events. In this overview of seismic investigations of

intraslab earthquakes in Mexico, we will limit ourselves

to the region between Tehuantepec in the southeast and

Jalisco in the northwest (94 ̊–105.5 ̊W). The region is

characterized by subduction of young slab (age less than

about 16 ma).

Plate tectonic framework and intraslab earthquakes

Figure 1 summarizes the plate tectonic framework of the

region. The figure gives age [Kostoglodov and Bandy,

1995] and relative plate convergence rate at the trench

according to NUVEL 1A model [DeMets et al., 1994],

epicenters and focal mechanisms of all well-located intraslab events, isodepth contour of the Benioff zone [Singh

and Mortera, 1991; Pardo and Suárez, 1995], and the

location of active volcanoes. The earthquakes shown in

the figure are based on a catalog of intraslab events (H≥ 35

km) which we have compiled from different sources. This

catalog contains all intraslab events for which focal mechanism is available. It is probably complete for Mw≥ 6.0

since 1963, and for Mw≥ 5.4 since 1976, a period for

which Harvard centroid-moment tensor (CMT) solutions

are available. Table 1 lists the largest earthquakes

(M≥ 6.9) from this catalog. Figure 1 also illustrates the

geometry of the Benioff zone along selected sections

[modified from Singh and Mortera, 1991; Suárez et al.,

1990; Singh and Pardo, 1993; Pardo and Suárez, 1995;

Kostoglodov et al., 1996]. From the figure, it may be noted

that the epicenters of the intraslab earthquakes cease

before reaching the volcanic front, and there is a significant along-strike variation of the geometry of the Benioff

zone and intraslab seismicity. In Jalisco, where the Rivera

plate subducts below Mexico with a relatively low convergence rate (NUVEL 1A), the Benioff zone has a usual

shape with a dip angle of ~45 ̊ [Bandy et al., 1999] (not

shown in Figure 1). However, intraslab events apparently

do not exceed Mw=5.0 there and our catalog contains

no such event. In Michoacan–Guerrero and in Guerrero

segments, the subducted Cocos plate bends with a small

dip, unbends becoming subhorizontal and reaching a

depth about 50 km at a distance of about 250 km from

the trench. Subduction becomes normal in the Tehuantepec segment.

Source parameters of recent intraslab earthquakes

Until recently it was thought that large intraslab events

below Mexico occurred downdip from the strongly coupled part of the interface at D>150 km. The great Oaxaca

earthquake of January 15, 1931 (M=7.8), a normal-faulting intraslab event which caused great destruction to the

City and the State of Oaxaca, seemed to be an exception

since it was located only 65 km from the coast [Singh et

al., 1985]. Recent earthquakes, however, suggest that large

intraslab events, which occur below or close to the downdip edge of the coupled plate interface, may be more

frequent than previously thought. For example the earthquakes of January 11, 1997 (Mw=7.1) occurred just

below the rupture area of the shallow-thrust Michoacan

earthquake of September 19, 1985 (Mw=8.0) [Mikumo

et al., 1999]. Similarly, the intraslab Oaxaca earthquake

of September 30, 1999 (Mw=7.4) occurred just below

the coast, near the downdip edge of the aftershock area

of the November 29, 1978 (Mw=7.7) thrust event [Singh

et al., 2000a]. Near-source and regional records of these

two earthquakes have been inverted to map slip history

on the fault. Both events show a strong directivity along

the strike, which parallels the coast, with a downdip component and relatively small slip near the hypocenter.
88 | Singh, Kostoglodov and Pacheco: Intraslab earthquakes in the subducting oceanic plates below Mexico

During the 1997 event, the rupture initiated at a depth

of 34 km, with most of the slip confined in the depth

range of 30–50 km [M. Santoyo, pers. comm., 2000].

The 1999 rupture initiated at a depth of 40 km and most

of the slip occurred within the depth range of 35–52 km

[Hernandez et al., 2000]. Since the plate interface in the

region lies at a depth of 15–20 km, we conclude that

these two events occurred in the oceanic mantle of the

subducted Cocos plate.

Aftershocks from intraslab earthquakes

Aftershock activity following intraslab earthquakes in

Mexico is very low. For example, only one aftershock

with Mw≥ 5.4 is reported for the 1997 event, while none

at this magnitude level occurred following the two large

earthquakes of 1999 listed in Table 1. Although a quantitative comparison is not available at present, at the

smaller magnitude level there is considerable variation

in the number of aftershocks from one shock to another.

For instance, the earthquake of 1973 (Mw=7.0, H=82

km) was followed by few aftershocks. The earthquake of

1980 (Mw=7.0, H=65 km) gave rise to a large number

of aftershocks. About 900 aftershocks were recorded by

a portable seismic network within nine days after the

mainshock. Of these, 300 events could be located

[Yamamoto et al., 1984]. The June 15, 1999 earthquake

(Mw=6.9, H=61 km) generated 33 aftershocks (2.3≤ M

≤ 3.7) within seven days. The 1999 earthquake (Mw=

7.1, H=34 km) produced far fewer aftershocks than the

September 30, 1999 earthquake (Mw=7.4, H=40 km).

To a first approximation, the number of aftershocks seems

to decrease with depth.

Relationship between intraslab

and interplate earthquakes

For the Mexican subduction zone, Lay et al. [1989]

reported that the large, shallow thrust events were preceded by tensional downdip intraslab events, in accordance with the general worldwide pattern, but a decrease

in the tensional events following the rupture of the interface was not clear from the available data. Singh et al.

[2000b] have analyzed intraplate events which occurred

in the subducting Cocos plate below northwest Guerrero

and Michoacan since 1964 (Mw≥ 5.4) and their possible

relationship with a sequence of four large/great shallow,

thrust earthquakes (7.3≤ Mw≤ 8.0) which ruptured the

entire plate interface along this part of the coast during

1979–1985 (a length of ~210 km). This sequence provides an excellent opportunity to study temporal evolution of intraslab seismicity during about 15 years before

and after the stress release caused by rupture of the plate

interface. The results are in agreement with the model

involving locking and unlocking of the coupled plate interface. A detailed study of the earthquake of 1997

(Mw=7.1) also shows that it occurred in the region of

maximum coseismic stress increase caused by the 1985

thrust earthquake [Mikumo et al., 1999]. Recent numerical modeling by Gardi et al. [2000] demonstrates that the

peculiar geometry of the Benioff zone in the region

strongly affects the state of intraslab stress and, hence,

the pattern of seismicity.

There is also evidence that in Mexico some large,

downdip, tensional earthquakes follow shallow thrust

earthquakes rather than precede it. For example, the 1931

earthquake of Oaxaca (M=7.8), a normal faulting,

downdip intraplate slab occurred after four thrust events

(7.4≤ M ≤ 7.8) ruptured the plate interface in the region

in 1928 [Singh et al., 1985].

Some damaging intraslab events

If Mexico City, with its abnormal soil characteristics and

large population, is excluded from the statistics, then the

damage and the loss of lives from interplate and intraslab

events in Mexico may be about equal, even though the

former type of earthquakes is more frequent than the

latter type. Some highly damaging intraslab earthquakes:

the earthquake of 1858 (M~7.7) which caused severe

damage to towns in Michoacan, including the City of

Date Time Latitude  ̊N Longitude  ̊E Depth (km) M Strike Dip Rake

19310115 1:50:40 16.340 -96.870 40 7.8 270 56 -90

19370726 3:47:13 18.450 -96.080 85 7.3 313 70 -95

19640706 7:22:13 18.215 -100.480 55 7.2 296 38 -63

19730828 9:50:41 18.250 -96.550 82 7.3 326 50 -76

19801024 14:53:35 18.030 -98.270 65 7.1 311 26 -66

19970111 20:28:26 18.340 -102.580 34 7.1 175 18 -28

19990615 20:42:04 18.130 -97.540 61 6.9 309 40 -83

19990930 16:31:13 16.030 -96.960 40 7.4 102 42 -103

TABLE 1: Large intraslab earthquakes of Mexico with known focal mechanism.
The Cascadia Subduction Zone and Related Subduction Systems | 89

FIGURE 1: Tectonic map of the region. Triangles indicate active volcanoes. All intraslab earthquakes with known focal mechanism are

shown. Relative plate convergence rates are given near arrows along the trench. The numbers above the arrow heads indicate the age

(Myr) of the oceanic lithosphere at the trench. Isodepth contours are shown by continuous lines when well defined by seismicity and by

dashed lines when inferred. Hypocenters of intraslab earthquakes projected on sections A to D are shown on the four graphs.
90 | Singh, Kostoglodov and Pacheco: Intraslab earthquakes in the subducting oceanic plates below Mexico

Morelia and Mexico City; the 1931 earthquake which

caused major destruction in the State and City of Oaxaca;

the Orizaba and Huajuapan de Leon earthquakes of

1973 and 1980; and the 1999 earthquake of Tehuacan

which caused unexpectedly large damage to colonial

structure in the City of Puebla. Part of the reason for

large damage during intraslab earthquakes is the high

population density in the Mexican altiplano.

Ground motion from intraslab earthquakes

A comparison of intensities during intraslab and interplate

earthquakes in Mexico suggest that the ground motions,

and hence the stress drop, involved in the former type

of earthquakes is higher than in the latter type. Extensive

recordings from recent intraslab events permits regression analysis of different ground motion parameters.

Based on hard-site, strong-motion recordings of ten

intraslab earthquakes (5.4≤ Mw≤ 7.4; 40≤ H≤ 65 km;

R≤ 400 km) we obtain the following preliminary relation

between horizontal Amax (in gal), Mw, and hypocentral

distance R (in kilometers):

log Amax= -0.148 + 0.623 Mw – log R - 0.0032 R

The corresponding relation for interplate Mexican earthquake is [Ordaz et al., 1989]:

log Amax = 1.76 + 0.300 Mw – log R - 0.0031 R

Note that the preliminary regression suggests a larger

dependence of Amax on Mw for intraslab events than for

interplate events. Figure 2 shows predicted Amax for

Mw=7.0 and 7.5 for intraslab and interplate earthquakes.

The expected Amax for an intraslab Mw=7.5 event is

four to five times greater than the corresponding value

for an interplate event. This difference decreases for

Mw=7.0. The figure also shows recorded Amax data

during the earthquakes of January 11, 1997 (Mw=7.1),

June 15, 1999 (Mw=6.9) and September 30, 1999

(Mw=7.4). The large scatter in the data is partially due

to the directivity of the sources, which has not been

accounted for in the regression.

Research needs

• A reliable catalog for realistic estimation of seismic

hazard. It is possible to improve the source parameters

of moderate and large intraslab events between 1906

and 1963 by analyzing seismograms and bulletins.

• A denser seismic network to detect small, intraslab

events to better define the intraslab seismicity, and the

geometry and the extent of the Benioff zone. How close

to Mexico City and other important cities in the Mexican altiplano can large intraslab earthquakes occur?

FIGURE 2: The peak horizontal acceleration (Amax) data versus hypocentral distance R for the last three large intraslab earthquakes of

Mexico (Table 1). Also shown are expected Amax values, based on regression analysis, for Mw=7.5 and 7.0 earthquakes; continuous

lines: intraslab earthquakes, dashed lines: interplate events.
The Cascadia Subduction Zone and Related Subduction Systems | 91

• A model to explain along-strike variation of the Benioff

zone and the intraslab seismicity, and the cause of large

intraslab earthquakes which occur just below the coupled plate interface at a depth of about 40 km.

• A resolution of the controversy regarding the origin of

the volcanism in Mexico (see, e.g., Verma, [1999] and

the references therein). Is there a direct association

between volcanism and subduction? Or is volcanism

related to past and present rifting and faulting?

References

Bandy W., V. Kostoglodov, A. Hurtado-Díaz,and M.

Mena, Structure of the southern Jalisco subduction

zone as inferred from seismicity and gravity, Geofisica

Internacional, 38, 127–136, 1999.

DeMets, C., R.G. Gordon, D.F. Argus, and S. Stein, Effects

of recent revisions to the geomagnetic reversal time

scale on estimates of current plate motions, Geophys.

Res. Lett., 21, 2191–2194, 1994.

Gardi, A., M. Cocco, A.M. Negredo, R. Sabadini, and

S.K. Singh, Dynamic modeling of the subduction zone

of central Mexico, Geophys. J. Int., 143, 809–820,

2000.

Hernandez, B., N.M. Shapiro, S.K. Singh, J.F. Pacheco, F.

Cotton, M. Campillo, A. Iglesias, V. Cruz, J. M. Gomez,

and L. Alcántara, Rupture history of September 30,

1999 intraplate earthquake of Oaxaca, Mexico

(Mw=7.5) from inversion of strong-motion data in the

frequency domain, Geophys. Res. Lett., 28, 363–366,

2001.

Kostoglodov, V., and W. Bandy, Seismotectonic constraints

on the convergence rate between the Rivera and

North American plates, J. Geophys. Res., 100, 17,977–

17,989, 1995.

Kostoglodov, V., W. Bandy, J. Domínguez and M. Mena,

Gravity and seismicity over the Guerrero seismic gap,

Mexico, Geophys. Res. Lett., 23, 3385–3388, 1996.

Lay. T., L. Astiz, H. Kanamori, and D.H. Christensen,

Temporal variation of large intraplate earthquakes in

coupled subduction zones, Phys. Earth Planet. Interiors, 54, 258–312, 1989.

Mikumo, T., S.K. Singh, and M. Santoyo, A possible stress

interaction between large thrust and normal faulting

earthquakes in the Mexican subduction zone, Bull.

Seism. Soc. Am., 89, 1418–1427, 1999.

Ordaz, M., J.M. Jara, and S.K. Singh, Riesgo sísmico y

espectros de diseño en el estado de Guerrero, Mem.

VIII Congreso Nac. de Ingeniería Sísmica, Acapulco,

Mexico, D40–D56, 1989.

Pardo, M., and G. Suárez, Shape of the subducted Rivera

and Cocos plates in southern Mexico: seismic and

tectonic implications, J. Geophys. Res., 100, 12,357–

12,373, 1995.

Singh, S.K., and F. Mortera, Source-time functions of large

Mexican subduction earthquakes, morphology of the

Benioff zone and the extent of the Guerrero gap, J.

Geophys. Res. 96, 21,487–21,502, 1991.

Singh, S.K., G. Suárez, and T. Domínguez, The great

Oaxaca earthquake of 15 January 1931: lithosphere

normal faulting in the subducted Cocos plate, Nature,

317, 56–58, 1985.

Singh, S.K. and M. Pardo, Geometry of the Benioff zone

and the state of stress in the overriding plate in central

Mexico, Geophys. Res. Lett., 20, 1483–1486, 1993.

Singh, S.K., M. Ordaz, J.F. Pacheco, R. Quaas, L.

Alcántara, S. Alcocer, C. Gutierrez, R. Meli, and E.

Ovando, A preliminary report on the Tehuacán,

México earthquake of June 15, 1999 (Mw=7.0),

Seism. Res. Lett., 70, 489–504, 1999.

Singh, S.K., J. Pacheco, T. Mikumo, and V. Kostoglodov,

Intraplate earthquakes in central México and their

relationship with large/great intraplate earthquakes,

Seism. Res. Lett., 71, 239, 2000a.

Singh, S.K., M. Ordaz, L. Alcántara, N. Shapiro, V.

Kostoglodov, J.F. Pacheco, S. Alcocer, C. Gutierrez,

R. Quaas, T. Mikumo, S. Alcocer, and E. Ovando,

The Oaxaca earthquake of 30 September 1999

(Mw=7.5): a normal-faulting event in the subducted

Cocos plate, Seism. Res. Lett., 70, 489–504, 2000b.

Suárez, G., T. Monfret, G. Wittlinger, and C. David, Geometry of subduction and the depth of the seismogenic zone in the Guerrero gap, Nature, 345, 336–

338, 1990.

Verma, S.P., Geochemistry of evolved magmas and their

relationship to subduction-unrelated mafic volcanism

at the volcanic front of the central Mexican volcanic

belt, J. Volcanol. Geotherm. Res., 93, 151–171, 1999.

Yamamoto, J., Z. Jimenéz, and R. Mota, El temblor de

Huajuapan de Leon, Oaxaca, México, del 24 de

Octubre de 1980). Geofis. Internacional, 23, 83–110,

1984.
92 |
The Cascadia Subduction Zone and Related Subduction Systems | 93

Seismicity and seismotectonics of Central America

with emphasis on intraslab seismic activity

Marino Protti

Observatorio Vulcanológico y Sismológico de Costa Rica, Universidad Nacional, Costa Rica

jprotti@una.ac.cr

Introduction

Although no large (Mw>7.5) intraslab earthquakes have

ever been recorded under the Central American arc,

small and moderate-size intermediate depth events are

frequent in this region. These events provide the key

information needed to map the geometry of the Wadati–

Benioff zone and have given important insights for the

understanding of tectonic processes at this subduction

zone. In Central America, and especially under Nicaragua and Costa Rica, it has been possible to obtain, in

great detail, the geometry of the subducted slab and to

make correlations with the structure and genesis of the

subducting Cocos plate.

Tectonic setting of Central America

Central America is located on the western margin of the

Caribbean plate (Figure 1). Here, the Cocos plate subducts

beneath the Caribbean plate along the Middle America

trench at a rate between 70 and 95 mm/year from Guatemala to southern Costa Rica, respectively (computed

from DeMets et al. [1990]). The northern boundary

between the Caribbean and the North American plate,

in Central America, is the Chixoy–Polochic–Motagua fault

system.

The subduction of young, warm lithosphere, together

with the collision of the Cocos Ridge in southern Costa

Rica, produced the fracturing of the southwestern part of

the Caribbean plate and the formation of a microplate,

the Panama block. This microplate includes the southeast

portion of Costa Rica and all of Panama. The northern

boundary between the Panama block and the Caribbean

plate is a convergent margin, the Panama thrust belt, that

extends from the Caribbean coast of Colombia to south

of Limón, Costa Rica [Silver et al., 1990]. Although a trench

has not yet developed, this convergent margin produces

large underthrust earthquakes and moderate size intermediate-depth seismic activity. In Costa Rica, the boundary between the Caribbean plate and the Panama block

consists of a diffuse left-lateral shear zone from Limón to

the middle American trench.

South of the border between Costa Rica and Panama

is the Panama fracture zone. This right-lateral transform

fault is the plate boundary between the Cocos and Nazca

plates.

Seismicity of Central America

Seismically, Central America is a very active region with

sources of different genesis and depths. Shallow events

(Z<30 km) occur: a) associated with the subduction of

the Cocos plate under the Caribbean plate and Panama

block; b) along the Chixoy–Polochic–Motagua fault

system; c) along the Panama fracture zone; d) as intraplate

faulting of the Panama block and the Cocos and Caribbean plates; e) as interplate activity between the Caribbean plate and the Panama block, both along the Panama

thrust belt and along a shear zone across central Costa

Rica; and f) associated with the volcanic arc.

Intermediate-depth earthquakes in Central America

occur as internal deformation of both the subducted

Cocos plate under the Caribbean plate and Panama

block, and the subducted Caribbean plate under the

Panama block, in northern Panama. The maximum depth

of intraslab earthquakes along the middle America trench

decreases from near 300 km under the border between

Guatemala and El Salvador, to around 50 km in southern Costa Rica.

Intermediate-depth seismicity of Central America

The intermediate-depth seismic activity in Central America

shows a very steeply dipping slab under Guatemala, El

Salvador and Nicaragua. Under northern Costa Rica the

angle begins to shallow and the seismic slab disappears in

southern Costa Rica.

Protti [1991] relocated all intraslab seismic activity

recorded by the Costa Rica Volcanological and Seismological Observatory (OVSICORI) under Costa Rica from

1984 to 1989 and constructed first-motion focal mechanisms for individual events, and composite focal mechanisms for clusters of events. He found that vertical to

subvertical P axes dominate at depths below 75 km and

vertical to subvertical T axes dominate at greater depths.

An analysis of P and T axes of intermediate-depth

seismicity recorded under Central America in the Harvard
94 | Protti: Seismicity and seismotectonics of Central America with emphasis on intraslab seismic activity

catalog was later made by Guendel and Protti, [1998].

They grouped the intraslab seismicity in three depth intervals (50–75 km, 75–100 km and >100 km) for two

regional areas: under Guatemala–El Salvador and under

Nicaragua–Costa Rica. In general, no major differences

exist in the orientation of the P axes nor in focal mechanisms for these regions and depth intervals (Figure 2). The

only important feature found in this study is the bimodal

distribution in the orientation of P axes for events with

focal depths between 75 and 100 km under Guatemala–

El Salvador.

Geometry of the Wadati–Benioff zone under

Nicaragua and Costa Rica, and its correlation with

the age of the subducted Cocos plate

The intermediate-depth seismicity under central Costa

Rica reveals a tear on the subducted Cocos plate (the

Quesada sharp contortion of Protti et al. [1995]) which

is recognizable below 70 km (Figure 3). To the northwest, the Cocos plate dips at 80° and seismicity reaches

depths of 220 km beneath the Nicaragua border to 135

km at the tear. To the southeast, the Cocos plate dips at

60° and the seismicity does not exceed 125 km [Protti et

al., 1995]. The projection of the tear to the trench coincides with an abrupt change in the bathymetry of the

Cocos plate along a northeast alignment (the rough/

smooth boundary of Hey [1977]). This bathymetry change

marks the boundary between lithosphere of the Cocos

plate created along the East Pacific Rise and that created

along the Galapagos spreading centre (Figure 4). The

northeast extension of the rough/smooth boundary, which

becomes the Quesada sharp contortion after subducting,

was interpreted as a relic of the original transform fault

along which the Farallon plate broke into Cocos and

Nazca plates. It is the contrast in age, and therefore density, which produces the tear at this boundary. The first

seismic event with a clear tear focal mechanism occurred

on March 7, 1992, precisely where the Quesada sharp

contortion begins to manifest. This earthquake, the

Naranjo earthquake, had a magnitude Mw=6.6 and a

focal depth of 78 km. No important damage was reported from this event; only falling objects.

FIGURE 1: Tectonic setting of Central America.
The Cascadia Subduction Zone and Related Subduction Systems | 95

Also in central Costa Rica, and almost parallel to the

84 ̊W of longitude, exists another abrupt change in age

of the subducting Cocos plate. This age contrast occurs

due to the subduction of the transform fault that offsets

the Costa Rica rift to the west from the Ecuador rift to

the east. The subduction of this feature produces the

abrupt termination, to the east, of intraslab seismicity.

There, the subducting slab may be so young and warm,

that its deformation is mainly plastic. Also, the buoyancy

and very low angle of subduction inferred for this youngest segment do not permit the penetration of the slab

into lower levels of the asthenosphere, inhibiting the internal phase changes in the slab that have been suggested

as a cause of intraslab seismicity.

Potential seismic hazard from intermediatedepth earthquakes in Central America

Because of the relatively large trench-arc distance in Guatemala, El Salvador and Nicaragua, together with the very

steep angle of the slab, most of the intermediate-depth

events occur either offshore or under the Pacific coast of

these countries. Given the depth and the relatively few

cities or towns along the Pacific coast of these countries,

the potential seismic hazard from intermediate-depth

earthquakes in that region is very low.

On the other hand, since under northern Costa Rica

the angle of the slab begins to shallow and the trench-arc

gap is smaller, intermediate-depth earthquakes tend to

have epicentres not only shallower than in the rest of

Central America, but right under important concentrations of population and infrastructure, increasing the seismic risk from this kind of activity in Costa Rica. Nevertheless, no large (Mw>7.0) intermediate earthquake has been

recorded under Costa Rica.

The largest intermediate-depth event recorded in the

Harvard catalog for Central America occurred on June

19, 1982 52 km off El Salvador and had a magnitude

Mw=7.3. An earthquake with magnitude M=7.0 and 80

km deep occurred under central Costa Rica on November 19, 1948. Although global catalogs locate this event

where no seismic slab has been identified with local data,

damage reports locate it further west right where the

FIGURE 2: Rose diagrams of the horizontal projection of P axes and triangular representation of focal mechanisms for intermediatedepth events under Central America from the Harvard catalog.
96 | Protti: Seismicity and seismotectonics of Central America with emphasis on intraslab seismic activity

seismic slab suddenly disappears. The other largest intermediate-depth event under Costa Rica was the

Naranjo, 1992, Mw=6.5 event, coincidentally at another

major segmentation of the seismic slab.

Future work

Two small projects are suggested here to contribute to

the understanding of intraslab events under Central

America. First, take the angles of the P and T axes obtained by Protti [1991] from intermediate-depth focal

mechanisms and add the projection of the slab dip to

them to set them within the slab frame. This will give a

better idea of the stress regime within the subducted plate.

This should also be done for all intermediate events contained in the Harvard catalog under Central America.

Second, compare the waveforms from the 1948 and

the 1992 earthquakes in central Costa Rica, as recorded

by the same stations and instruments to learn about their

relative locations, size and tectonic implications.

Postscript

A large (Mw=7.7), intermediate-depth (60 km) earthquake occurred off El Salvador on January 13th, 2001,

at 17:33 UTC. Despite its large magnitude and the fact

that it was felt all over the Central American isthmus,

from southern Guatemala to western Panama, the damage produced by this earthquake was relatively small. As

mentioned above, the large trench-arc gap, the steep

dip angle of the slab, the long hypocentral distance to

large cities and the path of the seismic waves through

the mantle wedge, contributed to the attenuation of the

radiated energy. The El Salvador January 13th, 2001

earthquake had a normal focal mechanism but when

the nodal planes are placed within the slab frame, the

FIGURE 3: Tectonic setting of southern Central America and geometry of the Wadati–Benioff zone (from Protti et al., [1995]). Isodepth

contours are at 20 km intervals, from 40 km. Filled tringles are active volcanoes and filled circles are the location of large (M>7.0)

earthquakes that occurred in the 20th centrury. QSC marks the projection of the Quesada sharp contortion.
The Cascadia Subduction Zone and Related Subduction Systems | 97

slip vector reflects compression within the slab resulting

in reverse faulting. There were only nine aftershocks with

magnitude above 5.0 in the following eight months and

the largest one reached only Mw=6.1. All these aftershocks occurred in a depth range from 55 to 75 km,

indicating an elastic thickness of 20 km for the slab and

they clustered on less than 80 km along the trench. Considering the large magnitude of this earthquake and the

relatively small rupture area, evidenced by the aftershocks, the El Salvador earthquake had a very large stress

drop, consistent with intraplate events.

References

DeMets, C., R.G. Gordon, D.F. Argus, and S. Stein, Current plate motions, Geophys. J. Int., 101, 425–478,

1990.

Guendel, F., and M. Protti, Sismicidad y sismotectónica

de America Central, Física de la Tierra, 10, 19–51, 1998.

FIGURE 4: Age distribution of the Cocos plate (in Ma). RSB: rough/smooth boundary; QSC: Quesada sharp contortion; CRR and ECR

are the Costa Rica and Ecuador rifts, respectively.

Hey, R.N. Tectonic evolution of the Cocos–Nazca spreading center, Geol. Soc. Am. Bull., 88, 1404–1420, 1977.

Protti, M., Correlation between the age of the subducting

Cocos plate and the geometry of the Wadati–Benioff

zone under Nicaragua and Costa Rica, M.sc. dissertation, 66 pp., University of California, Santa Cruz, 1991.

Protti, M., F. Güendel, and K. McNally, Correlation between the age of the subducting Cocos plate and the

geometry of the Wadati–Benioff zone under Nicaragua and Costa Rica, in Geologic and Tectonic Development of the Caribbean Plate Boundary in Southern

Central America: Boulder, Colorado, edited by P.

Mann, P., Geological Society of America Special Paper 295, 309–326, 1995.

Silver, E.A., D.L. Reed, J.E. Tagudin, and D.J. Heil, Implications of the north and south Panama thrust belt for

the origin of the Panama orocline, Tectonics, 9, 261–

281, 1990.
98 |
The Cascadia Subduction Zone and Related Subduction Systems | 99

Large, intermediate-depth earthquakes in Southwest Japan

Phil R. Cummins, Yoshiyuki Kaneda and Takane Hori

Frontier Research Program for Subduction Dynamics, JAMSTEC

2–15 Natsushima-cho, Yokosuka, 237–0061, Japan

cummins@jamstec.go.jp, kaneday@jamstec.go.jp, horit@jamstec.go.jp

Introduction

The subduction of the Philippine Sea plate beneath Southwest Japan leads to the repeated occurrence of great

megathrust earthquakes there, which are the focus of a

concerted earthquake disaster mitigation and prediction

effort. As shown by the 1995 Kobe earthquake, the seismic risk associated with crustal earthquakes is also substantial. There is, however, less discussion of the

earthquake risk associated with intermediate-depth earthquakes within the subducting slab. This is probably due

to the fact that almost all of the recent intraslab seismicity

has been associated with earthquakes less than M=6.5.

There were, however, two large intraslab earthquakes

which occurred in Southwest Japan in the very first part

of the 20th century. Three earthquakes of moderate size

(M= 6.6–6.8) occurred in 1952, 1968 and 1983. These

earthquakes are listed in Table 1 and the available information on them is reviewed in more detail below. (Much

of the information about damage and casualties is from

Usami, [1987]).

1905 Geiyo earthquake

This earthquake occurred beneath the Seto Inland Sea

near Hiroshima and was felt as far away as 600 km. Eleven

people were killed, 177 injured, and around 500 buildings were damaged or destroyed.

Japanese sources place the hypocenter at 40–50 km

depth, but Gutenberg and Richter [1954] and sources

based on their table report a depth of 100 km. The latter

is almost certainly in error, as suggested not only by the

intensity pattern but also by observed vertical surface

displacement as large as -4 cm.

There were a number of aftershocks, but none

located with sufficient precision to define the fault plane.

As far as we know, no focal mechanism for this earthquake has been obtained, but analysis of surface deformation is suggestive of normal faulting with strike

south–southeast. Recently, seismograms from various

types of Omori seismometers have been discovered at

the Earthquake Research Institute in Tokyo but their analysis is complicated by poor knowledge of the instrument

characteristics [Miyakawa et al., 1998].

1909 Hyuganada earthquake

The 1909 Hyuganada earthquake occurred beneath

northern Kyushu and was felt as far away as 635 km.

Despite its size, the great depth must have mitigated its

effects sufficiently that there were no deaths, but at least

five people were injured and a number of buildings damaged or destroyed.

Although this earthquake appears in a number of lists

of significant earthquakes, we have as yet found no more

detailed study. Its focal mechanism, as far as we have

been able to ascertain, is completely unknown. There

appear to have been no reported aftershocks.

1952 Yoshino earthquake

This earthquake occurred near the city of Nara in central Japan. It was felt as far away as 650 km, killed nine

people, injured 136, and damaged or destroyed about

350 buildings. There were only four reported aftershocks.

Remarkably, despite the extensive destruction caused

by this earthquake, we have been unable to find any

detailed study of its source mechanism. There is a Japan

Meteorological Agency (JMA) solution, however, which

is shown in Figure 1. We think surface deformation data

should exist, however, and hope to present these data at

the workshop.

1968 Bungo Strait earthquake

This earthquake occurred beneath the Bungo Strait separating Kyushu and Shikoku. It was felt as far away as 500

km, injured 22 people, and damaged or destroyed eight

buildings.

Because this earthquake occurred after the establishment of the World Wide Standardized Seismographic

Network (WWSSN), an extensive analysis of its source

mechanism was possible [Shiono and Mikumo, 1975]. A

substantial number of aftershocks were reported, defining a fault plane of about 20 x 20 km, with a strike of

about 20 ̊E of N and a dip of 72 ̊. The sense of faulting

was normal, with rupture initiating near the base of the

fault and propagating upward. Three centimeters of subsidence was observed near the epicenter that is consistent with this source mechanism.
100 | Cummins, Kaneda and Hori: Large, intermediate-depth earthquakes in Southwest Japan

1983 Kunisaki earthquake

This earthquake occurred beneath northern Shikoku and,

probably due to its depth, caused only one reported

injury and some light damage. Only a small number of

aftershocks were reported, but because this earthquake

occurred in (or almost in) the era of digital seismology, an

extensive study of the main shock is available [Imagawa

et al., 1985]. Assuming a 10 x 5 km fault plane based on

the limited number of aftershocks, they obtained an

average slip of 3.3 m, with rupture begining at the base of

the fault and propagating upwards. Unlike the other earthquakes discussed above, the mechanism appears to be a

high-angle thrust, but the horizontal projection of the T

axis is oriented along the strike of the subducting plate.

Conclusions

Recent intermediate-depth earthquakes in Southwest

Japan have been at most moderate in size, and this

together with their depth, has led to the impression that

the seismic risk associated with such earthquakes is small.

However, very early in the century two quite large intermediate-depth earthquakes occurred, so it seems that

the recent activity may give a misleading impression. Certainly an event similar to the 1905 Geiyo earthquake

occurring today would cause extensive damage and loss

of life.

It is not clear whether all the events listed here are

really intraslab events. Shiono and Mikumo [1975] and

Imagawa et al. [1985] characterize the 1968 and 1983

events respectively as occurring above the subducting

slab. Using our recently constructed model of the Philippine Sea plate boundary, however, we find that the

1968 event does lie within the slab and think this is very

likely true for the 1983 event as well (see Figure 1). Thus,

it seems likely that all of the events considered here are

intraslab events and moreover, are probably consistent

with the dehydration embrittlement model of Kirby et

al. [1996].

References

Gutenberg, B. and C.F. Richter, Seismicity of the Earth

and Associated Phenomena, Princeton Univ. Press,

Princeton, N.J., 310 pp., 1954.

Imagawa, K., K. Hirahara, and T. Mikumo, Source mechanisms of subcrustal and upper mantle earthquakes

around the northeastern Kyushu region, southwestern Japan, and their tectonic implications, J. Phys.

Earth, 33, 257–277, 1985.

Kirby, S., E.R. Engdahl, and R. Denlinger, Intermediatedepth intraslab earthquakes and arc volcanism as

physical expressions of crustal and uppermost mantle

metamorphism in subducting slabs, in Subduction: Top

to Bottom, G.E. Bebout et al., Eds., American Geophysical Union, Washington, D.C., pp. 195–214, 1996.

Miyagawa, K., I. Nakanishi, K. Miura, and S. Tanaka, The

1905 Geiyo earthquake (M=7.3) recorded by the

Omori seismometers and the digitization of the

records, Zishin, 51, 113–121, 1998.

Shiono, K. and T. Mikumo, Tectonic implications of

subcrustal normal faulting earthquakes in the western Shikoku region, Japan, J. Phys. Earth, 23, 257–

278, 1975.

Usami, T., Descriptive catalog of damaging earthquakes

in Japan, University of Tokyo Press, Tokyo, 1987.

Utsu, T., A catalog of large earthquakes (M≥ 6) and damaging earthquakes in Japan for the years 1885–1925,

in Historical Seismograms and Earthquakes of the

World, W.H.K. Lee et al., eds., Academic Press, San

Diego, pp. 150–161, 1987.

Earthquake Origin Time Lat Lon Depth Mag Deaths Injuries Damage* Aftershocks

Geiyo (U) 1905/06/02 05:39:07.? 34.1 132.5 40-100 7.2 22 177 500 several

Hyuganada (U) 1909/11/10 06:13:??.? 32.3 131.1 150 7.9 0 5 10+ ?

Yoshino (JMA) 1952/07/17 16:09:51.5 34.45 135.78 60 6.8 9 136 350 4

Bungo (SM) 1968/08/05 16:17:06.0 33.30 132.38 40 6.6 0 23 8

Kunisaki (IHM) 1983/08/25 21:23:34.3 33.55 131.60 116 6.8 0 1 light

Hyuganada (JMA) 1987/03/18 12:36:29.0 31.97 132.06 48 6.6 0 6 light ?

TABLE 1

* Approximate total no. damaged buildings as reported by Usami [1996].
The Cascadia Subduction Zone and Related Subduction Systems | 101

FIGURE 1: A model for the Philippine Sea plate and earthquake source mechanism solutions for large intermediate-depth earthquakes

in Southwest Japan.
102 |
The Cascadia Subduction Zone and Related Subduction Systems | 103

Large (M

w

≥ 5.5) intraplate slab earthquakes (1928–2000)

of the Prince William Sound region, Alaska

Diane Doser, Monique Velasquez, Wesley Brown and Annette Veilleux

Department of Geological Sciences, University of Texas at El Paso, El Paso, Texas 79968–0555, USA

doser@geo.utep.edu, monique@geo.utep.edu, wesley@geo.utep.edu, veilleux@geo.utep.edu

Earthquake relocations and body waveform modeling

studies have been used to examine changes in Mw≥ 5.5

seismicity before and after the 1964 great Alaskan earthquake. This portion of our study focuses on the region of

the Prince William Sound (PWS) asperity as defined by

Johnson et al. [1996] (boxes, Figure 1), while ongoing

work continues in the Kodiak Island region. In the Prince

William Sound region both the Yakutat block and Pacific

plate are being subducted beneath North America, with

the top of the Yakutat block forming the place interface.

Much of the lower-plate seismicity throughout the

study area for the past 70 years appears to cluster immediately downdip of the 1964 asperity. Hypocenters of

intraplate slab events in PWS show two persistent zones

of seismicity (Figure 2). These zones are shown in cross

section in Figure 3 with crustal structure based on the

refraction studies of Brocher et al. [1994 ] and Fuis et al.

[1991]. One zone of earthquakes (Columbia Bay [CB],

Figure 1) is located immediately downdip of the 1964

mainshock hypocenter (10–25 km) at 30–40 km depth

within the Pacific plate mantle. A second zone (Tazlina

[T], Figure 1) occurs 50–70 km from the main shock

hypocenter at 50–60 km depth in a region of the Pacific

plate mantle where the dip of the Pacific plate begins to

steepen [Brocher et al., 1994]. Both zones exhibit normal faulting with northwest–southeast directed T-axes

(Figure 2). Earthquakes of up to Mw=6.7 have occurred

within the Tazlina zone prior to the 1964 mainshock and

the largest intraslab events following the 1964 mainshock

(Mw=6.5–6.6) also occurred within this region.

A third region of intraslab earthquakes is located offshore in PWS near Montague Island (MI, Figure 2). Depth

control for these earthquakes is poor, but suggests they

are occurring either within the lower portion of the Yakutat

block or within the crust of the Pacific plate in a region

where there is significant flattening in the dip of the Pacific plate [Brocher et al., 1994]. In 1928, a Mw=7.0 earthquake occurred in this region and waveform modeling

of this event indicates reverse faulting (Figure 2), perhaps due to an upward bend in the Pacific plate.

Within Cook Inlet (Figure 4), intraslab seismicity is

complicated by the presence of a double Benioff zone

[e.g., Ratchkovsky et al., 1997] and significant contortion of the down-going slab [Doser et al., 1999]. Most

large events appear to be occurring within the Pacific

plate mantle, although slab geometry is not well constrained due to lack of seismic reflection/refraction studies such as those carried out in the PWS region. Events

at 40–60 km depth show normal to normal-oblique faulting (Figure 4) and the largest (Mw=6.7–6.8) historic events

are no deeper than 60 km. Of particular note is the

October 1954 (Mw=6.8) earthquake that was associated

with a maximum intensity of VIII [Stover and Coffman,

1993] in the Anchorage region. Since 1928 only the 1964

mainshock has produced a higher observed intensity in

Anchorage.

Below 60 km depth a wider variety of mechanisms is

observed. T-axes of Cook Inlet earthquakes show northwest to west-northwest oriented extension, regardless of

focal depth.

A higher level of seismicity within the subducting

plate in the south-central and eastern Kenai Peninsula

was observed prior to the 1964 mainshock. Since 1964,

no Mw>6.3 intraslab events have occurred within Cook

Inlet.

These observations suggest that intraplate slab earthquakes represent a significant source of seismic hazard

for the Anchorage region. The historic data also suggest

that the lower level of moment release within the subducting slab since the 1964 mainshock may not reflect the

longer-term average.

References

Brocher, T.M., G.S. Fuis, M.A. Fisher, G. Plafker, M.J. Moses, J.J. Taber, and N.I. Christensen, Mapping the

megathrust beneath the northern Gulf of Alaska using wide-angle seismic data, J. Geophys. Res., 99,

11,663–11,686, 1994.

Cohen, S.C., and J.T. Freymueller, Deformation of the

Kenai Peninsula, Alaska, J. Geophys. Res., 102,

20,479–20,488, 1997.

Doser, D.I., A.M. Veilleux, and M. Velasquez, Seismicity

of the Prince William Sound region for over thirty

years following the 1964 great Alaskan earthquake,
104 | Doser et al.: Large intraplate slab earthquakes of the Prince William Sound region, Alaska

Pure Appl. Geophys., 154, 593–632, 1999.

Fuis, G.S., E.L. Ambos, W.D. Mooney, N.I. Christensen,

and E. Geist, Crustal structure of accreted terranes in

southern Alaska, Chugach mountains and Copper

River basin, from seismic refraction results, J. Geophys.

Res., 96, 4187–4227, 1991.

Johnson, J.M., K. Satake, S.R. Holdahl, and J. Sauber,

The 1964 Prince William Sound earthquake: joint

inversion of tsunami and geodetic data, J. Geophys.

Res., 101, 523–532, 1996.

FIGURE 1: Lower plate earthquakes in south-central Alaska. Dots are Mw>5.0 earthquakes following the 1964 mainshock, diamonds

are events from 1928–1963. Boxes show slip along plate interface (in meters) during the 1964 mainshock [Johnson et al., 1996].

Yellow patches are regions with post seismic uplift >0.6 m between 1964 and 1995 [Cohen and Freymueller, 1997].

Ratchkovsky, N.A., J. Pujol, and N.N. Biswas, Relocation

of earthquakes in the Cook Inlet area, south-central

Alaska, using the joint hypocenter determination

method, Bull. Seismol. Soc. Am., 87, 620–636, 1997.

Stover, C.W., and J.L. Coffman, Seismicity of the United

States, 1568–1989 (revised), U.S. Geological Survey

Prof. Pap. 1527, 418 pp., 1993.
The Cascadia Subduction Zone and Related Subduction Systems | 105

FIGURE 2: Earthquakes in the Prince William

Sound area (1928–2000). Dots are post-1964

mainshock, diamonds are pre-1964 mainshock

and the square is the 1964 mainshock. Red denotes events below the plate interface. Dashed

lines show depth to plate interface. Cross section along South Richardson Highway (SRH)

shown in Figure 3. Focal mechanisms are for

Mw≥ 6.0 events. Numbers in brackets are

depths from waveform modeling.

FIGURE 3: Cross section along SRH (Southern Richardson Highway). Structure from

Brocher et al. [1994] and Fuis et al. [1991].

Error bars denote depth uncertainties. Triangles are events from the Harvard Centroid

Moment Tensor (CMT) catalog. Pacific plate

crust is yellow. Megathrust is bold line.
106 | Doser et al.: Large intraplate slab earthquakes of the Prince William Sound region, Alaska

FIGURE 4: Intraslab events of the Cook Inlet region (1928–2000). Symbols are as in Figure 2. Focal mechanisms are shown for Mw≥ 6.0

events. Numbers in brackets are focal depths from waveform modeling. Red mechanisms are CMT solutions.
The Cascadia Subduction Zone and Related Subduction Systems | 107

The radiated seismic energy and apparent stress of

interplate and intraslab earthquakes at subduction-zone

environments: Implications for seismic hazard estimation

George L. Choy1, John L. Boatwright2 and Stephen H. Kirby2

1 United States Geological Survey, PO Box 25046, Lakewood, Colorado, 80225, USA

2 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

choy@usgs.gov, boat@usgs.gov, skirby@usgs.gov

ABSTRACT

The radiated seismic energies (ES ) of 980 shallow subduction-zone earthquakes with magnitudes ≥ 5.8 are used

to examine global patterns of energy release and apparent stress. In contrast to traditional methods which have

relied upon empirical formulas, these energies are computed through direct spectral analysis of broadband seismic waveforms. Energy gives a physically different measure

of earthquake size than moment. Moment, being derived

from the low-frequency asymptote of the displacement

spectra, is a measure of the final static displacement. Thus,

moment is related to the long-term tectonic implication

of an earthquake. In contrast, energy, being derived from

the velocity power spectra, is more a measure of seismic

potential for damage to anthropogenic structures. There

is considerable scatter in the plot of ES–M0 for worldwide

earthquakes. For any given M0, the ES can vary by as much

as an order of magnitude about the mean regression line.

The global variation between ES and M0, while large, is not

random. When subsets of ES–M0 are plotted as a function

of seismic region, tectonic setting and faulting type, the

scatter in data is often substantially reduced. There are

two profound implications for the estimation of seismic

and tsunamic hazard. First, it is now feasible to characterize the apparent stress for particular regions. Second, a

given M0 does not have a unique ES. This means that M0

alone is not sufficient to describe all aspects of an earthquake. In particular, we have found examples of interplate

thrust-faulting earthquakes and intraslab normal-faulting

earthquakes occurring in the same epicentral region with

vastly different macroseismic effects. Despite the gross

macroseismic disparities, the MW’s in these examples were

identical. However, the Me’s (energy magnitudes) successfully distinguished the earthquakes that were more

damaging.

Introduction

There are two compelling reasons for using radiated seismic energy as a complement to moment in estimating

seismic and tsunamic hazard. First, energy gives a physically different measure of earthquake size than moment.

Energy is derived from the velocity power spectra, while

moment is derived from the low-frequency asymptote of

the displacement spectra. Thus, moment, being a measure of the final static displacement, is related to the longterm tectonic implication of an earthquake. In contrast,

energy, being strongly peaked about the corner frequency

of an earthquake, is more a measure of seismic potential

for damage to anthropogenic structures. Secondly, significant regional and tectonic variations in energy release

are suppressed by empirical formulas. Choy and Boatwright

[1995] demonstrated that systematic variations in the

release of energy and in apparent stress can now be identified that were previously undetectable because of the

lack of reliable energy estimates. In this paper, we first

review some of the important findings from Choy and

Boatwright [1995] and then proceed to identify classes

of earthquakes with unusually high intensity of energy

release relative to moment release. In particular, there is

a substantial difference in the energy radiated by normalfaulting intraslab earthquakes and by thrust-faulting

interplate earthquakes. Although M0 (or its derived magnitude MW) can characterize the area affected by the

rupture, ES (or its derived magnitude Me) is more capable

of describing macroseismic effects. Also, the highest

apparent stresses of any earthquake group are associated

with strike-slip faulting earthquakes occurring in oceanic

environments. Contrary to conventional wisdom, many

of these oceanic strike-slip earthquakes have been associated with tsunamis.

Previous methods of computing

radiated seismic energy, ES

The computation of the seismic energy radiated by an

earthquake simply requires an integration of radiated

energy flux in velocity-squared seismograms. However,

most methods of computing energy have historically
108 | Choy, Boatwright and Kirby: Radiated seismic energy and apparent stress of interplate and intraslab earthquakes

relied on empirical formulas because routine spectral

analysis was impractical with analog data. Two common

methods use the Gutenberg–Richter formulas which

derive ES from a magnitude

log ES = 5.8 + 2.4 mb

log ES = 11.8 + 1.5 MS

(in units of dyne cm). More recently, it has been suggested that

ES ~ 5 x 10-5 M0.

Notice that in all these formulas that ES is never actually computed. Instead it is predicated on another value

(mb, MS or M0). No new information is really obtained

about the earthquake.

Direct measurement of radiated seismic energy

Fortunately, theoretical and technological impediments

to the direct computation of radiated energy have now

been removed. The requisite spectral bandwidth is now

recorded digitally by a number of seismograph networks

and arrays with broadband capability. In addition, corrections for source mechanism and frequency-dependent wave propagation are better understood now than

at the time empirical formulas were first developed. We

briefly describe the method of Boatwright and Choy

[1986].

For shallow earthquakes where the source functions

of direct and surface-reflected body-wave arrivals may

overlap in time, the radiated energy of a P-wave group

(consisting of P, pP and sP) is related to the energy flux by

FIGURE 1: a) Radiated energy (ES) of 980 worldwide shallow-focus earthquakes that occurred between 1988 to 1999 as a function of

moment (M0

). The radiated energies are taken from the monthly listings of the PDE (Preliminary Determination of Epicenters) published

by the NEIC (National Earthquake Information Center) of the U.S.G.S. The slope of the least-squares linear regression (solid line) yields

a global average apparent stress for earthquakes of 4.8 bars. The 95% geometric spread (width of distribution) about the regression

line is indicated by the dashed lines. b) ES

–M0 plot for the subset of 355 shallow-focus earthquakes located near subduction zones with

predominantly thrust-faulting focal mechanisms. The geometric spread is significantly smaller for this subset than for the set of worldwide earthquakes.

where the P-wave energy flux, , is the integral of the

square of the ground velocity, corrected for frequencydependent attenuation; <F P>2 is the mean-square

radiation-pattern coefficient for P waves; RP is the P-wave

geometrical spreading factor; FgP is the generalized radiation pattern coefficient for the P-wave group defined as

where Fi are the radiation-pattern coefficients for i=P,

(a) (b)
The Cascadia Subduction Zone and Related Subduction Systems | 109

pP, and sP; and are plane-wave reflection coefficients of pP and sP at the free surface, respectively, and

corrected for free-surface amplification at the receiver;

and q is ratio of S-wave energy to P-wave energy. The

correction factors explicitly take into account our knowledge that the earthquake is a double-couple, that measurements of the waveforms are affected by interference

from depth phases, and that energy is partitioned

between P and S waves. The total radiated energy when

using the P-wave group is Es

=(1+q)EP

S

 .

Energy magnitude

From the radiated energies for a set of 378 global shallow

earthquakes Choy and Boatwright [1995] defined an

energy magnitude, Me,

log ES = 4.4 + 1.5 Me

or

Me = 2/3 log ES - 2.9

where ES is in units of Newton·meters. Note that magnitude Me is derived explicitly from energy (whereas in the

Gutenberg–Richter formula energy is derived from magnitude).

Although Me and Mw are magnitudes that describe

the size of an earthquake, they are not equivalent.

FIGURE 1 (cont’d) : c) ES–M0 for the subset of 76 shallow-focus earthquakes located near subduction zones with predominantly normalfaulting focal mechanisms. The geometric spread is smaller for this subset than for the set of worldwide earthquakes. d) ES

–M0 for the

subset of 27 shallow-focus earthquakes located near subduction zones with predominantly strike-slip focal mechanisms. The average

apparent stress of strike-slip earthquakes is a magnitude larger than that for the earthquakes with dip-slip mechanisms.

Because they measure different physical properties of

an earthquake, there is no a priori reason that they should

numerically equal for any given seismic event. Indeed,

in the following sections we show that earthquakes with

the same Mw can cause different macroseismic effects.

While the macroseismic effects cannot be distinguished

by Mw, they can be quantified by Me. The energy magnitude, Me

, is an essential complement to moment magnitude, Mw, for assessing seismic potential.

The energy and moment of an earthquake are related

by apparent stress, τa= μ Es / M0, where μ is the average

rigidity at the source. Apparent stress also serves as a good

indicator of the intensity of seismic energy radiation, Es,

relative to the size of the event as measured by the

seismic moment, M0. A plot of Es–M0 for the worldwide

population of earthquakes exhibits large scatter (Figure

1a). However, when data are separated into subsets based

on faulting mechanism and seismic region, the scatter in

the Es–M0 plots can be relatively small. The important

point is that no single empirical formula can hope to

predict Es from M0. Indeed, these differences can be exploited to detect patterns of energy and moment that

were previously masked by empirical formulas.

(c) (d)
110 | Choy, Boatwright and Kirby: Radiated seismic energy and apparent stress of interplate and intraslab earthquakes

FIGURE 2: a) ES–M0 for 46 shallow-focus thrust-faulting earthquakes that occurred between 1988 to 1999 along the Aleutian arc. The

characteristic apparent stress for this seismic region is centered narrowly about 2.3 bars. b) ES–M0 for 12 shallow-focus thrust-faulting

earthquakes that occurred between 1988 to 1994 along the Chile trench. The characteristic apparent stress for this seismic region is

centered narrowly about 3.3 bars.

FIGURE 3: Global patterns of energy release for all large (ES>1010 N•m) shallow thrust-fault earthquakes that occurred between 1988–

1991. The characteristic apparent stress, τc

, is indicated for those seismic regions where the geometric spread about the average

stress in the ES

–M0 plot is less than 15 bars.

(a) (b)
The Cascadia Subduction Zone and Related Subduction Systems | 111

Systematic variations in energy release

and apparent stress

As a function of seismic region and faulting type

The first step in sorting out the scatter in the ES–M0 plot is

to separate the populations by faulting mechanism.

Figure 1a shows the ES–M0 plot for the global set of shallow earthquakes with magnitude ≥ 5.8 that occurred

between 1988–1999. Figures 1b–1d show ES–M0 for subsets consisting of earthquakes whose focal mechanisms

are predominantly thrust, normal and strike-slip faulting

mechanisms, respectively. Although the scatter is still large,

the geometric spread about the average value of each

subset does seem to have been reduced. On a global

scale, the average apparent stress of strike-slip faulting

earthquakes is much greater than that for normal-faulting earthquakes, which in turn is slightly larger than that

for thrust-faulting earthquakes.

If we separate the data into subsets by region and

mechanism, or by tectonic setting and mechanism, we

find the reduction in the geometric spread is more

dramatic [Choy and Boatwright, 1995]. For most regions

of the world, the geometric spread is so narrow for a

given faulting type that the average apparent stress can

be regarded as the characteristic apparent stress (τc

) of

the region. Two examples are shown here (Figure 2), for

thrust-fault earthquakes at the Aleutian arc and the Chile

trench. A preliminary global map of the characteristic

apparent stress of thrust-faulting earthquakes for seismic

regions is shown in Figure 3. A τc was computed for a

seismic region only if τa’s from at least five earthquakes

were available. A similar mapping of characteristic

apparent stress for normal-faulting and strike-slip faulting earthquakes is not yet possible because of the relative paucity of such events. As more statistics on the

release of energy are accumulated, spatial and temporal

variations in energy release and apparent stress can be

refined.

As a function of tectonic setting and faulting type

The geometric spread in the ES–M0 plot is also small for

subsets based on earthquakes in specific tectonic settings.

Figure 4 shows the data for the set of continental intraplate

thrust-faulting earthquakes. Note that the average apparent stress for continental intraplate events is greater than

the average for subduction-zone thrust-faulting earthquakes.

Normal-faulting intraslab earthquakes

We found two normal-faulting intraslab earthquakes that

occurred in the immediate vicinity of interplate thrustfaulting events. For these two earthquakes, we could make

a direct comparison of the relative amount of moment

and energy associated with the different faulting types

and tectonic settings.

The source parameters and macroseismic reports of

two events that occurred in Chile are summarized in

Table 1. The epicenter of the interplate thrust-faulting

event of July 1997 is less than one degree from that of

the intraslab normal-faulting event of October 1997.

Despite the nearly identical MS’s and MW’s, as well as the

spatial and temporal proximity of the epicenters, the

intraslab normal-fault event caused extensive damage

and loss of life, whereas the effects of the interplate thrustfault event were minor. MW is unable to quantify the conspicuous difference in damage reports. On the other

hand, the Me’s are commensurate with the disparity in

macroseismic effects. The Me 7.6 of the intraslab earthquake, which caused much greater damage, is almost

one and one-half magnitude units larger than the Me 6.1

of the interplate event.

Another example of the disparity in damage between

interplate and intraplate earthquakes can be seen in a

suite of earthquakes that occurred in Kodiak, Alaska

(Table 2). Again, although the epicenters and MW’s are

similar (particularly between the first and third events),

the macroseismic effects associated with the normal-faulting intraslab earthquakes are far greater than those of

the interplate thrust-faulting earthquake. The differences

in macroseismic effects can not be distinguished if one

were to rely solely on differences in MW, but they are

FIGURE 4: ES–M0 for 17 continental intraplate earthquakes that

occurred between 1988 to 1999. The characteristic apparent

stress for earthquakes in this tectonic setting is centered narrowly about 4.7 bars, which is slightly higher than average for all

subduction-zone thrust earthquakes.
112 | Choy, Boatwright and Kirby: Radiated seismic energy and apparent stress of interplate and intraslab earthquakes

well quantified by the Me’s.

The relative paucity of normal-faulting events precludes

a similar comparison for most seismic regions. Nevertheless, radiated energy, apparent stress and Me are clearly

valuable considerations in evaluating seismic hazard.

Tsunamic potential

Choy and Boatwright [1995], using a set of 397 shallow

earthquakes distributed worldwide that occurred from

1987 to 1991, ranked earthquakes by energy and by

moment. When earthquake size was ranked by energy,

the list of the 20 largest events was dominated by strikeslip earthquakes (Table 3). When earthquake size was

ranked by seismic moment, the list of the largest events

was dominated by thrust-fault earthquakes. Two large

strike-slip earthquakes off the western coast of the U.S.,

for example, are ranked sixth and ninth by energy, but

are ranked 36th and 55th by moment. Thus, if a criterion

for seismic hazard is earthquake size, then it is important to note that the potential for damage predicted by

energy and moment would be different. A second feature in the comparison of rankings is that most of the

large (high Me) strike-slip earthquakes are associated with

tsunamis. Although strike-slip earthquakes are not generally thought of as such, they may be capable of generating tsunamis. Thus, Me may be useful in complementing MW and MS in evaluating tsunamic potential.

Estimates of radiated energy

for historical earthquakes

For earthquakes that predate the advent of broadband

digital instrumentation, ES can still be estimated if

seismograms recorded by broadband instruments can be

found and if such records are legible. However, unearthing such records and performing the digitization is both

a challenging and daunting task. Fortunately, there may

be an alternative. The preliminary results of Choy and

Boatwright [1995] show that many seismic regions have

a characteristic apparent stress, τc

=μ ES / M0. Solving for

ES and inserting the expression into the definition of Me ,

Me = 2/3 [log M0 + log ( τc / μ )] -2.9 .

This equation implies that Me, as well as ES, can be estimated for an historical earthquake in a given tectonic

setting and a specific faulting type if the M0 is known. In

contrast to ES, M0 for historical earthquakes can be

derived in numerous ways from geodetic and geologic

methods, as well as the spatial distribution of aftershocks.

1. Felt (III) at Coquimbo, La Serena, Ovalle and Vicuna.

2. Five people killed at Pueblo Nuevo, one person killed at Coquimbo, one person killed at La Chimba and another died of a heart attack at Punitaqui.

More than 300 people injured, 5000 houses destroyed, 5700 houses severely damaged, another 10,000 houses slightly damaged, numerous

power and telephone outages, landslides and rockslides in the epicentral region. Some damage (VII) at La Serena and (VI) at Ovalle. Felt (VI) at

Alto del Carmen and Illapel; (V) at Copiapo, Huasco, San Antonio, Santiago and Vallenar; (IV) at Caldera, Chanaral, Rancagua and Tierra

Amarilla; (III) at Talca; (II) at Concepcion and Taltal. Felt as far south as Valdivia. Felt (V) in Mendoza and San Juan Provinces, Argentina. Felt in

Buenos Aires, Catamarca, Cordoba, Distrito Federal and La Rioja Provinces, Argentina. Also felt in parts of Bolivia and Peru.

TABLE 1: Near the coast of central Chile.

Date Latitude ( ̊) Longitude (E) Depth (km) Me Mw mb Ms τa (bars) Faulting type

July 6, 19971 -30.06 -71.87 23.0 6.1 6.9 5.8 6.5 1 interplate-thrust

Oct. 15, 19972 -30.93 -71.22 58.0 7.5 7.1 6.8 6.8 44 intraslab-normal

TABLE 2: Kodiak Island region of Alaska.

Date Latitude ( ̊) Longitude (E) Depth (km) Me Mw mb Ms τa (bars) Faulting type

May 7, 19991 56.42 -152.94 20.0 5.8 6.3 5.7 6.1 2 interplate-thrust

Dec. 6, 19992 57.41 -154.49 52.0 7.0 7.0 6.8 - 17 intraslab-normal

Dec. 7, 19993 57.36 -154.10 45.0 6.4 6.5 6.5 6.1 15 intraslab-normal

July 11, 20004 57.60 -154.51 49.0 6.8 6.7 6.3 6.3 25 intraslab-normal

1. No felt reports.

2. Slight damage (VI) at Larsen Bay and Old Harbor. Felt strongly at Akhiok and Kodiak. Felt (III) at Homer and Chignik; (II) at Anchorage, Palmer

and Willow. Also felt at Dillingham, Eagle River, Fairbanks, Kasilof, Nelson Lagoon and Nikiski.

3. Felt at Anchorage, Ekwok, Homer, King Salmon, Kodiak and Old Harbor.

4. Some minor damage on Kodiak. Felt throughout southern Alaska and as far north as Fairbanks.
The Cascadia Subduction Zone and Related Subduction Systems | 113

Rank Region Date FM Me Mw Ms

1 Gulf of Alaska (T=85 cm) 1987 11 30 SS 8.42 7.91 7.6

2 Macquarie Islands (T) 1989 05 23 SS 8.12 8.10 8.2

3 Gulf of Alaska (T=10 cm) 1988 03 06 SS 8.12 7.79 7.6

4 Gulf of Alaska (T=10 cm) 1987 11 17 SS 7.74 7.21 6.9

5 Luzon, Philippines 1990 07 16 SS 7.65 7.74 7.8

6 Western Iran 1990 06 20 SS 7.49 7.43 7.7

7 Coast northern California (T=50 cm) 1991 08 17 SS 7.39 7.10 7.1

8 Aroe Islands (T) 1988 07 25 NR 7.35 6.96 6.7

9 Tonga Islands (T=25 cm) 1991 04 06 SS 7.22 6.69 6.7

10 Mariana Islands (T=24 cm) 1990 04 05 NR 7.23 7.47 7.5

11 Tonga Islands 1987 10 06 NR 7.26 7.30 7.3

12 Off coast of Oregon 1991 07 13 SS 7.24 6.88 6.9

13 Costa Rica (T=2 m) 1991 04 22 TH 7.14 7.68 7.6

14 Burma–China border 1988 11 06 SS 7.13 7.05 7.3

15 Banda Sea 1987 06 17 TH 7.03 7.12 --

16 East Papua New Guinea 1987 02 08 SS 7.04 7.36 7.4

17 South of Fiji 1989 08 14 SS 7.00 6.31 5.9

18 Macquarie Islands 1990 09 17 SS 7.00 6.44 6.0

19 Molucca Passage 1989 02 10 TH 6.99 7.15 6.8

20 Coast of north Chile (T=22 cm) 1987 03 05 TH 6.97 7.60 7.3

TABLE 3: Twenty largest earthquakes between 1986–1991 ranked by energy in terms of Me. The corresponding Ms and Mw are also

given. Three significant digits are used in Me and Mw in order to refine the ranking. T indicates a tsunami was generated; a wave height,

if reported, is also given. The type of faulting mechanism (FM) is indicated by TH (thrust), NR (normal) and SS (strike-slip).

Conclusions

Radiated seismic energy is a fundamental measure of

earthquake size that is different from seismic moment. It

can be used to complement seismic moment in evaluating seismic and tsunamic hazard.

• ES can be directly computed from seismic data. It need

not be an empirical quantity.

• Energy is a physically different measure of earthquake

size than moment. ES is a better estimate of the

potential for damage than M0.

• Me can be used to distinguish between earthquakes

with identical MW but different macroseismic effects.

• The correlation between ES and M0 varies systematically with faulting type, seismic region and tectonic

environment.

• In particular, the energy radiated by intraslab normalfaulting earthquakes is generally larger than that of

interplate thrust-faulting earthquakes. For some normal-faulting intraslab earthquakes, Me is more effective than MW in quantifying macroseismic effects.

• Oceanic strike-slip earthquakes can be followed by

tsunamis. Memay be a useful complement to MW and

MS in evaluating tsunamic potential.

• As statistics about radiated energy and moment are

accumulated, and their global variations become more

precise, the seismic hazard potential can be estimated

from historical as well as contemporary data.

References

Boatwright, J.L., and G.L. Choy, Teleseismic estimates of

the energy radiated by shallow earthquakes, J. Geophys.

Res., 91, 2095–2112, 1986.

Choy, G.L., and J.L. Boatwright, Global patterns of radiated seismic energy and apparent stress, J. Geophys.

Res., 100, 18,205–18,228, 1995.
114 | Choy, Boatwright and Kirby: Radiated seismic energy and apparent stress of interplate and intraslab earthquakes

TABLE 4: Twenty largest earthquakes between 1986–1991 ranked by moment in terms of Mw. The corresponding Me

 and Ms are also

given. T indicates a tsunami was generated; a wave height, if reported, is also given. The type of faulting mechanism (FM) is indicated

by TH (thrust), NR (normal) and SS (strike-slip).

Rank Region Date FM Me Mw Ms

1 Macquarie Islands (T) 1989 05 23 SS 8.12 8.10 8.2

2 Gulf of Alaska (T=85 cm) 1987 11 30 SS 8.42 7.91 7.6

3 Gulf of Alaska (T=10 cm) 1988 03 06 SS 8.12 7.79 7.6

4 Luzon, Philippines 1990 07 16 SS 7.65 7.74 7.8

5 Sulawesi 1990 04 18 TH 6.90 7.68 7.4

6 Costa Rica (T=2 m) 1991 04 22 TH 7.14 7.68 7.6

7 South of Fiji 1990 03 13 SS 6.95 7.65 7.4

8 Kuril Islands 1991 12 22 TH 6.61 7.63 7.4

9 Coast of north Chile (T=22 cm) 1987 03 05 TH 6.97 7.60 7.3

10 Solomon Islands 1988 08 10 TH 6.50 7.60 7.4

11 Mindanao, Philippines 1989 12 15 TH 6.64 7.59 7.3

12 Sulawesi 1991 06 20 TH 6.45 7.57 7.0

13 Mariana Islands (T=24 cm) 1990 04 05 NR 7.23 7.47 7.5

14 Macquarie Islands 1987 09 03 NR 6.76 7.43 7.3

15 East coast Honshu (T=56 cm) 1989 11 01 TH 6.85 7.43 7.4

16 Western Iran 1990 06 20 SS 7.49 7.43 7.7

17 New Britain (T=13 cm) 1987 10 16 TH 6.37 7.41 7.4

18 Taiwan 1986 11 14 TH 6.94 7.41 7.8

19 East Papua New Guinea 1987 02 08 SS 7.04 7.36 7.4

20 Costa Rica 1990 03 25 TH 6.36 7.36 7.0
The Cascadia Subduction Zone and Related Subduction Systems | 115

Earthquakes having high apparent stress

in oceanic intraplate lithosphere

Art McGarr1 and George L. Choy2

1 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

 2 United States Geological Survey, PO Box 25046, Lakewood, Colorado, 80225, USA

mcgarr@usgs.gov, choy@usgs.gov

Choy and Boatwright [1995] studied shallow earthquakes

distributed worldwide and concluded that the highest

apparent stresses (μEs/M0, where Es is the radiated energy,

μ is the modulus of rigidity, and M0 is the seismic

moment) are associated with strike-slip earthquakes that

occur at oceanic ridge–ridge transform faults and in

intraplate environments seaward of island arcs. Our most

recent survey of source parameters for earthquakes

during the period 1987 through 1998 revealed that the

highest apparent stresses found anywhere are associated

with events located in the depth range 10–25 km in

intraplate oceanic lithosphere. Apparent stresses occasionally exceed 20 MPa for these earthquakes, which

share two or more of the following characteristics: 1)

their mechanisms are strike-slip; 2) their P axes are perpendicular to the nearest transform fault; 3) their T axes

are perpendicular to the nearest spreading ridge or subduction zone; and 4) they are often located near triple

junctions, especially where intraplate deformation is

thought to be taking place.

Two of the highest apparent stresses, 27 and 25 MPa,

are associated with two large earthquakes in 1987 and

1988 of moment magnitude 7.2 and 7.9 at depths of 18

and 20 km in the Gulf of Alaska. These shocks involved

strike-slip faulting along a north–south oriented zone [Lahr

et al., 1988]; the sequence also includes a third event of

Mw=7.8 having the same mechanism and an apparent

stress of 13.5 MPa. This sequence of intraplate earthquakes is considered to indicate fragmentation of the

northeast corner of the Pacific plate according to Lahr et

al. [1988].

Similarly, earthquakes having exceptionally high

apparent stresses near the southeast corner of the Gorda

plate, where the Mendocino escarpment, the San

Andreas fault, and the Cascadia subduction zone form a

triple junction, are thought to reflect intraplate deformation along northeast and northwest oriented strike-slip

faults (e.g., Wilson, [1989]; Wang et al., [1997]). One of

the magnitude 6.6 aftershocks of the April, 1992, M=7.1

Petrolia earthquake was located about 25 km seaward

of the mainshock at a depth of 22 km in the oceanic

lithosphere and had an apparent stress of 16 MPa. This

event, in contrast to the mainshock and its thrust faulting mechanism, involved right-lateral slip on a northwestoriented fault plane and a high-level of ground motion

onshore [Oppenheimer et al., 1993]. Generally, the southeastern portion of the Gorda plate produces intraplate

earthquakes of very high apparent stress quite frequently.

The largest oceanic intraplate earthquake ever

recorded, which occurred near the Balleny Islands

between Australia and Antarctica, had a moment magnitude of 8.2 [Wiens et al., 1998] and a high apparent

stress of 13 MPa. This event, which was located within

the Antarctic plate at a depth of about 18 km, indicates

strike-slip deformation within the Antarctic plate. Indeed,

Conder and Forsyth [2000] have ruled out the possibility

that this earthquake occurred along the boundary of a

microplate that might have existed to accommodate part

of the motion between the Antarctic and Australian

plates.

The depth distribution of apparent stresses for earthquakes located in oceanic lithosphere shows a steep increase in maximum values 6–10 km below the ocean

bottom, peak values in the neighborhood of 25 MPa at

10–20 km depth, followed by a steep decrease at 20–24

km depth. Apparent stresses in excess of 10 MPa are all

found in the depth range 10–22 km. Interestingly, this

depth distribution of peak apparent stress is qualitatively

consistent with estimates of oceanic lithospheric strength

inferred from laboratory measurements of rock strength.

That is, Brace and Kohlstedt [1980], Kirby and Kronenberg

[1987] and Kohlstedt et al. [1995] have all proposed that

frictional failure on faults (Byerlee’s Law) limits the strength

of the uppermost oceanic lithosphere to depths of the

order of 20–30 km, below which temperature-dependent creep of olivine results in a steep decline in strength.

Thus, the laboratory evidence on lithospheric strength indicates at least the possibility of a peak in the depth range

10–20 km.

 To make the comparison between apparent stresses

and lithospheric strength more quantitative, we first relate the peak values in apparent stress τa to the total shear
116 | McGarr and Choy: Earthquakes having high apparent stress in oceanic intraplate lithosphere

stress τ causing fault slip using [McGarr, 1999)] τa/τ ≤ 0.06

and then we make the common assumption [e.g., Grasso

and Sornette, 1998] that the applied shear stress is close

to the strength limit. Laboratory-based strength estimates,

multiplied by 0.06, are in reasonably good agreement

quantitatively with our distribution of peak apparent

stresses. The curve of Kohlstedt et al. [1995] for oceanic

lithospheric strength at 20 km depth, for instance, is 350

MPa, which, if multiplied by 0.06, yields a peak apparent stress of 21 MPa. Thus, we conclude that maximum

apparent stresses for earthquakes in intraplate oceanic

lithosphere are consistent with laboratory-based estimates

of strength in the same depth range. If so, then these are

the highest apparent stresses to be found anywhere because the uppermost mantle of the oceanic lithosphere

is the strongest material within the earth.

We have not attempted to determine the sources of

such high levels of shear stress, roughly 350 MPa at depths

near 20 km, except to note that typical estimates of

“ridge-push” forces, if supported mostly by the strongest

part of the oceanic lithosphere, provide deviatoric stresses

of the same order as the strength estimates just reviewed.

More detail on the causative forces is beyond the scope

of our study.

Perhaps the most important conclusion of our study

with regard to seismic hazard is that earthquakes in

intraplate oceanic lithosphere can produce exceptionally high levels of strong ground motion for their size

inasmuch as there is quite a close correspondence

between apparent stress and high-frequency ground

motion. The large intraplate aftershocks of the 1992

Petrolia earthquake near the southeast corner of the

Gorda plate [Oppenheimer et al., 1993] provided a good

example of this effect.

References

Brace, W.F., and D.L. Kohlstedt, Limits on lithospheric

stress imposed by laboratory experiments, J. Geophys.

Res., 85, 6248–6252, 1980.

Choy, G.L., and J.L. Boatwright, Global patterns of radiated seismic energy and apparent stress, J. Geophys.

Res., 100, 18,205–18,228, 1995.

Condor, J.A., and D.W. Forsyth, Do the 1998 Antarctic

plate earthquake and its aftershocks delineate a plate

boundary?, Geophys. Res. Lett., 27, 2309–2312, 2000.

Grasso, J.-R. and D. Sornette, Testing self-organized

criticality by induced seismicity, J. Geophys. Res., 103,

29,965–29,987, 1998.

Kirby, S.H., and A.K. Kronenberg, Rheology of the lithosphere: selected topics, Rev. Geophys., 25, 1219–

1244, 1987.

Kohlstedt, D.L., B. Evans, and S.J. Mackwell, Strength of

the lithosphere: constraints imposed by laboratory

experiments, J. Geophys. Res., 100, 17,587–17,602,

1995.

Lahr, J.C., R.A. Page, C.D. Stevens, and D.H. Christensen,

Unusual earthquakes in the Gulf of Alaska and fragmentation of the Pacific plate, Geophys. Res. Lett.,

15, 1483–1486, 1988.

McGarr, A., On relating apparent stress to the stress causing earthquake fault slip, J. Geophys. Res., 104, 3003–

3011, 1999.

Oppenheimer, D., and 19 coauthors, The Cape Mendocino, California, earthquakes of April 1992: subduction at a triple junction, Science, 261, 433–438, 1993.

Wang, K., J. He, and E.E. Davis, Transform push, oblique

subduction resistance, and intraplate stress of the Juan

de Fuca plate, J. Geophys. Res., 102, 661–674, 1997.

Wiens, D.A., M.E. Wysession, and L. Lawver, Recent oceanic intraplate earthquake in Balleny Sea was largest

ever detected, EOS, Trans. AGU, 79, 353–354, 1998.

Wilson, D.S., Deformation of the so-called Gorda plate,

J. Geophys. Res., 94, 3065–3075, 1989.
The Cascadia Subduction Zone and Related Subduction Systems | 117

Energy-to-moment ratios for damaging intraslab

earthquakes: Preliminary results on a few case studies

Emile A. Okal1 and Stephen H. Kirby2

1 Department of Geological Sciences, Northwestern University, 1847 Sheridan Road, Evanston, Illinois, 60208, USA

2 United States Geological Survey, 345 Middlefield Road, Menlo Park, California, 94025, USA

emile@earth.northwestern.edu, skirby@usgs.gov

ABSTRACT

We use the energy-to-moment ratio, as introduced by

Newman and Okal [1998] to examine the source characteristics of normal-faulting intraslab earthquakes, compared to nearby interplate thrust events, based on recent

case studies in central Chile and southeastern Mexico. In

Chile, we find that the 1997 intraslab event had an exceptionally large E/M0 ratio, 30 times greater than the

nearby interplate shock. This suggests a very fast strain

release at the source as the origin of the particularly destructive character of intraslab events. While the difference is less sharp in Mexico, we find a similar trend, which

is in agreement with the observation that the locally most

damaging earthquakes are indeed the intraslab events. We

also document on the 1939 Chilean earthquake the feasibility of extending this approach to historical earthquakes

for which high-quality analog records have been archived.

Introduction

Records of instrumental seismicity over the last century

show that maximum earthquake size, measured in terms

of seismic moment M0, is achieved by interplate thrust

earthquakes at subduction zones, as exemplified by the

two largest seismic moments ever measured, the 1960

Chilean and 1964 Alaskan earthquakes (two to five times

1030 dyn-cm and 8.2x1029 dyn-cm, respectively)

[Kanamori, 1970; Kanamori and Cipar, 1974; Cifuentes

and Silver, 1989]. In this context, and for the purpose of

evaluating seismic risk in a given subduction zone, the

largest interplate thrust event expected in the region has

generally been used as a benchmark to define greatest

possible hazard. For example, in the Pacific Northwest

region of the United States, this reference event has been

taken as a so-called “mega-thrust” interplate shock [Heaton

and Kanamori, 1984; Rogers, 1988], rupturing the entire

plate boundary from Cape Mendocino to the Juan de Fuca

Strait, as proposed by Satake et al. [1996] for the source

of the transpacific tsunami of January 26, 1700.

However, in a recent contribution, Kirby [1999] has

pointed out that a number of intraplate earthquakes,

occurring within the down-going slab, can result in significantly more damage than expected from their reported magnitudes. Examples include the Olympia,

Washington event of April 13, 1949 (MPAS=7), which resulted in eight deaths and more than 150 million 1949-

dollars of damage and the Chilean earthquake of January

25, 1939 (MPAS=8.3), which killed upwards of 25,000

people, i.e., at least five times as many as the 1960 event,

in spite of an estimated moment 200 times smaller. A

similar case is that of the Peruvian earthquake of May

31, 1970 which, in addition to casualties directly due to

strong motion damage, also triggered a landslide burying the Yangay Valley and resulted in a total of 66,000

deaths. It was determined to be an intraplate event occurring within the descending slab, in a region featuring

no great interplate thrust earthquakes [Plafker et al., 1971;

Abe, 1972]. Such intraslab earthquakes are usually deeper

than 40 km, and displaced laterally away from the trench,

often beneath populated regions. They frequently, but

not always, feature normal faulting.

In comparing the hazards associated with a gigantic

interplate thrust earthquake and a smaller intraslab event

in the same region, one must bear in mind that the latter

may have greater recurrence rate by virtue of its smaller

size (expressed either as magnitude or moment). The

combination of being closer geographically to populated

areas, having a potential for greater damage and a greater

probability of occurrence may outweigh the intrinsically

smaller nature of the intraplate event and make it, rather

than the interplate “megathrust”, a potentially greater

contributor to seismic risk in the region. This scenario

would be particularly relevant in the Pacific Northwest,

where the period of recurrence of a 1700-type megathrust event may range from 600 to 1600 years, based

on the analysis of sand deposits in the intertidal zone

[Atwater, 1992], while at least three destructive intraslab

earthquakes took place in the past 52 years (Satsop,

Washington, 1999; Puget Sound, 1965 [Langston and

Blum, 1977]; and Olympia, 1949 mentioned above).

There can be a priori two explanations for the enhanced damage (in physical terms, higher accelerations)
118 | Okal and Kirby: Energy-to-moment ratios for damaging intraslab earthquakes

observed during intraslab earthquakes. One is intrinsic—

the strain release at the source would be faster than usual,

resulting in a source spectrum richer in higher frequencies. Another model would involve a receiver effect,

namely that the areas most affected by the rupture, which

are located directly above the source and hence displaced

inland with respect to an interplate event sited at the

trench, could be set in a geological environment particularly conducive to the local amplification of ground

motion. This could be the case, for example, in the sediment-filled Central Valley of Chile, under which the 1939

earthquake took place.

By studying the teleseismic characteristics of damaging intraslab earthquakes, one can hope to eliminate local site effects and discriminate between the two models.

We report here on a preliminary analysis of a few case

studies in two subduction provinces: the Oaxaca region

of southern Mexico and central Chile. We reject site effects as the source of enhanced damage and conclude

that the three intraslab earthquakes studied did indeed

exhibit a faster than usual seismic source.

Data analysis

In each of the two regions studied, we selected a pair of

earthquakes, one underthrusting interplate shock and one

intraslab event, of comparable sizes and locations. In

particular, the latter constraint ensures that site effects at

teleseismic receivers will be comparable for both events

in each pair, since seismic rays under the receiver will

travel along essentially identical paths.

In Chile, we consider the interplate earthquake of

July 6, 1997 (30.06 ̊S, 71.87 ̊W; h=19 km; M0=1.9 x

1026 dyn-cm) and the intraslab event of October 15, 1997

(30.93 ̊S, 71.22 ̊W; h=58 km; M0=4.9 x 1026 dyn-cm).

In Mexico, we study the large 1978 Oaxaca interplate

event (16.01 ̊N, 96.60 ̊W; h=18 km; M0=5.3 x 1027

dyn-cm) and the more recent normal faulting earthquake

of September 30, 1999 (16.06 ̊N, 96.93 ̊W; h=60 km;

M0=1.7 x 1027 dyn-cm; [Singh et al., 2000]). Our analysis of these events parallels that by G.L. Choy [pers.

comm.] using a generally similar methodology. The source

locations listed above are hypocentral parameters given

by the USGS. Centroid depths inverted at Harvard are

16 km (interplate) and 47 km (intraslab) in Mexico; in

Chile, the intraslab event is given at 70 km, but the depth

of the interplate one is unresolved and was constrained

at 15 km in the inversion. The distances separating the

epicenters in each pair are only 114 km in Chile and 36

km in Mexico (Figures 1 and 2).

We then use the energy-to-moment ratio to characterize the events through their “slowness” parameter

Θ = log10 E/M0,

as introduced by Newman and Okal [1998] following

the work of Boatwright and Choy [1986]. General scaling laws predict Θ=-4.90, in excellent agreement with

large worldwide datasets, such as Newman and Okal’s

[1998] and Choy and Boatwright’s [1995]. Anomalously

slow rupture, such as observed during so-called “tsunami

earthquakes”, leads to values of Θ deficient by as much

as 1.5 to 2 units. Our results are summarized on Figure

3. We use energy estimates corrected for available focal

mechanisms. For the recent (1997 and 1999) earthquakes, we analyzed between five and six broadband

IRIS records at stations providing good azimuthal coverage. For the 1978 Mexican earthquake, we were able to

find four short-period vertical records at SRO stations.

In Chile, we find Θ=-5.50 for the interplate event,

which characterizes it as mildly slow. The intraplate earthquake, on the other hand, has Θ=-3.76, making its

source one of the fastest analyzed by our technique. We

also tested the possibility of using historical earthquakes,

by processing the Pasadena–Benioff 1–100 record of the

1939 Chilean event (see Figure 4). This instrument can

be regarded as the ultimate prototype of the modern

broadband seismometer. We obtain a radiated energy

of 8 x 1023 ergs (corrected for the focal mechanism given

by Beck et al. [1998]), leading to Θ=-4.04, for a probFIGURE 1: Map of central Chile, showing the events analyzed in

the present study. Depths plotted are for centroid solutions, either from the Harvard centroid moment tensor (CMT) catalog or

as estimated by Beck et al. [1998] for the 1939 event. The focal

mechanisms are color-keyed to the value of Θ (see Figure 3).
The Cascadia Subduction Zone and Related Subduction Systems | 119

FIGURE 2: Map of southcentral Mexico, showing the events analyzed in the present study. Depths plotted are for centroid solutions,

compiled from the CMT catalog. The focal mechanisms are color-keyed to the value of Θ (see Figure 3).

FIGURE 3: Summary of results obtained in this study. This figure plots radiated seismic energy against seismic moment, in logarithmic

units (after Newman and Okal [1998]). The dashed diagonals are lines of constant Θ, with the thick line (Θ=−4.9) showing the relation

predicted by scaling laws; they are color-keyed from fast (blue) to slow (red) sources. The smaller gray symbols are the dataset

examined in Newman and Okal [1998] with bull’s eye symbols referring to the four recent tsunami earthquakes (T: Tonga, 1982; N:

Nicaragua, 1992; J: Java, 1994; P: Peru, 1996). The larger symbols show the events studied here. Interplate thrust earthquakes are

labeled in roman type, intraslab ones in italics. Note the large disparity in Θ values in Chile and the smaller one in Mexico.
120 | Okal and Kirby: Energy-to-moment ratios for damaging intraslab earthquakes

able moment of 1 x 1028 dyn-cm, as suggested by the G2

waves recorded on the Pasadena strainmeter. This indicates that the event is definitely fast.

In Mexico, the situation is not as clear-cut, with the

deeper event only slightly faster than the interplate shock.

The 1999 event has Θ=-4.93, which is essentially the

value predicted theoretically on the basis of scaling laws

for shallow earthquakes. For the 1978 event, we find

Θ=-5.37, giving it a weak trend towards slowness. It is

possible that the average character of the 1999 event is

due to its relatively shallow centroid depth, as compared

to its Chilean counterpart. While the difference in Θ between the two Mexican events is smaller than in Chile

and the results are not as conclusive, the trend upholds

the observation that the 1999 earthquake was more destructive (with 31 people killed locally) than the 1978

event (one local fatality). Similarly, the great normal-faulting intraslab earthquake of January 15, 1931 was by far

the most destructive in Oaxaca province this century

[Ordoñez, 1931; Singh et al., 1985]. Higher energy-tomoment ratios for shallow normal faulting events (in various configurations including intraslab) were also found

by Choy and Boatwright [1995] and G.L. Choy [pers.

comm. to Steve Kirby, 2000].

Conclusion

This very preliminary study, in which we have documented resolvable differences in source slowness between

interplate and intraslab earthquakes, suggests the systematic use of the parameter Θ, as introduced by Newman

and Okal [1998], in the evaluation of the source characteristics of intraslab earthquakes worldwide. It further

demonstrates on the case of the 1939 Chilean earthquake

the feasibility of extending the technique to historical

earthquakes through the use of high-quality analog data.

Our work also confirms trends involving these types of

earthquakes reported by Choy and Boatwright [1995] and

G.L. Choy [pers. comm. 2000].

In the case of Chile, the comparative analysis of P

waves from nearby earthquake pairs (one interplate and

one intraslab) at teleseismic distances proves that the large

accelerations responsible for the destruction during the

intraslab event were an inherent characteristic of its

source, rather than a local effect due to the structure of

the Central Valley of Chile. This is in agreement with

similar results reported by Singh et al. [this volume] in

Mexico.

Postscript

Since this research was presented at the Intraslab Workshop and this report written up, three major intraslab

earthquakes took place: on January 13, 2001 in El Salvador; February 28, 2001 in Washington (the "Nisqually"

event); and March 24, 2001 near Hiroshima, Japan, all

three featuring higher values of Θ [Okal and Kirby, 2001].

References

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1966 and 1970 Peru earthquakes, Phys.Earth Planet.

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the past 2000 years along the Copali River, southern

FIGURE 4: Top: Close-up of the P-wave arrival of the 1939 Chilean intraslab earthquake, as recorded on the Pasadena–Benioff 1–100

north-south instrument (tick marks are minutes). Bottom: After optical magnification, such records can be successfully digitized at a

rate of 10 samples per second and processed through Newman and Okal’s [1998] algorithm.
The Cascadia Subduction Zone and Related Subduction Systems | 121

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122 |
The Cascadia Subduction Zone and Related Subduction Systems | 123

Thermal structure and metamorphism of subducting

oceanic crust: Insight into Cascadia intraslab earthquakes

Simon M. Peacock1, Kelin Wang2 and Aaron M. McMahon1

1 Department of Geological Sciences, Arizona State University, Box 871404, Tempe, Arizona, 85287–1404, USA

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

peacock@asu.edu, wang@pgc.nrcan.gc.ca, Aaron.McMahon@asu.edu

Most of the world’s earthquakes, including the largest

earthquakes ever recorded, occur in subduction zones

where cool oceanic lithosphere descends into the mantle. Subduction-zone earthquakes relieve complex stresses

generated by the interaction between the forces that drive

and resist subduction, slab flexure, thermal expansion and

metamorphic densification reactions [e.g., Isacks and

Barazangi, 1977; Spence, 1987; Kirby et al., 1996].

Recent efforts to integrate thermal-petrologic models of

subduction zones with seismological observations suggest

that metamorphic reactions cause, or at least control, most

subduction-zone earthquakes, particularly intraslab earthquakes that occur at depths >40 km where high confining pressure makes normal frictional sliding difficult.

Intermediate-depth, intraslab earthquakes can be very

large (e.g., 1963 Banda Sea MW=8.3) and pose a significant seismic risk. In this century, intermediate-depth earthquakes have killed tens of thousands of people (e.g., 1970

Central Peru, MW=7.5–8.0, depth=40–66 km; 2001 El

Salvador, MW =7.6, depth=39–57 km) and have caused

the most destruction in the Pacific Northwest (1949 Olympia, M=7.1, depth=54 km).

The thermal structure of a subduction zone is determined primarily by the convergence rate and the

thermal structure (related to age) of the incoming lithosphere [e.g., Peacock, 1996]. Compared to most subduction zones, the Cascadia subduction zone is relatively

warm because of the relatively slow convergence rate

(~40 mm/yr), the young age (5–10 Ma) of the incoming

Juan de Fuca plate, and the thick (3–3.5 km) blanket of

insulating sediments which acts to increase temperatures

in the underlying oceanic crust. The calculated thermal

structure of the Cascadia subduction zone is very similar

to other warm subduction zones like southwest Japan

(Nankai) and southern Mexico (Figure 1).

Although the Cascadia subduction zone is relatively

warm, the subduction of the Juan de Fuca plate beneath

North America cools the Cascadia forearc. The low surface heat flux through the Cascadia forearc reflects the

downward advection of isotherms by the subducting lithosphere [Blackwell and Steele, 1992]. Hyndman and Wang

[1995] constructed two-dimensional thermal models

across the Cascadia forearc. The model which best fits

the observed heat flow data requires negligible frictional

heating which is consistent with the apparent lack of shear

heating in the warm Nankai subduction zone. Hyndman

and Wang’s [1995] thermal models predict thrust temperatures of ~250°C at the deformation front increasing

to ~525°C at 40 km depth. The downdip limit of the

Cascadia seismogenic zone (where shallow-thrust earthquakes occur) may be controlled by the temperatureinduced transition (350–450°C) from stick-slip to stable

sliding behavior [Hyndman and Wang, 1995].

Intermediate-depth, intraslab earthquakes represent

brittle failure within the subducting slab (intraplate) rather

than along the plate interface (interplate). Focal mechanisms for intraslab earthquakes commonly show normal

faulting with downdip extension although there are many

exceptions. Several lines of evidence suggest that most

intraslab earthquakes occur within the subducting

oceanic crust: 1) seismic tomography studies show that

intermediate-depth earthquakes occur near the top of

the high seismic-velocity slab [Zhao et al., 1994]; 2) shallow intraslab earthquakes merge with the downdip limit

of the interplate earthquakes indicating that intraslab

earthquakes occur very near the plate interface within

the uppermost few kilometers of the subducting slab

[Kirby et al., 1996]; and 3) in well-studied subduction

zones, intraslab earthquakes occur within 5–10 km of

the plate interface as defined independently by converted

seismic phases (e.g., northeast Japan; Hasegawa et al.,

[1994]) or seismic reflection studies (e.g., Cascadia;

Creager et al., [this volume]). Intraslab earthquakes may

also occur within the subducting mantle. For example,

two recent intraslab earthquakes in southern Mexico

occurred 15–30 km beneath the plate interface as defined by interplate thrust events in the same region [Singh

et al., this volume]. Intermediate-depth earthquakes in

several subduction zones, most notably northeastern

Japan [Hasegawa et al., 1994], define two seismic zones

separated by 20–40 km. The lower seismic zone clearly

occurs within the subducting mantle; such intraslab earth-
124 | Peacock, Wang and McMahon: Thermal structure and metamorphism of subducting oceanic crust

FIGURE 1: Calculated P-T paths and metamorphic conditions encountered by oceanic crust subducted beneath Cascadia, southern

Mexico, southwest Japan and northeast Japan [Peacock and Wang, 1999; Currie et al., 2000; this study]. Metamorphic facies: EA=

epidote-amphibolite; EB=epidote blueschist; GS=greenschist; Px Gran=pyroxene granulite. Hydrous minerals stable in the eclogite

facies shown in italics [Poli and Schmidt, 1995]: amph=amphibole; ctoid=chlorotoid; lawsonite; zoi=clinozoisite. Partial melting

reactions in red (see references in Peacock et al., [1994]).

quakes may be triggered by serpentine dehydration reactions [Peacock, 2001].

Kirby et al. [1996] proposed a model whereby intermediate-depth earthquakes are triggered by metamorphic dehydration reactions that embrittle the subducting

oceanic crust. In this model, the transformation of hydrated oceanic crust (basalt, gabbro) to eclogite releases

substantial amounts of H2O that increases pore pressures

and promotes brittle reactivation of pre-existing faults.

Kirby et al. [1996] noted that the maximum depth of

intermediate-depth earthquakes correlates with subduction-zone temperatures; cooler subducting slabs have

deeper earthquakes. In order to test this model, Peacock

and Wang [1999] constructed subduction-zone thermal

models that extend to depths >200 km by incorporating flow of the mantle wedge induced by the subducting

slab. In a detailed study of the warm southwest Japan

and cool northeast Japan subduction zones, we demonstrated that striking differences in intermediate-depth

seismicity, as well as differences in seismic-velocity structure and arc magmatism, can be linked to contrasting

metamorphic reactions and temperatures in the subducting oceanic crust [Peacock and Wang, 1999].

Subducting sediment contains substantial amounts of

pore fluids that are expelled at shallow depths <10 km

(Figure 2) [Moore and Vrolijk, 1992]. Most of the water

liberated from subducting slabs at depths >10 km is

derived from the variably hydrated basalts and gabbros

of the subducting oceanic crust (Figure 2) [e.g., Peacock,

1990]. At temperatures of perhaps 200–400°C, substantial amounts of pore water may be expelled from the

uppermost basaltic section by porosity collapse. Alterna-
The Cascadia Subduction Zone and Related Subduction Systems | 125

tively, this pore water may react with unaltered basalt to

form low-temperature minerals, such as zeolites, which

subsequently dehydrate. The progressive metamorphism

of metabasalts involves complex, continuous reactions.

The most important reactions in subducting oceanic crust

involve the transformation to eclogite, a relatively dense,

anhydrous rock consisting primarily of garnet and omphacite (Na–Ca clinopyroxene). In a given subduction zone,

the depth and nature of eclogite formation and associated dehydration reactions depends on the pressure–temperature (P–T) conditions encountered by the subducting

oceanic crust. In relatively warm subduction zones like

Cascadia, southwest Japan and southern Mexico, the

transformation to eclogite occurs at ~50 km depth (Figure 1). In contrast, calculated P–T paths for relatively cool

subduction zones like northeast Japan pass through the

blueschist facies and eclogite may not form until depths

>100 km.

Predicting the amount and location of H2

O release

from the subducting crust is difficult because of the uncertain amount and distribution of H2

O in the Juan de

Fuca plate prior to subduction (Figure 2). Anhydrous portions of the crust will not release any H2

O whereas fully

hydrated portions may release as much as 5–10 wt. %

H2O during subduction. Calculated P–T paths for the

upper basaltic part of the subducting Juan de Fuca crust

pass through a complex sequence of metamorphic facies:

prehnite–actinolite→ greenschist→ epidote–amphibolite

→ eclogite facies (Figures 1 and 2). Dehydration of fully

hydrated portions of the basaltic crust will occur preferentially at facies boundaries, although there is likely to

be some dehydration within facies. Important hydrous

minerals which will dehydrate during subduction include

prehnite (4.4 wt. % H2

O), chlorite (13 wt. % H2O), clinozoisite (2 wt. % H2

O) and amphibole (2 wt. % H2O). Our

calculations suggest that Cascadia intraslab earthquakes

FIGURE 2: Predicted thermal structure and fluid production for Cascadia subduction zone. A) Calculated maximum H2O flux expelled

from the subducting Juan de Fuca plate as a function of distance from the deformation front. Note that calculated fluid fluxes are subject

to large uncertainties. Sediment compaction dominates the fluid flux near the deformation front whereas dehydration reactions in

subducted metabasalt dominates the fluid flux at ~50 km depth where intraslab earthquakes occur. Important fluid release processes:

a= sediment porosity collapse (compaction); b= sediment mica dehydration (800°C); c= basalt porosity collapse (200–400°C); d=

basalt garnet-forming dehydration reactions (~500°C); e= basalt high-T dehydration reactions (500–700°C). B) Calculated thermal

structure for the Cascadia subduction zone at 42°N based on a preliminary finite-element model.
126 | Peacock, Wang and McMahon: Thermal structure and metamorphism of subducting oceanic crust

occur where metabasalts in the Juan de Fuca plate undergo garnet-forming and higher temperature dehydration reactions (Figure 2). In contrast, the lower gabbroic

section of the subducting Juan de Fuca crust at the deformation front lies in the high-temperature amphibolite

facies where hornblende (amphibole) may be the only

stable hydrous mineral. Both the basaltic and gabbroic

sections are predicted to transform to eclogite at 40–50

km depth which correlates with the rapid increase in slab

dip. Significant amounts of H2

O is released during transformation to eclogite but small amounts of H2

O-bearing

amphibole and clinozoisite may persist to ~80 km depth.

If intermediate-depth seismicity is triggered by metamorphic dehydration reactions in the subducting crust,

why does the Oregon sector generally lack intermediate-depth earthquakes? Possible explanations for the lack

of seismicity include: 1) the Juan de Fuca crust entering

the Oregon part of the subduction zone is essentially

anhydrous; 2) H2

O is released by metamorphic dehydration reactions but the hydrogeology in this region (controlled by the overlying Siletz terrane?) prevents fluid

pressures from reaching levels necessary to trigger faulting; 3) slab stresses in this region are not sufficient to

trigger seismicity (no slab pull?); or 4) the slab in the

Oregon sector may be unusually warm due to interaction with the Newberry hot spot.

Acknowledgements

We wish to express our thanks to the Geological Survey

of Canada, the United States Geological Survey and the

organizers of the Cascadia Intraslab Earthquake workshop for arranging a very stimulating and productive conference in Victoria, British Columbia. This research was

supported by NSF grant EAR 98-09602 (SMP).

References

Blackwell, D.D., and J.L. Steele, Geothermal Map of North

America: Decade of North American Geology, DNAG,

Continent-Scale Map-006, Geological Society of

America, Boulder, CO, 1992.

Creager, K.C., L.A. Preston, R.S. Crosson, T. Van Wagoner, A.M. Tréhu, and the SHIPS98 working group,

Three-dimensional reflection image of the subducting

Juan de Fuca Plate, in The Cascadia Subduction Zone

and Related Subduction Systems, edited by S.H. Kirby,

K. Wang, and S.G. Dunlop, pp. 37–42, U.S. Geological Survey Open-File Report 02–328, Geological

Survey of Canada Open File 4350, 2002.

Currie, C.A., K. Wang, and R.D. Hyndman, Thermal

constraints on megathrust earthquakes of the Mexican subduction zone, EOS Trans. AGU (Fall meeting

suppl.), 81, F910, 2000.

Hasegawa, A., S. Horiuchi, and N. Umino, Seismic structure of the northeastern Japan convergent plate

margin: a synthesis, J. Geophys. Res., 99, 22,295–

22,311, 1994.

Hyndman, R.D., and K. Wang, The rupture zone of

Cascadia great earthquakes from current deformation and the thermal regime, J. Geophys. Res., 100,

22,133–22,154, 1995.

Isacks, B.L., and M. Barazangi, Geometry of Benioff

zones: lateral segmentation and downwards bending of the subducted lithosphere, in Island Arcs, Deep

Sea Trenches and Back Arc Basins (Maurice Ewing Series Vol. 1), edited by M. Talwani and W.C. Pitman,

pp. 99–114, American Geophysical Union, Washington, D.C., 1977.

Kirby, S., E.R. Engdahl, and R. Denlinger, Intermediatedepth intraslab earthquakes and arc volcanism as

physical expressions of crustal and uppermost mantle metamorphism in subducting slabs, in Subduction:

Top to Bottom, AGU Monograph 96, edited by G.E.

Bebout et al., pp. 195–214, American Geophysical

Union, Washington, D.C., 1996.

Moore, J.C., and P. Vrolijk, Fluids in accretionary prisms,

Rev. Geophysics, 30, 113–135, 1992.

Peacock, S.M., Fluid processes in subduction zones, Science, 248, 329–337, 1990.

Peacock, S.M., T. Rushmer, and A.B. Thompson, Partial

melting of subducting oceanic crust, Earth Planet. Sci.

Lett., 121, 227–244, 1994.

Peacock, S.M., Thermal and petrologic structure of subduction zones, in Subduction: Top to Bottom, AGU

Monograph 96, edited by G.E. Bebout et al., pp. 119–

133, American Geophysical Union, Washington,

D.C., 1996.

Peacock, S.M., Are double seismic zones caused by serpentine dehydration reactions in subducting oceanic

mantle? Geology, 29, 299–302, 2001.

Peacock, S.M., and K. Wang, Seismic consequences of

warm versus cool subduction metamorphism: examples from southwest and northeast Japan, Science,

286, 937–939, 1999.

Poli, S., and M.W. Schmidt, H2

O transport and release

in subduction zones: experimental constraints on

basaltic and andesitic systems, J. Geophys. Res., 100,

22,299–22,314, 1995.

Singh, S.K., V. Kostoglodov, and J.F. Pacheco, Intraslab earthquakes in the subducting oceanic plates below Mexico,

in The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 87–92, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

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22,313–22,329, 1994.
The Cascadia Subduction Zone and Related Subduction Systems | 127

Heat sources in subduction zones:

Implications for slab seismicity and arc volcanism

Jason R. McKenna and David D. Blackwell

Department of Geological Sciences, Southern Methodist University, PO Box 750395, Dallas, Texas 75275

jmckenna@post.smu.edu, blackwel@passion.isem.smu.edu

Introduction

The shallow (upper 30–50 km) thermal structure of subduction zones is well-constrained, and with few exceptions, can be modeled utilizing a variety of numerical

[e.g., Hyndman and Wang, 1993] or even analytical solutions [e.g., Molnar and England, 1995]. It is stressed

that the intermediate thermal structure of subduction

zones remains largely unknown because the mechanism

responsible for arc-magma genesis in the depth range of

80–120 km remains subject to debate. Attempts to investigate the source region of intermediate-deep slab

seismicity are likewise hampered by the poorly-constrained thermal structure at these depths. Constraints

on the conditions on the subducting slab face are therefore necessary to any critical assessment of seismogenic

conditions within the interior of the slab.

Most comprehensive thermal models have concluded

that additional heat-mass transfer from outside the mantle wedge is necessary to explain the occurrence of arc

volcanism and back-arc spreading [Minear and Toksoz,

1970; Tatsumi and Eggins, 1995]. Early thermal models

by McKenzie and Schlater [1968], Oxburgh and Turcotte

[1968; 1970] and Minear and Toksoz [1970] tested various heat sources to explain the occurrence of arc volcanism including conductive heating from a stagnant mantle

wedge, exothermic phase changes, brittle shear-heating

and viscous dissipation. All were ultimately rejected because they failed to satisfy the surface heat-flow over the

subduction zone or generate the high-eruption temperature basalts observed in subduction zones (see Davies

and Stevenson [1992] for a discussion).

However, Thatcher and England [1998] have demonstrated that significant ductile shearing localized in the

lower crust/upper mantle along the San Andreas fault

can explain the observed California Coast Ranges heatflow anomaly. Because the viscosity of a ductile medium

is temperature dependent, ductile shearing is rapidly

concentrated in a narrow zone about the slip zone [Yuen

et al., 1978] and is virtually independent of slip velocity

[Thatcher and England, 1998].

Thermal modeling results

The thermal structure of the shallowest 30–50 km of the

subduction zone can be modeled satisfactorily using the

forearc surface heat-flow profile as the primary constraint

[e.g., Honda, 1985; Hyndman and Wang, 1993]. However, recent work by McKenna and Blackwell [2002] has

demonstrated that the shallow thermal regime of even

young subduction zones is entirely conductive and can

be successfully modeled without consideration of the

deeper processes that influence the thermal structure of

the subduction zone (i.e., arc magmatism and mantle

wedge convection).

A thermal model specific to the subduction of the

Juan de Fuca plate near 40°N (northern California) has

been constructed assuming a subducting plate velocity

of 40 mm/yr (the details of which will be presented elsewhere). Qualitatively, the model is comprised of moderately sedimented ~6 Ma Gorda lithosphere subducting

beneath the North American backstop (composed of

Mesozoic Franciscan complex and Klamath granite) and

is overlain by a localized basin of Eel River sediments

[e.g., Smith et al., 1993]. This structure was extended

eastward to the volcanic arc using crustal structure from

Mooney and Weaver [1989]. This thermal model is

typical of our estimates for the thermal structure in the

upper ~40 km of the southern Cascadia subduction zone

system. Figure 1 illustrates the resultant temperature–

depth curve along the subducting plate top for the initial

model. Also presented is a temperature–depth curve

representing the Sierra Nevada paleo-subduction zone

to emphasize the difference between the cold subduction that occurred in Mesozoic North America (Sierra

Nevada) and the much warmer thermal conditions in

the present-day Cascadia subduction zone. The most

interesting feature of the Cascadia curves is that temperature along the subducting slab face increases rapidly, attaining temperatures in excess of 450°C by the

relatively shallow depth of five kilometers. As depth increases, the temperature along the subducting slab face

increases only moderately.

The effect of viscous shear-heating on the subduc-
128 | McKenna and Blackwell: Heat sources in subduction zone: implications for slab seismicity and arc volcanism

tion zone thermal structure is simulated by specifying a

heat source within a thin (two kilometers) boundary layer

blanketing the slab face for a 200 km horizontal extent

(beginning about 115 km landward of the deformation

front). The strength of the additional heat source within

the boundary layer is between 100 and 200 mW/m2.

We feel that while rather high, these values are not

unreasonable estimates for the additional heat generated by viscous dissipation because in the presence of a

non-linear (i.e., temperature-dependent) mantle viscosity, even moderate shear velocities can produce large temperature gradients across a thin boundary layer [Thatcher

and England, 1998]. The resultant temperature–depth

curves for the models that include the effects of viscous

dissipation are presented in Figure 1. The point at which

the heat source is “turned on” is visible as the bifurcation of the initial model curve.

Basalts are volumetrically the dominant lava type in

the Cascades [Blackwell et al., 1982; 1990]. Models that

satisfy the surface heat-flow profile and still produce the

requisite temperature (~1100°C) in the mantle wedge

source region are considered successful. Utilizing this criteria, it is clear that only the curve representing the model

with an additional heat input of 200 mW/m2 begins to

approach the temperature necessary for arc-basalts.

Figure 2 presents the results of the thermal calculations as pressure–temperature paths along the slab face.

It has been suggested that the densification to eclogite

via metamorphic dehydration reactions may be responsible for intermediate-depth intraslab seismicity [Kirby

et al., 1996; Peacock et al., this volume]. Again, we show

the blueschist generating P–T conditions in the Sierra

Nevada paleo-subduction zone to illustrate the distinct

thermal conditions in present-day Cascadia. Here, even

the initial model P–T path (no additional heat sources)

almost reaches the eclogite stability field. Any additional

heat introduced downdip through viscous dissipation

promotes conditions in which the slab face readily transforms to eclogite by the relatively shallow depth of ~45

km, about 100–150 km west of the volcanic arc. We conclude from this analysis that seismicity below this depth

is probably not related to the densification to eclogite,

but most likely mechanically generated, and must therefore, be restricted to the cooler interior of the slab.

In addition to facilitating potentially seismogenic P–T

conditions at shallow depths, the addition of viscous

dissipation through a boundary layer overlying the subducting slab can induce local convective upwelling via

boundary layer separation. In fluid dynamics, the Reynolds

Number describes the stability of some particular flow

FIGURE 1: Temperature–depth curves along the subducting slab face, northern California. Note the high temperatures encountered

along the slab face, even without the additional heat generated by viscous dissipation, when compared to the Sierra Nevada paleosubduction zone.
The Cascadia Subduction Zone and Related Subduction Systems | 129

field and is defined as Re=UL/ν, where U is the freestream velocity, L is the characteristic length of the flow,

and ν is the kinematic viscosity. When Re<<1, viscous

forces dominate the flow and the velocity is inversely

proportional to (dynamic) viscosity. Furthermore, any

inertial contributions to the flow may be neglected,

greatly reducing the complexity of the analysis. Most, if

not all, subduction zones can be characterized by an

extremely small Re due to the initial large mantle viscosity, typically, 1x1019 or 1x1020 Pa·s in the uppermost

mantle. For the northern California subduction zone,

U=40 mm/yr, ν~1x1019 Pa·s, and L=200 km, yielding a

Re~2x10-23.

Along the subducting slab, parameters change much

faster normal to the slab than along it, so we may expect

that a boundary layer of thickness d to be present along

the slab. If we define our coordinate system so that the

horizontal flow parallels the slab dip, the vertical component of velocity is normal to the slab and should be

laminar, so long as: d<<L and Re<<1. This is generally

true if viscosity is constant. However, if the material properties are temperature-dependent, rapid departures from

laminar flow may occur. We have determined the velocity in a boundary layer two kilometers thick parallel to

the subducting Gorda slab in northern California. We

allowed the viscosity to depend on temperature as

follows: μ=μo·exp[Q/RT], where μo=1x1019·exp[-Q/

R(T+1073.15)]. Q, the activation energy is 522 kJ/m3, R,

the universal gas constant is 8.317 J/K·mol, and T is

absolute temperature. Once the velocity field is known

within the boundary layer, the point at which the boundary layer separates from the slab is simply where the shearstress is zero on the slab: (du/dz)

slab=0.

Figure 3 illustrates the distance at which boundary

layer separation occurs for the thermal models incorporating viscous dissipation. The initial thermal model does

not attain the vertical velocity gradient reversal necessary for separation within the model geometry. In the

model geometry, the position of the volcanic arc corresponds to a distance of 100 km. It is clear that only the

models with an additional heat source produce the

upwelling beneath the position of the volcanic arc.

Conclusions

1. Reasonable heat source estimates (100–200 mW/m2

)

in a thin boundary layer blanketing the subducting

slab-face raises the ambient mantle temperature by

at least 200–400°C beneath the volcanic arc. Viscous

FIGURE 2: Pressure–temperature paths, northern California, curves are the same as in Figure 1. The potentially seismogenic phase

transition to eclogite occurs around 45 km depth, and about 100–150 km west of the volcanic arc. If the densification to eclogite is

largely completed by this depth, any observed seismicity is likely mechanically generated, and must therefore, be restricted to the

cooler interior of the slab.
130 | McKenna and Blackwell: Heat sources in subduction zone: implications for slab seismicity and arc volcanism

dissipation of heat produced in this layer can produce the minimum source temperature (>1100°C)

necessary for the generation of the volumetrically significant basalts observed in the Cascades.

2. High temperatures along the subducting slab predict

the seismogenic phase transition to eclogite at relatively shallow depths, even without viscous dissipation, resulting in little intraslab seismicity below ~45

km in the subducting Gorda slab.

3. Because the high slab face temperature restricts seismicity in the upper portion of the slab, the subduction

zone thermal structure is a boundary condition on the

occurrence of intraslab seismicity that can aid in the

interpretation of the subducting slab morphology.

4. The separation of the boundary layer from the slab is

a direct result of the strong temperature-dependence

of the viscosity and occurs approximately beneath the

volcanic arc. This separation should produce local convective upwelling. However, the interaction of the

separating boundary layer and any deeper, “cornerflow” remains uncertain. Future work will attempt to

model this interaction and quantify the mass-flux off

the slab in this region.

References

Blackwell, D.D., R.G. Bowen, D.A. Hull, J. Riccio, and

J.L. Steele, Heat flow, arc volcanism, and subduction

in northern Oregon, J. Geophys. Res., 87, 8735–8754,

1982.

Blackwell, D.D., J.L. Steele, S.H. Kelley, and M.A. Korosec,

Heat flow in the state of Washington and thermal

conditions in the Cascade Range, J. Geophys. Res.,

95, 19495–19516, 1990.

Davies, J.H., and D.J. Stevenson, Physical model of source

region of subduction zone volcanics, J. Geophys. Res.,

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Japan: a case study for understanding the detailed thermal structure of the subduction zone, Tectonophysics,

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Hyndman, R.D., and K. Wang, Thermal constraints on

the zone of major thrust earthquake failure: the

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2039–2060, 1993.

Kirby, S.H., E.R. Engdahl, and R. Denlinger, Intermediate-depth intraslab earthquakes and arc volcanism

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Scholl, S. Kirby, and J.P. Platt, pp. 195–214, American Geophysical Union, Washington D.C., 1996.

McKenna, J.R., and D.D. Blackwell, Shallow thermal

conditions in the middle and southern Cascadia subduction zone, J. Geophys. Res., in press, 2002.

McKenzie, D.P., and J.C. Schlater, Heat-flow inside the

island arcs of the northwest Pacific, J. Geophys. Res.,

73, 3173–3179, 1968.

Minear, J.W., and M.N. Toksoz, Thermal regime of a

downgoing slab and new global tectonics, J. Geophys.

Res., 75, 1397–1419, 1970.

Molnar, P. and P. England, Temperatures in zones of steady

state underthrusting of young oceanic lithosphere,

FIGURE 3: Boundary layer separation in the northern California subduction zone resulting from viscous dissipation. When the vertical

velocity gradient reverses sign, the thin boundary layer that blankets the subducting slab begins to peel off from the slab, generating

convective upwelling. Here, the separation point for 100 mW/m2 is about 105 km, while for 200 mW/m2

, it is closer to 125 km. Both

distances are almost directly below the volcanic arc, suggesting that boundary layer separation may contribute mass-flux to the

volcanic arc system.
The Cascadia Subduction Zone and Related Subduction Systems | 131

Earth Planet. Sci. Lett., 131, 57–70, 1995.

Mooney, W.D., and C.S. Weaver, Regional crustal structure and tectonics of the Pacific coast states: California, Oregon and Washington, in Geophysical

Framework of the Continental United States, Geological Society of America Mem. 172, edited by L.C.

Pakiser, and W.D. Mooney, pp. 493–522, Geological

Society of America, 1989.

Oxburgh, E.R., and D.L. Turcotte, Problem of high heat

flow and volcanism associated with zones of descending mantle convective flow, Nature, 218, 1041–1043,

1968.

Oxburgh, E.R., and D.L. Turcotte, Thermal structure of

island arcs, Geol. Soc. Amer. Bull., 81, 1665–1688,

1970.

Peacock, S.M., The importance of blueschist–eclogite dehydration reactions in subducting oceanic crust, Geol.

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Peacock, S., K. Wang, and A.M. McMahon, Thermal

structure and metamorphism of subducting oceanic

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132 |
The Cascadia Subduction Zone and Related Subduction Systems | 133

Theoretical mineralogy, density, seismic wavespeeds

and H2

O content of the Cascadia subduction zone,

with implications for intermediate-depth seismicity

and earthquake hazard

Bradley R. Hacker1, Geoffrey A. Abers2 and Simon M. Peacock3

1 Department of Geological Sciences, University of California, Santa Barbara, California, 93106–9630, USA

2 Department of Earth Sciences, Boston University, 675 Commonwealth Avenue, Boston, Massachusetts, 02215, USA

3 Department of Geological Sciences, Arizona State University, Box 871404, Tempe, Arizona, 85287–1404, USA

 hacker@geology.ucsb.edu, abers@bu.edu, peacock@asu.edu

ABSTRACT

We calculate physical properties of rocks in subduction

zones using a compilation of mineral physical property

measurements, a new set of phase diagrams and subduction-zone thermal models, and apply our results to

Cascadia. Observed P-wave speeds of the Juan de Fuca

plate west of the Cascadia trench are best matched by a

greenschist-facies upper crust and metastable, partially

hydrated lower crust and mantle. The predicted positions

of dehydration reactions in the subducting crust show a

spatial link to observed seismicity, suggesting that dehydration may be the cause of the earthquakes and that the

predicted location of dehydration reactions in the slab can

be used to predict hazardous seismicity the length of

Cascadia.

Introduction

Much about subduction zones remains unknown because all our information comes either from indirect geophysical observations or observation of fossil subduction

zone rocks. In particular, we would like to understand:

the hydration state of the mantle wedge and the incoming slab; where and why slabs devolatilize and how this

contributes to arc volcanism; slab buoyancy and shape;

seismic wavespeeds and their relation to mineralogy; and

intermediate-depth earthquakes and their possible relationship to phase transformations. This study uses a synergistic thermal–petrological–seismological approach to

address some of these questions. Our approach comprises six specific steps:

1) Compile and assess physical properties of minerals

relevant to subduction zones.

2) Construct phase diagrams apropos to subduction zone

rock types and physical conditions.

3) Compute pressures (P) and temperatures (T) for a specific subduction zone.

4) Superimpose calculated phase relations onto the P–T

model.

5) Superimpose rock physical properties onto the P–T

model.

6) Compare predictions to observations.

1. Compiling mineral properties

We performed an extensive literature search to obtain

physical properties of minerals relevant to subduction

zones. The necessary physical properties are: formula

weight, molar volume, H2

O content, expansivity α, ∂α/

∂T, the isothermal bulk modulus KT, ∂KT/∂P, the shear

modulus μ, ∂μ/∂T, ∂μ/∂P, the Grüneisen parameter γth and

the second Grüneisen parameter δT. From these we used

a second-order finite-strain approximation (which yields

negligible error at these pressures for which P/KT < 1) to

calculate the adiabatic bulk modulus, shear modulus,

density, seismic wavespeeds and Poisson’s ratios for each

mineral as a function of P and T. Each of these values

was examined in some detail to ensure that the calculated values were in agreement with values measured

directly on minerals or single-phase mineral aggregates

at elevated P and T.

2. Constructing phase diagrams

We chose to simplify the rock compositions under consideration to basalt/gabbro, lherzolite, harzburgite and

serpentinite—i.e., the entire overriding plate crust and

subducting plate crust was treated as basalt/gabbro, and

the remainder was considered ultramafic. We then calculated the P–T stability fields of different minerals and

the reactions that bound the various fields. These phase

diagrams permit direct assessment of phase proportions,

reaction stoichiometries, H2

O content, and, in combination with the rock physical property calculations described above, density, VP

, and VS of mafic and ultramafic

rocks in subduction zones.
134 | Hacker, Abers and Peacock: Dehydration and seismicity of the Cascadia subduction zone

To create a phase diagram for mafic rocks, we performed an extensive literature search to identify studies

that reported all three of the following: 1) bulk rock compositions; 2) mineral modes; and 3) mineral compositions. From these studies we chose only rocks whose bulk

compositions differed from mid-ocean ridge basalts

(MORB) by less than ~10%. Assuming that the entire

crust is of MORB composition is clearly incorrect for the

lower crust, as detailed below. We then compiled the

mineral modes and mineral compositions into various

metamorphic facies. Next, the P–T range over which each

mineral assemblage is stable was defined using

Thermocalc [Holland and Powell, 1998]. Most boundaries are thus defined by discontinuous reactions involving the appearance or disappearance of at least one phase.

Obvious outliers were discarded, and a mean mineral

mode and set of mineral compositions was computed

for each metamorphic facies.

Phase diagrams for lherzolite, harzburgite and

serpentinite were constructed principally from Thermocalc [Holland and Powell, 1998] using Mg-endmember

reactions in the Ca–Fe2+–Mg–Al–Si–H–O system and

mineral compositions reported in the literature for naturally metamorphosed ultramafic rocks. Thermocalc’s calculations of phase relations at P >5 GPa were modified

in light of recent experiments by Luth [1995], Ulmer and

Trommsdorff [1995], Wunder and Schreyer [1997], Bose

and Navrotsky [1998], Wunder [1998] and Pawley [2000].

3. Computing subduction-zone pressure

and temperature

For Cascadia, we used the thermal model of Wang et al.

[1995] calculated for southern Vancouver Island (Figure

1a). Pressures for the same cross section were calculated

using appropriate rock densities.

4. Superimposing phase relations

Onto a subduction-zone cross section depicting P and T

and rock composition, we overlayed the different metamorphic mineral assemblages computed in step 2. Figures 1b and 1c show the results for mafic rocks and harzburgite, respectively. We have assumed that the activity

of H2

O = 1 (or that PH2O= Plithostatic) and that equilibrium

obtains. These assumptions cannot be correct everywhere

and are addressed partially in a later section.

5. Superimposing rock physical properties

From the mineral physical properties calculated at elevated P and T in step 1, we derived rock physical properties using a Voigt–Reuss–Hill average weighted by

mineral proportions determined in step 2. The results

are shown in Figures 1d–1g.

6. Comparing predictions to observations: Cascadia

Beneath Cascadia, the 5–9 Ma Juan de Fuca plate

subducts at a rate of 41 mm/yr [Hyndman and Wang,

1993], rendering it among the youngest, most slowly

subducting, and, consequently, warmest slabs in the

world. The slab has a thick sediment blanket and upper

plate volcanism is weak to moderate.

Seismicity: the correlation to dehydration

In the vicinity of southern Vancouver Island, seismicity

along and near the intraplate subduction thrust occurs

in two distinct pulses—one at 30–40 km depth and another around 70 km depth [Rogers et al., 1990] (Figure

1a). It is a primary observation of this study that there is

a spatial correlation between this seismicity and the predicted dehydration reactions in the subducting crust

(Figure 1b). The upper nest of earthquakes at 30–40 km

depth falls within the lower half of the epidote amphibolite field and stops abruptly at the down dip end at

the transition to zoisite–amphibole eclogite. This corresponds to a predicted change from 2.1 to 0.7 wt. %

crystallographically bound H2

O. The lower nest of earthquakes at ~70 km depth coincides with the transition

from zoisite–amphibole eclogite to anhydrous eclogite—

a reduction in H2O content from 0.3 to 0.1 wt. %. In

other words, along the south Vancouver Island transect,

the bulk of the seismicity occurs at depths coincident

with predicted dehydration in the downgoing crust. This

same spatial correlation between phase transformations

in the crust and intermediate-depth seismicity holds for

the Nankai (southern Japan) and Tohoku (northern

Japan) subduction zones where thermal conditions are

different. We propose, therefore, that calculating the

position of dehydration reactions in the slab could be an

inexpensive and effective means of predicting future

hazardous seismicity; certainly this hypothesis can easily

be tested with additional modeling.

The idea that intermediate-depth slab seismicity might

be related to dehydration reactions in the upper part of

the slab was explored in considerable detail by Kirby et

al. [1996], Peacock and Wang [1999] and Kirby [2000].

Our calculations for Cascadia provide a very strong confirmation of this hypothesis. Moreover, the fact that the

seismicity correlates well with phase changes predicted

from equilibrium considerations implies that at least part

of the subducted crust is undergoing mineralogical change

in a near-equilibrium fashion. Abundant geologic evidence from localities around the globe demonstrates that

fine-grained, H2O-rich basalts transform to high-pressure

mineralogies faster than coarse-grained, H2O-poor

gabbros [Hacker, 1996]. Thus, in a subducting slab, we

expect phase transformations in the gabbroic layer to lag

those in the volcanic layer and produce dehydration at

greater depths than predicted by equilibrium considerations. Therefore it is likely that: 1) the observed seismicity is generated principally by dehydration reactions in

the volcanic layer of the subducted crust, which trans-
FIGURE 1: Calculated properties of the Cascadia subduction zone along a transect through southern Vancouver Island. a) Geology [Hyndman and Wang, 

et al., 1995]. b) Calculated phase relations in mafic crust and observed seismicity [Rogers et al., 1990]. c) Calculated phase relations in ultramafic roc

of minerals by percentage: anth, anthophyllite (amphibole); atg, antigorite (serpentine); br, brucite; clin, clinochlore (chlorite); en, enstatite (orthopyrox

sp, spinel; py, pyrope (garnet). d) Calculated H2O content. e) Calculated densities. f) Calculated 

P-wavespeeds. g) Calculated 

S-wavespeeds.
136 | Hacker, Abers and Peacock: Dehydration and seismicity of the Cascadia subduction zone

forms on a near-equilibrium basis to eclogite; and 2) the

subducted plutonic layer is nearly anhydrous, generates

few earthquakes and transforms to eclogite at greater

depth than predicted by equilibrium considerations.

Slab velocities and mineralogy

An ocean-bottom seismometer survey offshore and parallel to Vancouver Island by Clowes [this volume] can be

compared with our predicted wavespeeds for the Juan

de Fuca plate. Clowes’ interpretation (Figure 2a) is that

the Juan de Fuca plate includes: a one kilometer thick

layer 2 with VP=4.0 km/s; a four kilometer thick layer

3A with VP=6.2–6.9 km/s; a layer 3B split into a one

kilometer layer with VP=7.0–7.3 km/s above a two

kilometer layer with VP =7.3–8.0 km/s; and an underlying mantle with a P-wave speed of 8.2 km/s.

In our equilibrium model (Figure 2b), layers 2 and

3A have a greenschist-facies mineralogy and a P-wave

speed of 6.6 km/s. Layer 3B is at amphibolite facies and

has VP=6.5 km/s. The immediately underlying mantle is

predicted to be an anthophyllite–chlorite-bearing harzburgite with a P-wave speed of 7.6 km/s.

Thus, we do not successfully model layer 2, but this

may be because we do not consider the effects of porosity and fluids. At greater pressures where fluids are expelled and pores closed, our model may be correct. The

faster-than-predicted speeds (>6.6 km/s) in the base of

layer 3A imply that the lower part of that layer is not at

greenschist facies. Studies of ophiolites imply that the most

likely candidate is unaltered basalt/gabbro. If we model

layers 3A and 3B as unaltered, metastable basalt/gabbro

of MORB composition, we calculate a speed of 6.9 km/s

(Figure 1c). Speeds of 7.0–7.3 km/s in the upper one kilometer of layer 3B are also faster than in our model, implying that layer 3B is not amphibolite-facies gabbro.

Moreover, the observed speeds are faster than our calculated speeds for metastable gabbro of MORB composition. Such speeds are, however, compatible with calculated speeds for pyroxene gabbro or olivine gabbro. Speeds

of 7.3–7.8 km/s in the lower one kilometer of layer 3B

can be modeled as a mix of olivine gabbro, wehrlite and

clinopyroxenite. Our calculated speed for anthophyllite–

chlorite-bearing harzburgite is slow compared to the observed seismic wave speed and the reported VP implies a

mix of anhydrous harzburgite and dunite.

Trying to match the measured seismic P-wavespeeds

with specific homogeneous rock types is fraught with

uncertainty. However, the best match is attained assuming that: 1) layer 3A is greenschist facies at the top and

grades downward into an anhydrous, metastable basalt;

2) layer 3B is not MORB-composition gabbro; and 3)

the mantle is metastable, anhydrous harzburgite. Alternatively, the velocity gradient observed by Clowes [this

volume] at the bottom of the crust may reflect lateral

and/or vertical resolution limits of the seismic data and a

more complex crust-mantle boundary than assumed.

Conclusions

Our model produces calculated physical properties of

rocks in subduction zones using a compilation of mineral

FIGURE 2: Observed and calculated P-wavespeeds for the Juan de Fuca plate offshore Vancouver Island. a) Seismic section from

Clowes [2000]. b) P-wavespeeds calculated assuming aH2O=1. c) P-wavespeeds calculated assuming aH2O=0. Calculations in b) and c)

are based on pressures and temperatures from most-outboard portion of Hyndman and Wang’s [1993] thermal model. Best fits to

seismic observations are indicated with boldface type.
The Cascadia Subduction Zone and Related Subduction Systems | 137

physical property measurements, a new set of phase diagrams and subduction-zone thermal models. Comparison of predicted vs. observed P-wavespeeds indicates that

the lowermost crust and mantle of the Juan de Fuca plate

west of the Cascadia trench is not at mineralogical equilibrium with the ambient P and T and, specifically, is not

fully hydrated. A distinctive correlation between the predicted positions of dehydration reactions in the subducting

crust and observed seismicity suggests a causal link between the two and that future hazardous seismicity the

length of Cascadia can be predicted via our method.

Acknowledgements

Thanks to Steve Kirby and Kelin Wang and their respective institutions for hosting the Victoria Workshop. Steve

was also the first to suggest the use of mapping dehydration reactions in the slab as an indicator of future

seismicity.

References

Bose, K., and A. Navrotsky, Thermochemistry and phase

equilibria of hydrous phases in the system MgO–SiO2

–

H2O; implications for volatile transport to the mantle, J. Geophys. Res., 103, 9713–9719, 1998.

Clowes, R.M., Crustal structure of northern Juan de Fuca

plate and Cascadia subduction zone—new results,

old data, in The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K.

Wang, and S.G. Dunlop, pp. 55–58, U.S. Geological

Survey Open-File Report 02–328, Geological Survey

of Canada Open File 4350, 2002.

Hacker, B.R., Eclogite formation and the rheology, buoyancy, seismicity and H2O content of oceanic crust, in

Dynamics of Subduction, edited by G.E. Bebout, D.

Scholl, S.H. Kirby, and J.P. Platt, pp. 337–246, American Geophysical Union, Washington, D.C., 1996.

Holland, T.J.B., and R. Powell, An internally consistent

thermodynamic data set for phases of petrological

interest, J. Metamorph. Geol., 16, 309–343, 1998.

Hyndman, R.D., and K. Wang, Thermal constraints on

the zone of possible major thrust earthquake failure

on the Cascadia margin, J. Geophys. Res., 98, 2039–

2060, 1993.

Kirby, S.H., Taking the temperature of slabs, Nature, 403,

31–33, 2000.

Kirby, S.H., E.R. Engdahl, and R. Denlinger, Intermediate-depth intraslab earthquakes and arc volcanism as

physical expressions of crustal and uppermost mantle

metamorphism in subducting slabs, in Subduction Top

to Bottom, Geophysical Monograph, edited by G.E.

Bebout, D. Scholl, and S. Kirby, pp. 195–214, American Geophysical Union, Washington, D.C., 1996.

Luth, R.W., Is phase A relevant to the Earth’s mantle?,

Geochimica et Cosmochimica Acta, 59, 679–682,

1995.

Pawley, A., Stability of clinohumite in the system MgOSiO2

-H2O, Contributions to Mineralogy and Petrology, 138, 284–291, 2000.

Peacock, S.M., and K. Wang, Seismic consequences of

warm versus cool subduction metamorphism: examples from southwest and northeast Japan, Science,

286, 937–939, 1999.

Rogers, G.C., C. Spindler, R.D. Hyndman, and J.L. Varsek,

Seismicity along the Vancouver Island Lithoprobe

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Ulmer, P., and V. Trommsdorff, Serpentine stability related to mantle depths and subduction-related magmatism, Science, 268, 858–861, 1995.

Wang, K., T. Mulder, G.C. Rogers, and R.D. Hyndman,

Case for very low coupling stress on the Cascadia subduction fault, J. Geophys. Res., 100, 12,907–12,918,

1995.

Wunder, B., Equilibrium experiments in the system MgOSiO2-H2

O (MSH); stability fields of clinohumite-OH

[Mg9Si4O16(OH)2], chondrodite-OH ([Mg5Si2O8(OH)2]

and phase A (Mg7Si2O8(OH)6), Contributions to Mineralogy and Petrology, 132, 111-–20, 1998.

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138 |
The Cascadia Subduction Zone and Related Subduction Systems | 139

Slab deformation at the edge of subduction zones:

An example from Cascadia

Uri S. ten Brink

 United States Geological Survey, 384 Woods Hole Road, Woods Hole, Massachusetts, 02543, USA

TenBrink@usgs.gov

Slab deformation

Understanding the interaction between the subducting

plate and the overlying forearc is a key to a successful

evaluation of the seismogenic hazard above subduction

zones. Here we focus on the interaction, which takes

place when subduction ceases. In particular, the cessation of subduction at the southern end of the Cascadia

subduction zone offers insights to the processes that operate in a warm subduction zone and the coupling between the plates. We show that the legacy of subduction

remains in the thermal structure of California even after

subduction ceased.

Plate kinematic interpretations predict a gap in the

underlying subducted slab caused by the northward migration of the Pacific–North America (NOAM)–Juan de

Fuca (JdF) triple junction (MTJ). Upon cessation of subduction, the weight of the already subducted plate and

of the adjacent subducting section continued to pull the

slab downward, opening a gap into which sublithospheric

material (asthenosphere) upwelled under the former

forearc [Dickinson, 1997; Dickinson and Snyder, 1979a].

Various predictions have been made on the basis of the

slab gap hypothesis. The slab gap was envisioned to be

the locus of decompression melting in the asthenosphere,

resulting in the accretion of a several kilometers thick

layer of mafic material to the base of the California crust

[Liu and Furlong, 1992]. Estimates of hydrocarbon maturation were made based on the thermal regime of a slab

gap with asthenoshperic upwelling [Heasler and Surdam,

1985]. The Pacific–NOAM plate boundary was envisioned

to be subhorizontal at depth and to migrate inland over

time as the slab gap annealed [Furlong et al., 1989].

The slab-gap model is based on the configuration of

three independently-moving rigid plates, Pacific, NOAM,

and JdF, but, in fact, the Pacific plate is the only rigid

plate. Non-rigid deformation of the JdF plate, known as

the Gorda deformation zone, is manifested by curved

magnetic anomalies and pervasive faulting [Wilson,

1986]. The NOAM plate in California also does not behave as a single rigid plate. It consists of two blocks, the

Sierra Nevada–Great Valley Klamath Mountains block

and the Coast Ranges. The latter is essentially a deformation zone between two more competent blocks, the

Pacific plate and the Sierra Nevada block [Walcott, 1993].

The slab gap model also assumes that the three plates

are moving independently. This has not been the case

for the JdF plate after the fragmentation of the Farallon

plate. Starting 18 Myr ago, the JdF plate has rotated clockwise, seafloor spreading slowed [Atwater, 1989], and the

JdF plate motion is progressively more aligned with that

of NOAM with respect to the Pacific.

We propose an alternative model based on kinematics of non-rigid plate boundaries [ten Brink et al., 1999].

In this model, continuous deformation, rotation and

stretching of the JdF plate fill the geometrical gap without large-scale holes and tears. The large and rigid Pacific plate is acting as an indenter into the non-rigid JdF

plate. It is moving northwestward in a hot spot reference

frame pushing at the JdF plate across the Mendocino

fracture zone and causing it to move northward. Although

the fracture zone may be weak in shear, north–south

oriented normal compressive stresses can be transmitted to the southern part of the JdF plate [Wang et al.,

1997]. Landward of the MTJ, however, the subducting

Juan de Fuca plate is not pushed by the Pacific plate and

is not confined. This allows the plate to spread into the

unconfined region.

How much does the subducted plate have to stretch

to fill the geometrical gap? Using two different plate kinematic configurations, the estimated size of the gap to

be filled is only 27% or 37% of the JdF plate area subducted during the past 4 and 3.2 Myr respectively. If the

JdF slab stretches to fill this gap, it will have to thin by a

factor of only 1.27 or 1.37. The age of the subducted

plate is 6–10 Ma [Wilson, 1986]; hence, the thickness of

the thermal lithosphere (to a temperature of 1300°C) is

25–30 km. Thinning by a factor of 1.37 reduces the slab’s

thermal thickness to 18.1–21.7 km, which is equivalent

to the thickness of a 4 Ma slab. (Thinning of the subducted

slab by a factor of 1.27 reduces the thermal thickness of

the lithosphere to 19.7–23.6 km, which is equivalent to

the thickness of a 4–5 Ma slab.) Thinning of the mantle

at these shallow depths (40–50 km) by this factor is ex-
140 | ten Brink: Slab deformation at the edge of subduction zones: an example from Cascadia

pected to generate only negligible decompression melting of the upper mantle [McKenzie and Bickle, 1988].

Strain may not be uniform, however, so larger amounts

of thinning may take place locally as the tomographic

images suggest.

Thermal and magmatic consequences

With the slab stretch model in mind, we propose that

the thermal evolution of the lithosphere results from thermal re-equilibration following the cessation of subduction and some stretching of the slab. The overlying plate

is the forearc region of the subduction zone, which is

insulated during subduction by the cold surface of the

subducting plate beneath the forearc. When subduction

and stretching stop, thermal reequilibration begins, and

temperatures at the base of the forearc rise. We approximate this process by calculating the geothermal gradients as a function of time after cessation of subduction

(time 0) for subducted slabs with a thermal age of 10

and 4 Ma overlain by 20 and 30 km thick forearc.

There are two important consequences of thermal reequilibration: 1) melting of the subducted crust; and 2)

elevated heat flow in the overlying plate. Minor volcanic

activity occurred in the wake of the northward passage

of the MTJ, reflected in the younger volcanic outcrops

toward the north [Dickinson, 1997]. The most recent

activity (<2 Ma) is located at Clear Lake. It has been suggested that the volcanism resulted from a slab gap under

the forearc [Dickinson and Snyder, 1979b; Johnson and

O’Neil, 1984]. Quantitative estimates of asthenospheric

upwelling into the gap, however, predict voluminous

magmatic production [Liu and Furlong, 1992], far larger

than the geologic data suggest. In the slab stretch model,

we explain the volcanism as mainly arising from crustal

anatexis. Pressure–temperature conditions of the

amphibolite solidus (temperatures ≥ 630°C at depths ≥ 25

km) [Wyllie and Wolf, 1993] can be attained within the

subducted oceanic crust during the first 5 Myr after the

cessation of subduction of a hot young (~4 Ma) oceanic

lithosphere overlain by 20–30 km thick crust. Magmatism

in California is typically short lived and produces small

volumes of magma. It occurs mainly in the eastern part

of the Coast Ranges [Johnson and O’Neil, 1984] where

the overlying crustal thickness reaches 20–25 km

[Henstock et al., 1997]. A seismic profile across northern

California where the MTJ was located <2 Myr ago shows

“bright spots” representing melt lenses [Levander et al.,

1998] that cluster throughout the fossil oceanic crust. The

volcanic rocks are dominantly intermediate to silicic, have

relatively high 87Sr/86Sr ratios, relatively high δ18O values

(7.5–11.5), and relatively high Th contents [Johnson and

O’Neil, 1984], suggesting that crustal anatexis played a

dominant role in their generation.

Surface heat flow in western California increases rapidly immediately south of the MTJ, reaching a maximum

several hundreds of kilometers to the south and then

decreasing slightly. Our calculations best fit the data with

an overlying crustal thickness of 20 km and an underlying stretched oceanic slab with an equivalent thermal

age of 2–4 Ma. The pattern of increasing heat flow south

of the MTJ was previously cited as evidence for a slab

gap [Lachenbruch and Sass, 1980]. Interestingly, however, these authors also pointed out that the heat flow

variations are best fit with cooler than astheonospheric

temperatures (1000°C) emplaced instantaneously at the

base of a 20-km-thick crust. Our proposed geothermal

gradient also matches thermobarometric estimates from

spinel lhezorlites from Coyote Reservoir [Jove and

Coleman, 1998].

A broad zone of high heat flow across the entire Coast

Ranges averaging 30 mW m-2 higher than is typical for

continental regions was explained as a result of either: 1)

shear heating on horizontal planes due to viscous drag of

the upper part of the lithosphere over its substratum

[Lachenbruch and Sass, 1973]; 2) shear heating on vertical planes in the lithospheric upper mantle along the Pacific–North American diffuse plate boundary [Molnar,

1992]; or 3) the effect of the slab gap [Lachenbruch and

Sass, 1980]. Central California is probably underlain by

the stalled Monterey microplate [Nicholson et al., 1994],

and therefore a slab gap never existed in central California [Atwater and Stock, 1998]. We argue that shear heating is also not required to explain the elevated heat flow.

The higher heat flow is simply the consequence of reequilibration of the geothermal gradient following the

cessation of subduction of the Monterey microplate. The

stalled Monterey slab underlying the Coast Ranges was

probably less than 4 Myr old, assuming a convergence

rate of 30–40 km Myr-1. The predicted heat flows for a 2-

and 4 Myr old subducted slabs 18 Myr after subduction

matches well the observations from this area [Sass et al.,

1997]. Note that the Monterey microplate behavior prior

to cessation was similar to the JdF plate, namely, clockwise rotation and decreased subduction rate.

References

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and western North America, in The Eastern Pacific

Ocean and Hawaii, edited by E.L. Winterer, D.M.

Hussong, and R.W. Decker, pp. 29–72, Geol. Soc. Am.,

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Atwater, T., and J. Stock, Pacific–North America plate

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an update, Int. Geol. Rev., 40, 375–402, 1998.

Dickinson, W.R., Tectonic implications of Cenozoic

volcanism in coastal California, Geol. Soc. Am. Bull.,

109, 936–954, 1997.

Dickinson, W.R., and W.S. Snyder, Geometry of

subducted plates related to San Andreas transform,

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USGS OPEN FILE REPORT #: The Cascadia Subduction Zone and Related Subduction Systems | 141

Dickinson, W.R., and W.S. Snyder, Geometry of triple

junctions related to San Andreas transform, J.

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Furlong, K.P., W.D. Hugo, and G. Zandt, Geometry and

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maturation, AAPG Bull., 69, 1386–1400, 1985.

Henstock, T.J., A. Levander, and J.A. Hole, Deformation

in the lower crust of the San Andreas fault system in

northern California, Science, 278, 650–653, 1997.

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rocks in western California, Earth Planet. Sci. Lett.,

71, 241–262, 1984.

Jove, C.F., and R.G. Coleman, Extension and mantle

upwelling within the San Andreas fault zone, San

Francisco Bay area, California, Tectonics, 17, 883–

890, 1998.

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aspects of the San Andreas fault system, in Tectonic

Problems of the San Andreas Fault System, edited by

R.L. Kovach, and A. Nur, pp. 192–205, Stanford

University Press, Stanford, Calif., 1973.

Lachenbruch, A.H. and J.H. Sass, Heat flow and

energetics of the San Andreas fault zone, J. Geophys.

Res., 85, 6977–6989, 1980.

Levander, A., T.J. Henstock, A.S. Meltzer, B.C. Beaudoin,

A.M. Trehu, and S.L. Klemperer, Fluids in the lower

crust following Mendocino triple junction migration:

active basaltic intrusion? Geology, 26, 171–174, 1998.

Liu, M., and K.P. Furlong, Cenozoic volcanism in the California Coast Ranges: numerical solutions, J. Geophys.

Res., 97, 4941–4951, 1992.

McKenzie, D., and M.J. Bickle, The volume and composition of melt generated by extension of the lithosphere,

J. Petrol., 29, 625–679, 1988.

Molnar, P., Brace–Goetze strength profiles, the partitioning of strike-slip and thrust faulting at zones of oblique

convergence, and the stress–heat flow paradox of the

San Andreas fault, in Fault Mechanics and Transport

Properties of Rocks, edited by T.F. Wong and B. Evans,

pp. 435–459, Academic, San Diego, Calif., 1992.

Nicholson, C., C.C. Sorlien, T. Atwater, J.C. Crowell, and

B.P. Luyendyk, Microplate capture, rotation of the

western Transverse Ranges, and initiation of the San

Andreas transform as a low-angle fault system, Geology, 22, 491–495, 1994.

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and F.V. Grubb, Thermal regime of the San Andreas

fault near Parkfield, California, J. Geophys. Res., 102,

27,575–27,585, 1997.

ten Brink, U.S., N. Shimizu, and P.C. Molzer, Plate deformation at depth under northern California, slab gap

or stretched slab? Tectonics, 18, 1084–1098, 1999.

Walcott, D., Neogene tectonics and kinematics of western North America, Tectonics, 12, 326–333, 1993.

Wang, K., J. He, and E.E. Davis, Transform push, oblique

subduction resistance, and intraplate stress of the Juan

de Fuca plate, J. Geophys. Res., 102, 661–-674, 1997.

Wilson, D.S., A kinematic model for the Gorda deformation zone as a diffuse southern boundary of the

Juan de Fuca plate, J. Geophys. Res., 91, 10,259–

10,269, 1986.

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405–416, Geol. Soc. of London, London, 1993.
142 |
The Cascadia Subduction Zone and Related Subduction Systems | 143

Stresses in the Juan de Fuca plate and the role

of mantle resistance to horizontal slab motion

Kelin Wang1, Jiangheng He1, Earl E. Davis1 and Chris Goldfinger2

1 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

2 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97331, USA

wang@pgc.nrcan.gc.ca, he@pgc.nrcan.gc.ca, davis@pgc.nrcan.gc.ca, gold@oce.orst.edu

The Juan de Fuca plate is a small oceanic plate between

the Pacific and North America plates. In the southernmost

region, referred to as the Gorda deformation zone, the

maximum compressive stress σ1 constrained by earthquake focal mechanisms is north–south. Off Oregon, and

possibly off Washington, northwest trending left-lateral

faults cutting the Juan de Fuca plate indicate a σ1 in a

northeast–southwest to east–west direction. The magnitude of differential stress increases from north to south;

this is inferred from the plastic yielding and distribution

of earthquakes throughout the Gorda deformation zone.

To understand how tectonic forces determine the stress

field of the Juan de Fuca plate, we have modeled the

intraplate stress using both elastic and elastic perfectlyplastic plane-stress finite element models. We conclude

that the right-lateral shear motion of the Pacific and North

America plates is primarily responsible for the stress pattern of the Juan de Fuca plate. The most important roles

are played by a compressional force normal to the

Mendocino transform fault, a result of the northward push

by the Pacific plate, and a horizontal resistance operating against the northward, or margin-parallel, motion of

the slab. Margin-parallel slab resistance results in a large

north–south compression in the Gorda deformation zone

because the force is integrated over the full length of the

Cascadia subduction zone. The Mendocino transform

fault serves as a strong buttress that is very weak in shear

but capable of transmitting large strike-normal compressive stresses. Internal failure of the Gorda deformation

zone potentially places limits on the magnitude of the

fault-normal stresses being transmitted and correspondingly on the magnitude of strike-parallel subduction resistance. Transform faults and oblique subduction zones

in other parts of the world can be expected to transmit

and create stresses in the same manner. Margin-parallel

slab resistance is comparable to mantle’s updip resistance to slab motion and therefore is comparable to the

slab-pull force. It must play a major role in controlling

the state of stress in the slab and the occurrence of

intraslab earthquakes.
144 |
The Cascadia Subduction Zone and Related Subduction Systems | 145

Phase changes, fluids and the co-location of the deep

and shallow seismicity beneath

Puget Sound and southern Georgia Strait

Garry C. Rogers

 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

rogers@pgc.nrcan.gc.ca

The majority of contemporary seismicity in the northern

half of the Cascadia subduction zone occurs in one location: straddling the international border, adjacent to

the change in trend of the North American coastline from

north–south to northwest–southeast. This is true in both

the subducting oceanic Juan de Fuca plate and in the

overlying continental North American plate. Why is the

region of crustal seismicity located directly above the

subcrustal seismicity and why is the areal extent very

similar, even though the two regions are decoupled and

have very different stress regimes? The upward migration of fluids offers a potential explanation. In the

subducting oceanic crust, the phase change from basalt

to eclogite is coincident with the region of seismicity.

This phase change creates extensional stress in the

subducting oceanic crust and alters the subducting plate

buoyancy from positive to negative, adding to the tension near the upper surface. The phase change from

hydrous metabasalt to eclogite also releases water. In

warm subduction zones such as Cascadia, this phase

change takes place over a very narrow range in the downdip direction. The broad arch in the subducting plate,

created by the change in the orientation of the subducting

margin, allows margin parallel migration of hot fluids to

the highest region: beneath Puget Sound and southern

Georgia Strait. The crustal seismicity in the North America

plate is characterized by patterns and lineaments, but

there is little evidence of fault planes aligning with spatial trends of epicentres. Instead, most crustal seismicity

seems to be occurring on random faults, all responding

to the same regional stress. Seismicity induced by fluid

injection often exhibits this characteristic.

FIGURE 1: This cartoon explains the hypothesis for the colocation of the deep and shallow seismicity. Studies in the

oceanic crust offshore have shown that the flow of water

in hydrothermal systems can operate over great many tens

of kilometers. The tensional environment created by the

phase change provides paths for the abundant water supplied by the phase change process to migrate to the top of

the arch in the subducting plate, where it enhances the earthquake generating process. From the top of the arch, the

water bleeds upwards into the North American plate where

it causes a region of increased micro-seismicity.
146 |
The Cascadia Subduction Zone and Related Subduction Systems | 147

The role of phase changes

in the development of forearc basins

Garry C. Rogers

 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

rogers@pgc.nrcan.gc.ca

As the oceanic lithosphere is subducted into the mantle,

the changing pressure and temperature conditions cause

phase changes that result in denser material. The largest

density change takes place in the oceanic crust as basalts

and gabbros change to eclogite. This causes the thinner

crustal portion of the descending lithosphere to contract

significantly more than the mantle portion. The net result is a vertical contraction of the subducted crust. In

warm subduction zones, this phase change takes place

in a very narrow range in the down-dip direction and

appears to be reflected as a down-drop in the topography, forming a basin in the forearc region. Evidence for

such down-drops can be seen as linear inland seas paralleling the subduction front in the warm subduction

regimes of Cascadia, the Nankai region of Japan, and in

southern Chile. In Cascadia, the Georgia Strait–Puget

Sound–Willamette Valley lowland is directly above the

expected location of the basalt to eclogite phase change

region. The rapid onset of the basin formation 85 million years ago coincides with a dramatic decrease in the

age of oceanic lithosphere being subducted as the

Farallon–Kula triple junction migrated along the coast,

changing the subduction regime from a cold to a warm

one. This phase change process may be an important

mechanism for the initiation of forearc basins.

FIGURE 1: Forearc basins above

phase-change regions in warm subduction zones can be seen in the

forearcs of Cascadia (Georgia Strait/

Puget Sound/Willamette Valley),

southern Chile (Gulf of Corcovado/

Great Longitudinal Valley), and in the

southeast region of Japan where the

Philippine Sea plate subducts (Seto

Inland Sea). The light green/white regions indicate the extent of modern

forearc basins.
148 |
The Cascadia Subduction Zone and Related Subduction Systems | 149

ABSTRACT

In-slab earthquakes (earthquakes within the subducting

Juan de Fuca plate) make the major contribution to seismic hazard for the Strait of Georgia region of British Columbia. These earthquakes dominate the hazard, despite

their depth, because they have a higher rate and cause

stronger shaking than the crustal earthquakes. Key knowledge of in-slab earthquakes needed to improve seismic

hazard estimates for southwestern British Columbia includes the constraints on the spatial distribution, rate and

maximum size of the earthquakes, the ground motions

to be expected, the nature of the earthquake sources and

the structure and properties of the lithosphere through

which the waves propagate.

Introduction

Seismic hazard for the Strait of Georgia region of British

Columbia (including Vancouver, Victoria and a substantial fraction of B.C.’s population) comes from three sources:

crustal seismicity in the North American plate, great earthquakes of the Cascadia subduction zone on the interface between the North American and subducting Juan

de Fuca plate, and deep earthquakes within the subducting slab (“in-slab” earthquakes). It is, however, dominated by the contribution from in-slab earthquakes. In

Canada’s fourth generation seismic hazard model (see

Adams et al., 1999a, 1999b, 2000), these earthquakes

dominate the hazard despite their greater depth, firstly

because they occur at a rate up to five-fold higher per

unit area than the shallower crustal earthquakes, and

secondly, because their predicted shaking is stronger than

crustal events of the same size (see below). Thus when

attempts are made to improve the estimation of seismic

hazard for southwestern B.C., a great deal can be gained

by better understanding these earthquakes.

We raise the following series of questions to highlight

the knowledge of in-slab earthquakes we believe is

needed to improve seismic hazard estimates for southwestern British Columbia. Some of the differences that

result from the current level of uncertainty are demonstrated on Figure 1, a comparison of the GSC and USGS

deaggregated hazard for Bellingham, Washington. Clear

differences are seen in the fraction of the hazard coming

from in-slab versus crustal earthquakes, and in the contribution from earthquakes larger than magnitude 7.

What is the spatial distribution likely for future

earthquakes within the slab?

The GSC’s fourth generation hazard maps use three

source zones to model deep earthquake distribution,

Georgia Strait (GEO) and Puget Sound (PUG) for one

probabilistic model and Georgia Strait/Puget Sound (GSP)

for the other, reflecting uncertainty in the future location

of damaging deep earthquakes (Figure 2). What is needed

are geological or geophysical reasons to constrain the

updip, downdip, northern and southern extent of the

deep seismicity. Although the two probabilistic models

attempt to model the range of possible distributions, the

level of hazard is strongly controlled by the active PUG

zone. This is especially important for Vancouver, as the

northern boundary of PUG lies under the city and generates a steep gradient in hazard across the city (Figure

3). Fairly large changes in hazard for communities in this

gradient zone could result from slight adjustments to the

source zone boundary, perhaps as the result of new significant earthquake activity outside the currently-defined

PUG zone, or a recognition that certain regions within

the boundary are (and will continue to be) aseismic. Improved geological/geophysical constraints might identify

these regions and so refine the hazard estimates.

What is the rate of activity?

The rate of large earthquakes is a function of the rate of

activity for small earthquakes (a-value or alpha for the

magnitude-recurrence curve) and the slope of the magnitude-recurrence curve (b-value or beta). Alpha is quite

well determined in aggregate, but it is unknown whether

it truly varies in space (as it appears to during the historical record), and if so, why it should vary. The GSC uses a

source zone approach which assumes uniform rates

within each source zone (which may not be valid); the

USGS uses spatial smoothing of past activity, which assumes that the locations of future large earthquakes will

Knowledge of in-slab earthquakes needed to improve

seismic hazard estimates for southwestern British Columbia

John Adams and Stephen Halchuk

Geological Survey of Canada, 7 Observatory Crescent, Ottawa, Ontario, K1A 0Y3, Canada

adams@seismo.nrcan.gc.ca, halchuk@seismo.nrcan.gc.ca
150 | Adams and Halchuk: Knowledge of in-slab earthquakes needed to improve seismic hazard estimates

precisely mimic the smoothed distribution of the small

earthquakes (which may not be valid either).

The slope of the curve (beta) represents the relationship between the number of small and big earthquakes.

For PUG, it is distinctively flatter than for most crustal

source zones such as the crustal earthquake zone, Cascade Mountains R model (CASR), which overlies it (Figure 4). Two curves are shown for the crustal earthquakes

(CASR), one representing the mathematical fit to the observed rates and the other (dashed) accommodating the

observed higher rate of M>6.5 earthquakes. The value

used in the Canadian hazard model is much lower

(beta=1.01, b=0.44) than that used by the USGS

(beta=1.5, b=0.65) for its deep earthquakes. No sound

explanations exist for the different empirical values of

beta, though a study of worldwide in-slab earthquakes

might confirm the reasonableness of the value chosen,

and provide insight into the reasons for such a low value.

Together, the magnitude recurrence parameters explain some of the difference in hazard. Figure 4 shows

the activity rates of PUG and the overlying CASR crustal

earthquake source. At magnitude 6, the predicted rate

of in-slab earthquakes is three to ten times the rate of

crustal earthquakes, thus accounting for the larger hazard contribution from the former. On Figure 4, the curve

representing the USGS slope is drawn through the magnitude 4 data point on our PUG magnitude recurrence

curve. As to be expected, the steeper USGS slope predicts a rate of M>6 earthquakes only one-third the GSC

rate, and thus explains some of the hazard difference.

How large can the in-slab earthquakes get?

The largest historical in-slab event occurred in 1949, of

moment magnitude about 6.9. Compared with recent

earthquakes, almost nothing is known about the rupture

parameters of this earthquake, such as its depth extent,

fault length or stress drop. Some geophysical constraints

such as temperature in the slab are believed to limit the

thickness of brittle rock thus restricting fault width; larger

earthquakes therefore require greater fault lengths or

greater slip (or both). The GSC model currently allows

an upper bound magnitude of 7.3 for PUG (with an uncertainty range of 7.1–7.6) as shown on Figure 5, presuming that a future large earthquake could extend

deeper into the slab, or have larger displacement, or rupture a longer fault (perhaps through cascading rupture

segments as demonstrated during the Lander’s earthquake). In 1997, the USGS adopted an upper bound

magnitude of 7.0. Because of the high rate for these large

earthquakes (due to the small b value), their contribution to the total seismic hazard is not trivial (for the GSC’s

results, about 14–24% of the seismic hazard, dependent

on model, comes from earthquakes larger than the 1949

one). More work in understanding the 1949 and 1965

earthquakes together with the geological/geophysical

conditions might allow tighter constraints on the largest

possible earthquakes.

How reliable are the current

strong ground motion relations?

Both the GSC and USGS use the Youngs et al. [1997]

relations to compute seismic hazard from the in-slab

earthquakes. These relations concluded that in-slab earthquakes produce ground motions 40% larger than ground

motions from adjacent subduction interface earthquakes

(Figure 6), but this is not completely accepted. On the

one hand, the Youngs et al. [1997] relations have been

criticized as being based on rather sketchy data and upon

no long period data at all [Atkinson and Boore, this volume]; on the other, the qualitative differences in damage between interface and in-slab earthquakes [e.g., Okal

and Kirby, this volume] argue that there is almost certainly a quantitative difference in excitation, perhaps even

larger than 40%. Considerable work is needed to deterFIGURE 1: Seismic hazard deaggregations of 0.2 second spectral

acceleration values at 2%/50 years for Bellingham show the GSC

results are dominated by the contribution from in-slab earthquakes, unlike the 1997 USGS results.
The Cascadia Subduction Zone and Related Subduction Systems | 151

FIGURE 2: Selected in-slab earthquakes (>35 km) in the Puget

lowlands/southwestern B.C. and the alternative source zones used

to model them for the GSC’s fourth generation seismic hazard

maps.

FIGURE 3: Hazard map from the GSC model (‘H’) using the Puget

Sound (PUG) source. Contours, for 0.2 second spectral acceleration and 2%/50 year probability, are in % g. Note how the steep

gradient near Vancouver is dependent on the position of the PUG

boundary.

mine if the 40% “premium” for in-slab earthquakes is

realistic, implausible, or too small, and whether the premium applies to all periods or just to the shorter ones.

The comparison of the in-slab and crustal (using the Boore

et al., [1997] relations) earthquake motions (Figure 6) indicates that at essentially all the distances significantly

contributing to the hazard, the ground motions from a

50-km-deep in-slab earthquake are expected to exceed

those from a similar sized 10-km-deep crustal earthquake.

What are the typical seismic sources

we have to contend with?

Our knowledge of the seismic source can affect our decision on which strong ground motion relations to use.

Most earthquakes will probably have normal faulting

mechanisms, but undetermined is the degree to which

rupture directivity effects are important, particularly if

ruptures tend to rupture upwards from their nucleation

point (Figure 7).

If as a first approximation, the in-slab earthquakes

are described as Brune sources, what are their stress

drops? If as a refinement, they are described as realistic,

elasto-dynamic sources, what are the key parameters (e.g.

rupture velocity, source elongation, complete or fractional

stress drop, source complexity/episodic rupture, fault

roughness, etc.) that affect the spectral shapes of the

source as radiated towards the overlying urban areas?

Do in-slab source acceleration spectra have intermediate (omega-1) slopes, and if so, over what frequency band?

Haddon [1996] showed that typical Mw=6 eastern earthquake sources have omega-1 slopes for about one decade of frequency above a lower corner, and that the high

frequency (f>1 Hz) levels exceed those associated with

a Brune model for a Mw=6, 100 bar stress drop event by

a factor of three, and approach those for a Brune model

source a full magnitude larger (see the velocity spectra

on Figure 8). The intermediate slopes are consequent

on high rupture velocities, rupture directivity effects involving asymmetrical ruptures, episodic ruptures and partial stress drop events. Therefore, given records of small

earthquakes, source scaling parameters correctly incorporating these factors are needed to synthesize the

ground motions for potentially damaging earthquakes.
152 | Adams and Halchuk: Knowledge of in-slab earthquakes needed to improve seismic hazard estimates

What are the crustal/mantle properties (e.g. Q,

velocity layering, dipping layers) that affect the

radiated energy between the source and the site

where hazard is needed?

A reliable interpretation of crustal and mantle properties

is needed to assess and adjust the strong ground motion

relations and to perform forward modeling to determine

the consequences of scenario earthquakes. For example, if crustal conditions differ significantly from Mexico,

a source of much in-slab earthquake data, how do we

adjust strong ground motion parameters derived from a

worldwide dataset?

What earthquake scenario should be

adopted for Vancouver and Victoria?

How can the use of empirical Green’s functions

improve hazard estimates?

Deaggregations like Figure 1 indicate the magnitude and

distance of the earthquakes contributing to the seismic

hazard and are the starting point for design earthquake

scenarios. Use of empirical Green’s functions can improve hazard estimates (by effectively accounting for all

path complexity), but still require much knowledge about

the seismic source so that the source scaling can be done

FIGURE 4: Magnitude-recurrence curves and observed activity

rates (dots with error bars) for in-slab (red) and crustal (black)

earthquakes for the Puget Sound. Both CASR curves have been

reduced by a factor of 6.2 to account for the larger area of CASR

relative to the PUG (see inset). The scaled USGS relation (blue)

for deep earthquakes is also shown.

FIGURE 5: Magnitude-recurrence curve for PUG (like Figure 4),

showing the upper and lower uncertainty bounds and range of

upper bound magnitudes. For comparison, the curve used for the

USGS calculations is shown in blue.

appropriately. Hence, future improvements will depend

critically on our ability to understand what will happen

during the larger earthquakes, and our best insight to

that will come from analysis of the past large Puget Sound

earthquakes.

Conclusions

Different assumptions were adopted by the USGS in 1997

and the GSC in 1994–1999 and resulted in different estimates of seismic hazard for the U.S. and Canadian territory overlying these in-slab earthquakes. Reconciling these

estimates and refining them towards the true hazard will

involve better answers to the questions raised above.

References

Adams, J., D.H. Weichert, and S. Halchuk, Trial seismic

hazard maps of Canada 1999: 2%/50 year values for

selected Canadian cities, 107 pp., Geological Survey

of Canada Open File 3724, 1999a.

Adams, J., D.H. Weichert, and S. Halchuk, Lowering

the probability level—fourth generation seismic hazard results for Canada at the 2% in 50 year probability level, in Proceedings 8th Canadian Conference on

Earthquake Engineering, Vancouver 13–16th June

1999, pp. 83–88, 1999b.
The Cascadia Subduction Zone and Related Subduction Systems | 153

.

.

.

FIGURE 6: A comparison of expected ground motions from adjacent interface and in-slab earthquakes, to those from similar-sized

crustal earthquakes (black).

FIGURE 7: Ground motions above a typical in-slab normal faulting

earthquake may depend critically on the rupture plane location,

rupture plane dip direction and the asymmetry of the rupture relative to the hypocenter, all contributing to directivity effects.

FIGURE 8: According to Haddon, the velocity spectra of typical eastern earthquake sources are flat for about one decade of frequency

above their lower corner, and their high frequency (f>1 Hz) levels exceed those associated with the corresponding Brune model event

by a factor of three, approaching the shaking of a Brune event one magnitude larger (redrawn from Haddon, [1996]).
154 | Adams and Halchuk: Knowledge of in-slab earthquakes needed to improve seismic hazard estimates

Adams, J., S. Halchuk, and D.H. Weichert, Lower probability hazard, better performance? Understanding

the shape of the hazard curves from Canada’s Fourth

Generation seismic hazard results, Paper 1555, 12th

World Conference on Earthquake Engineering, Auckland, 30th January–4th February 2000, 8 pp., 2000.

Atkinson, G.M., and D.M. Boore, Empirical ground motion relations for subduction-zone earthquakes in The

Cascadia Subduction Zone and Related Subduction

Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 155–158, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Boore, D.M., W.B. Joyner, and T.E. Fumal, Equations for

estimating horizontal response spectra and peak acceleration from western North America earthquakes:

a summary of recent work, Seism. Res. Lett., 68, 128–

153, 1997.

Haddon, R.A.W., Earthquake source spectra for Eastern

North America, Bull. Seism. Soc. Amer., 86, 1300–

1313, 1996.

Okal, E., and S.H. Kirby, Energy–to–moment ratios for

damaging intraslab earthquakes: preliminary results in

The Cascadia Subduction Zone and Related Subduction Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 117–122, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Youngs, R.R., S.-J. Chiou, W.J. Silva, and J.R. Humphrey,

Strong ground motion relationships for subduction

zone earthquakes, Seism. Res. Lett., 68, 58–73, 1997.
The Cascadia Subduction Zone and Related Subduction Systems | 155

Empirical ground motion relations

for subduction-zone earthquakes

Gail M. Atkinson and David M. Boore

Department of Earth Sciences, Carleton University, 1125 Colonel By Drive, Ottawa, Ontario, K1S 5B6, Canada

gma@ccs.carleton.ca

Ground motion relations for subduction-zone earthquakes are an important input to seismic hazard analyses

for regions affected by the Cascadia subduction zone

(Washington, Oregon and British Columbia). There is a

significant hazard potential from great thrust events along

the subduction interface and from large events within

the subducting slab. We have compiled a response spectra database from thousands of strong motion recordings from events of moment magnitude (M) 5 to 8.3

occurring in subduction zones around the world, including both interface and in-slab events. The distribution of

the database in magnitude and distance (where closest

distance to fault surface is used) is shown in Figure 1.

We are currently refining the database and performing

regression analyses to determine empirical ground motion relations for interface and in-slab events. The large

dataset enables better determination of attenuation parameters and magnitude scaling, for both types of events,

than has previously been possible. Soil response parameters, including nonlinear effects at strong levels of shaking, can also be determined from the data.

Preliminary regression results suggest that in-slab events

attenuate more rapidly with distance than do interface

events. For both in-slab and interface events, ground

motion amplitudes depend on focal depth, with deeper

events showing larger amplitudes. In-slab events cause

greater ground motion amplitudes at close distances than

do subduction (interface) events but the overall effect is

countered by the more rapid attenuation of in-slab motions with distance.

Figure 2 compares the preliminary relation for a subduction thrust (interface) event of M=7.9 at a focal depth

of about 20 km, to response spectra data in the range

from M=7.5 to 8.3 (focal depths<35 km), for a frequency

of 1 Hz (5% damped pseudo-acceleration [PSA] in cm/s2,

horizontal component). This type of event is a major

contributor to the 1 Hz hazard in the Cascadia region.

The relation of Youngs et al. [1997] is also plotted. The

Youngs et al. [1997] subduction relations are commonly

used for seismic hazard analysis in the region; they were

developed in the early 1990s from the much more limited dataset that was available at that time. Figure 2 shows

that the AB2000 and Y97 relations are similar for this

type of event and that both tend to underestimate the

motions in this particular magnitude range. Because the

regression relations fit a large range of magnitudes, distances, depths and event types, it is not an unusual circumstance to obtain such a misfit within a specified range

of parameters. However, this magnitude range is important and refinements in the regression results to address

this mismatch are being pursued.

Figure 3 compares the preliminary relation for an inslab event of M=6.8 at a focal depth of about 60 km for

a frequency of 5 Hz, to response spectra data in the range

from M=6.5 to 7.1 (focal depths 40–60 km). This type of

event is also important for hazard analysis. The preliminary relation agrees reasonably well with the overall trend

of the data but underpredicts the sparse near-fault amplitudes. Again, refinements to better fit this important

magnitude-distance range are in progress.

Looking at a broad range of magnitudes and distances,

the preliminary results suggest that the subduction ground

motion relations of Youngs et al. [1997] may tend to overestimate the motions caused by moderate-to-large in-slab

events (at all distances) but underestimate the motions

from great subduction thrust events at large distances. This

finding, if verified by more detailed regression analyses

currently in progress, could have significant implications

for seismic hazard analysis in the Cascadia region.

References

Youngs, R., S. Chiou, W. Silva, and J. Humphrey, Strong

ground motion attenuation relationships for subduction-zone earthquakes, Seism. Res. L., 68, 58–73,

1997.
156 | Atkinson and Boore: Empirical ground motion relations for subduction-zone earthquakes

FIGURE 2: AB2000: Atkinson and Boore, this paper; Y97: Youngs et al., 1997.

FIGURE 1
The Cascadia Subduction Zone and Related Subduction Systems | 157

FIGURE 3: AB2000: Atkinson and Boore, this paper; Y97: Youngs et al., 1997.
158 |
The Cascadia Subduction Zone and Related Subduction Systems | 159

Deep intraslab earthquakes and how they contribute

to seismic hazard in the Pacific northwestern U.S.

Mark D. Petersen and Arthur D. Frankel

United States Geological Survey, PO Box 25046, Lakewood, Colorado, 80225, USA

mpetersen@usgs.gov, afrankel@usgs.gov

ABSTRACT

We investigate the impact of two different source and

attenuation models for deep intraslab earthquakes to the

seismic hazard along the northwestern coast of the United

States. Different source models for deep intraslab earthquakes may cause either localized hazard in the lower

Puget Sound of Washington and northwestern California,

or hazard that is spread out parallel to the Cascadia subduction zone. These source models result in hazard estimates for 2% probability of exceedance in 50 years that

differ by as much as 25% near Portland, Oregon. Applying maximum magnitudes (M=7.0–7.4) to the intraslab

source model results in differences of less than 10% for

the hazard estimates in the Pacific Northwest. Attenuation models [e.g., Youngs et al., 1997; Atkinson and Boore,

this volume] that relate the earthquake magnitude and

distance to the ground motion, lead to differences of 50%

in hazard for one second spectral acceleration near the

earthquake sources. The variability in hazard for different

source models and attenuation relations indicates the importance of identifying the physical mechanisms that drive

the intraslab events and understanding the variability in

the ground motions.

Introduction

Seismic hazard along the northwestern coast of the U.S.

is dominated by contributions from crustal faults, shallow

crustal seismicity, deep intraslab seismicity and great earthquakes along the Cascadia subduction interface [Frankel

et al., 1996]. The deep intraslab events contribute significantly to the hazard in the Puget lowlands in the state of

Washington, as can be observed from the deaggregation

of the seismic hazard (Figure 1) [Harmsen and Frankel,

2001]. In fact, the deep intraslab earthquakes contribute

31% and 24% of the hazard at 0.2 and 1 second for the

5% damped spectral acceleration, respectively, for 10%

probability of exceedance in 50 years. For the 2% probability of exceedance in 50 years hazard, the contribution from these deep events is not as great in Seattle

because the crustal faults become more important.

The hazard questions that arise from observations of

deep earthquakes are:

1) Why are the historical deep intraslab events restricted

to the Puget lowlands and northwestern California and

are not observed beneath the state of Oregon?

2) What is the long-term rate of intraslab earthquakes?

3) What is the maximum size of earthquakes in the

subducting crust?

4) What is the variability in ground motions allowed by

strong ground motion recordings of intraslab earthquakes?

The answers to these questions are critical in determining the characteristics of earthquake sources that will

be used by engineers to design earthquake resistant structures. Therefore, it is imperative that we try to answer

these questions since they will impact public safety in

this region.

In this article we present two source models that take

into account various assumptions about the physical

mechanism of strain release in intraslab earthquakes associated with the Cascadia subduction zone. We test these

models to understand the sensitivity of the different source

model parameters to the hazard. In addition, we evaluate the impact of the maximum magnitude and two different attenuation relations for deep intraslab earthquakes

on the seismic hazard in the Pacific Northwest.

Source models for deep intraslab earthquakes

The 1996 national seismic hazard maps [Frankel et al.,

1996] and the Oregon Department of Transportation

report [Geomatrix, 1995] both recognized the importance

of deep intraslab earthquakes in the northwestern United

States. The national seismic hazard map implemented a

source model for the intraslab earthquakes that is based

on smoothed seismicity of historical events since 1949.

The depths of these earthquakes are poorly constrained,

so all of the depths were fixed at 45 km for this hazard

analysis. Earthquake sources were smoothed using a

Gaussian smoothing operator with a correlation distance

of 50 km [Frankel et al., 1996]. This smoothing process

assumes that future earthquakes will occur near the locations of historical earthquakes and results in localized

hazard beneath the Puget lowlands and northwestern
160 | Petersen and Frankel: Deep intraslab earthquakes and how they contribute to seismic hazard

California. The resulting hazard maps are shown in Figures 2a and 2c.

An extreme opposite model is to consider that deep

intraslab earthquakes could occur with equal likelihood

anywhere along the Cascadia subduction zone. Because

we have not observed deep earthquakes beneath Oregon,

this model would imply that we have not recorded earthquakes for long enough to recognize deep earthquakes in

that region. The hazard resulting from this extended source

zone model is lower in areas where we have observed

historical intraslab earthquakes and is higher in areas above

the slab where we have no evidence of such earthquakes

(compare Figures 2a and 2b). For example, if we smooth

the earthquakes over a large zone parallel to the Cascadia

subduction zone, it raises the overall hazard in parts of

western Oregon from about 0.35 g to 0.45 g for 2% probability of exceedance in 50 years. However, this model

also decreases the hazard in the places where we have

observed historic intraslab earthquakes such as portions

of the Puget lowlands (compare Figures 2c and 2d). Therefore, questions arise whether this extended source model

is reasonable for public policy hazard maps.

Maximum magnitude

Two deep intraslab earthquakes have caused damage in

the Seattle–Tacoma–Olympia region during the past century. These earthquakes occurred in 1949 (Mb

=7.1) and

in 1965 (Mb

=6.5), and indicate the hazard from future

deep events in the Pacific Northwest. For calculation of

the national seismic hazard maps, a maximum magnitude of Mw=7.0 was assumed (note that this is moment

magnitude). The question remains “Is the 1949 earthquake the largest possible intraslab earthquake?” We have

tested the sensitivity of the maximum magnitude of the

intraslab earthquake to the hazard by varying the maximum magnitude from Mw=7.0 to 7.4. Results of this sensitivity study indicate that increasing the maximum

magnitude to Mw=7.4 raises the hazard by 10% when

only considering the intraslab earthquakes, and less than

5% in the overall national seismic hazard maps including all sources. Therefore, this parameter by itself does

not make a large difference in the current hazard estimates for the Pacific Northwest of the United States.

Attenuation relations

Once the source model is constructed, the ground

motion from each earthquake is estimated through an

attenuation equation that relates the magnitude and distance to a particular ground motion parameter (e.g., peak

horizontal ground acceleration or spectral acceleration).

The 1996 hazard model applied only the Youngs et al.

[1997] attenuation relationship to calculate the ground

motion from the intraslab earthquakes, since it was the

only one available at that time. Recently another attenuation relationship has been generated by Atkinson and

Boore [this volume] that uses additional strong ground

motion measurements to update the relationship. We

applied the Youngs et al. [1997] and the Atkinson and

Boore [this volume] attenuation relations for rock site

conditions at one second spectral acceleration to calculate the hazard in the Pacific Northwest. Atkinson and

Boore [this volume] do not report an uncertainty so we

used the uncertainty given in the Youngs et al. [1997]

relationship that is about a factor of two in size but is

dependent on earthquake magnitude. We applied a fixed

depth of 45 km for the deep earthquakes. The hazard

for one second spectral acceleration at 2% probability of

exceedance in 50 years is about 50% lower when using

the Atkinson and Boore [this volume] relation than when

using the Youngs et al. [1997] relation (Figure 3).

Discussion and conclusions

The differences between the hazard calculated using the

Atkinson and Boore [this volume] attenuation relation

for one second spectral acceleration differs significantly

from the hazard generated using the Youngs et al. [1997]

FIGURE 1: Deaggregation of hazard in the 1996 National Seismic

Hazard Map for 10% probability of exceedance in 50 years at a

site in Seattle, Washington. A) shows the hazard for the 0.2 second spectral acceleration and B) shows the hazard for the 1 second spectral acceleration.
The Cascadia Subduction Zone and Related Subduction Systems | 161

FIGURE 2: Seismic hazard maps produced using: A) 1996 intraslab source model; B) extreme opposite model with extended intraslab

source zone; C) full 1996 model with intraslab source model shown in A; and D) 1996 model with intraslab source model shown in B.
162 | Petersen and Frankel: Deep intraslab earthquakes and how they contribute to seismic hazard

equation. The hazard may vary up to 50% or more in

many areas near the sources. It is important to determine what level of ground shaking will occur from future deep earthquakes in this area and correctly weight

the attenuation relationships in the hazard analysis. This

is one of the most important considerations in calculating the hazard for deep earthquakes in this region.

Near Seattle Washington, the hazard is higher when

assuming a model that uses the historic seismicity and

the hazard is slightly lower when assuming a model that

spreads the seismicity along the Cascadia zone. However, near Portland Oregon, the hazard is 25% higher

when assuming the model that spreads the seismicity out

along the deep Cascadia slab. The latter model gives an

upper bound for hazard in Portland and western Oregon

from deep earthquakes. We need to focus our scientific

efforts on understanding the physical mechanisms of these

earthquakes to be able to understand the validity of these

models. Questions still remain regarding why the deep

earthquakes tend to be correlated with shallow crustal

earthquakes and why these deep earthquakes tend to

occur near the ends of the subduction zone or where the

fault changes strike or dip. The answer to these questions

would allow us to constrain our source models for deep

intraslab earthquakes and in the resulting hazard maps.

Acknowledgments

We want to thank Steve Harmsen for providing the

FIGURE 3: Seismic hazard maps for 2% probability of exceedance in 50 years for deep earthquakes in the Pacific Northwest on a

uniform rock site condition. Hazard calculated using the one second spectral acceleration of A) Youngs et al. [1997] and B) Atkinson

and Boore [this volume].

deaggregation data used to produce Figure 1 and Gail

Atkinson for providing a preliminary version of her attenuation relation for deep intraslab earthquakes. We

thank Chris Cramer for a thoughtful review.

References

Atkinson, G.M., and D.M. Boore, Empirical ground motion relations for subduction-zone earthquakes in The

Cascadia Subduction Zone and Related Subduction

Systems, edited by S.H. Kirby, K. Wang, and S.G.

Dunlop, pp. 155–158, U.S. Geological Survey OpenFile Report 02–328, Geological Survey of Canada

Open File 4350, 2002.

Harmsen S.C., and A.D. Frankel, Geographic deaggregation of seismic hazard in the United States, Bull.

Seismol. Soc. Am., 91, 13–26, 2001.

Frankel A., C. Mueller, T. Barnhard, D. Perkins, E.V.

Leyendecker, N. Dickman, S. Hanson, and M. Hopper, National Seismic Hazard Maps, June 1996 Documentation, U.S. Geological Survey Open File Report

96–532, 1996.

Geomatrix Consultants, Seismic Design Mapping State of

Oregon, prepared for Oregon Department of Transportation Transportation Building Salem, OR 97310, 1995.

Youngs R.R., S.J. Chiou, W.J. Silva, and J.R. Humphrey,

Strong ground motion attenuation relationships for

subduction zone earthquakes, Seis. Res. Lett., 68, 58–

73, 1997.
The Cascadia Subduction Zone and Related Subduction Systems | 163

The proposed NEPTUNE seafloor observatory

on the Juan de Fuca plate: Applications to

studies of intraplate earthquakes

William S.D. Wilcock1, Kelin Wang2, Roy D. Hyndman2, Anne M. Tréhu3 and Kenneth C. Creager4

1 School of Oceanography, University of Washington, Seattle, Washington, 98195, USA

2 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

3 College of Oceanic and Atmospheric Sciences, Oregon State University, Corvallis, Oregon, 97330, USA

4 Department of Earth and Space Sciences, University of Washington, Seattle, Washington, 98195–1650, USA

wilcock@ocean.washington.edu, wang@pgc.nrcan.gc.ca, hyndman@pgc.nrcan.gc.ca,

trehu@oce.orst.edu, kcc@ess.washington.edu

ABSTRACT

NEPTUNE is a plan to deploy a long-term fiber-optic observatory on the seafloor of the Juan de Fuca plate. The

network will support a wide range of scientific observations including a regional broadband seismic network and

multidisciplinary experiments to study hydrothermal

activity on the Juan de Fuca ridge and its flanks. Such

observations will contribute to our understanding of the

hydration of oceanic lithosphere and improve seismic images of the Cascadia subduction zone. In the long term,

NEPTUNE will contribute to a better understanding of the

mechanism of intraslab earthquakes.

Introduction

Oceanography is in the midst of a revolution that is fundamentally changing the means by which scientists observe many processes in and beneath the oceans [Ryan

et al., 2000]. Historically, ship-based expeditions have

characterized the oceans using data collected onboard

and with short deployments of autonomous instruments.

As scientific knowledge has advanced, it has become

increasingly apparent that many problems cannot be

addressed by time-limited visits alone. A new operational

paradigm is emerging made possible by advances in computational sophistication, communication and power

technologies, robotic systems and sensor design. Ocean

research is moving towards a mode of operations in which

ship-based studies will be complemented by long-term

observations from a network of coastal, regional and

global seafloor observatories.

The NEPTUNE project (http://www.neptune.-

washington.edu) is an ambitious plan to deploy a regional

submarine fiber-optic network on the Juan de Fuca plate

off the coast of the Pacific Northwest (Figure 1) [Delaney

et al., 2000; Delaney and Chave, 2000; NEPTUNE Phase

1 Partners, 2000]. The network will host scientific instrumentation to study a wide variety of geological, oceanographic and ecological processes. NEPTUNE will link

scientists and the public interactively to a wide range of

sensors, sensor networks and autonomous mobile platforms in, on and above the seafloor (Figure 2). It will

operate both large long-term experiments managed by

scientific communities and smaller experiments developed by individual scientists. NEPTUNE will also support extensive educational and outreach activities.

A conceptual design and engineering feasibility study

[NEPTUNE Phase 1 Partners, 2000] demonstrates that

there are no insurmountable technical obstacles. It is

envisioned that the NEPTUNE backbone will comprise

3000 km of cable connecting about 30 evenly distributed

primary nodes. Secondary branch cables can extend to

any location on the plate. At each node a junction box

will provide standard power, communication and timing interfaces for scientific instruments. The complete

network will carry 10 GB per second of data, deliver up

to 100 kW of power and have an operational life span of

at least 30 years. The node topology (Figure 1) includes a

branch that runs near the base of the continental shelf,

extending the length of the subduction zone from the

Mendocino transform fault to the Explorer plate, and

branches along the Juan de Fuca spreading center and

Blanco transform faults. There are two shore landings at

Nedonna Beach, Oregon and Victoria, British Columbia.

Seismic studies in NEPTUNE

In the field of marine seismology, the need for seafloor

observatories is well documented. Many seismology studies rely on global and regional networks of land-based

seismometers that have operated for several decades. Data

from these networks have been used to image the earth’s

interior and they provide a long-term view of seismic
164 | Wilcock et al.: The proposed NEPTUNE seafloor observatory on the Juan de Fuca plate

deformation. However, these networks have two major

limitations. First, images of the earth obtained from global networks are severely limited by the lack of instruments on the seafloor. A permanent seafloor seismometer

was recently attached to an abandoned telecommunications cable midway between Hawaii and the west coast

of the U.S. [Chave et al., 1998], but about 20 additional

sites will be required to fill all the oceanic gaps in global

networks [Orcutt and Purdy, 1995]. Second, most of the

Earth’s plate boundaries are located offshore and are

poorly observed by regional land networks. Several

cabled seismometers have recently been deployed off

the coast of Japan [e.g., Suyehiro and Fukao, 1998], but

studies of oceanic plate boundary seismicity are generally limited to teleseismic earthquakes or data collected

with short deployments of autonomous ocean bottom

seismometers.

NEPTUNE can help address this second limitation.

The Juan de Fuca–Gorda–Explorer plate system incorporates a remarkable array of plate tectonic features within

a relatively small area, including all the major types of

oceanic plate boundary. Although the subduction zone

plate boundary is seismically quiet, observations from land

and with the Navy’s Sound Surveillance System (SOSUS)

hydrophone arrays show extensive seismicity in the region (Figure 3). The seismic and geodetic component of

the proposed NEPTUNE network will provide an opportunity to study a wide range of geophysical processes

including the seismic potential of the Cascadia subduction zone, the mechanisms of plate deformation and

interactions, the structure and evolution of the lithosphere/

asthenosphere system, and the characteristics of earthquakes and geological processes at plate boundaries [NEPTUNE Phase 1 Seismology Working Group, 2000].

NEPTUNE and studies of intraslab earthquakes

In the past century, the most damaging earthquakes in

the Pacific Northwest have been intraslab events, earthquakes that occur within the subducted Juan de Fuca plate

at depths of 30–70 km. Examples include the magnitude

7.1 Olympia earthquake in 1949 and the magnitude 6.5

SeaTac earthquake in 1965. These earthquakes are believed to be related to metamorphic processes in the slab

[Kirby et al., 1996]. The oceanic crust created at the midocean ridges is hydrothermally altered and hydrated.

During subduction, the hydrated basalts and gabbros are

transformed into the higher density eclogite and this

induces deviatoric stresses. The metamorphic transformation also releases a significant amount of water (1–2%

by weight) that elevates pore fluid pressure. The combination of the stress and fluid pressure increases may

reactivate some pre-existing faults and cause earthquakes.

It has been well documented that the subduction

zone thermal regimes have a strong influence on intraslab

earthquake processes. Peacock and Wang [1999] studied the northeast Japan and southwest Japan subduction

zones, typical examples of cool and warm subduction

zones, and concluded that the higher temperature in

the southwest Japan subduction zone caused the basalteclogite transformation and hence intraslab earthquakes

to take place at much shallower depths. Other consequences include a shallow termination of a low seismic

velocity layer representing the untransformed crust,

sparse arc volcanism and possible melting of the slab.

The Cascadia subduction zone is a warm subduction zone

because of the young age of the subducting Juan de Fuca

plate and is very similar to southwest Japan.

NEPTUNE will provide an opportunity to improve

our understanding of intraslab earthquakes and will

allow us to assess their hazard along the Cascadia

margin. Studies of seafloor hydrothermal activity on the

Juan de Fuca ridge and its flanks will contribute to our

understanding of the degree and depth extent of the

hydration of the oceanic lithosphere prior to subduction. Long term monitoring of seafloor seismicity will yield

information on the distribution and extent of faults in

the subducting slab. In addition to constraining the characteristics of the plate prior to subduction, seafloor seismic observations can be used to study the slab after

FIGURE 1: Candidate scenario for the NEPTUNE network. A fiber

optic cable will connect primary nodes spaced about 100 km

apart. Power and communications will be routed through two

shore landings.
The Cascadia Subduction Zone and Related Subduction Systems | 165

subduction. For this margin, the intraslab stress regime is

rather poorly constrained at depths between 10 and 50

km, partly due to the lack of seismic observations offshore. By recording and analyzing seismic waves that

are generated by intraslab events and that propagate

along the slab, it will be possible to constrain the seismic

velocity structure of the slab and thus infer its thermal

structure and petrology. Such information is essential if

we are to constrain downdip and along-strike variations

in the metamorphism and earthquake processes in the

slab. In the long term, a better understanding of the

mechanism of intraslab earthquakes and the slab thermal and mechanical structure will facilitate a more rigorous hazard study for these earthquakes.

A NEPTUNE seismic network

Figure 4 illustrates a possible design for a NEPTUNE seismic network [NEPTUNE Phase 1 Seismology Working

Group, 2000]. It is based on a modified version of the

candidate node scenario (Figure 1) developed for the

conceptual design and feasibility study. It should be

emphasized that the final design will require extensive

planning to determine community priorities and ensure

that the network is optimized for seismic detection and

characterization. It is envisioned that a broadband seismometer will be deployed at each primary node. Because the Gorda and Explorer plates are seismically active,

the node topology of Figure 1 has been modified to

improve coverage in these regions. An additional 30 to

40 broadband instruments will be required to fill large

gaps in the network and span the plate boundaries. Strong

motion accelerometers and hydrophones will also be

deployed throughout the region. Local networks of

intermediate band sensors will be required at some nodes

for high-resolution observations in regions of particular

geological interest.

Future plans

Upon completion of the conceptual design and feasibility

study in 2000, NEPTUNE commenced a phase of fundraising and detailed design that will lead to deployment

as early as 2005. A consortium of five partners (University

of Washington, Jet Propulsion Laboratory, Monterey Bay

Aquarium Research Institute, Woods Hole Oceanographic

Institution and the Institute of Pacific Ocean Science and

Technology [Canada]) are leading the effort to design and

FIGURE 2: Essential elements of NEPTUNE. Land based scientists and the public are linked in real time interactively with sensors,

sensor networks and various mobile platforms in, on and above the seafloor, NEPTUNE's fiber optic/power cable and associated

technology provide the enabling network infrastructure.
166 | Wilcock et al.: The proposed NEPTUNE seafloor observatory on the Juan de Fuca plate

prototype the power, communications and data management systems. Test beds are planned for Monterey Bay

and Georgia Basin to evaluate the system performance

and to test scientific sensors.

The feasibility study supported a number of small ad

hoc scientific working groups but the detailed engineering design must be guided by much more extensive

scientific planning. About ten open workshops will be

held to develop the motivation and priorities for community experiments. These deliberations will lead to the

formation of working groups with members drawn from

the whole community. Each working group will be

responsible for the design and implementation of an experiment. The workshops on seismology and geodynamics, ridge crest processes and crustal hydrology will be

of interest to those studying intraplate earthquakes.

References

Chave, A.D., F.K. Duennebier, R. Butler, R.A. Petitt Jr., F.B.

Wooding, A.D. Bowen, D. Harris, and D.R. Yoerger,

H2O: the Hawaii-2 Observatory, Eos Trans. AGU, 79,

F65, 1998.

Delaney, J.R., and A.D. Chave, NEPTUNE: a fiber-optic

‘telescope’ to inner space, Oceanus, 42, 10–11, 2000.

Delaney, J.R., G.R. Heath, B. Howe, A.D. Chave, and H.

Kirkham, NEPTUNE: real-time ocean and earth sciences at the scale of tectonic plate, Oceanography,

13(2), 71–79, 2000.

FIGURE 3: Earthquake epicenters for the period August 29, 1991 to December 31, 1997, determined using T-phases recorded by

underwater hydrophones in the U.S. Navy's Sound Surveillance System (data courtesy of Chris Fox, Pacific Marine Environmental

Laboratory, National Oceanic and Atmospheric Administration).
The Cascadia Subduction Zone and Related Subduction Systems | 167

Kirby, S., E.R. Engdahl and R. Denlinger, Intermediatedepth intraslab earthquakes and arc volcanism as

physical expressions of crustal and uppermost mantle

metamorphism in subducting slabs, in Subduction: Top

to Bottom, edited by E. B. Bebout et al., pp. 195–214,

American Geophysical Union, Washington DC, 1996.

NEPTUNE Phase 1 Seismology Working Group, Opportunities for Seismology and Geodynamics in NEPTUNE,

Online report, http://gore.ocean.washington.edu/

neptune, 2000.

NEPTUNE Phase 1 Partners (University of Washington,

Woods Hole Oceanographic Institution, Jet Propulsion Laboratory, Pacific Marine Environmental Laboratory), Real Time, Long-term Ocean and Earth Science

Studies at the Scale of a Tectonic Plate—NEPTUNE

Feasibility Study (prepared for National Oceanographic Partnership Program), 106 pp., University of

Washington, 2000.

Orcutt, J.A., and G.M. Purdy, Broadband Seismology in the

Oceans—Towards a Five-year Plan, 105 pp., Joint Oceanographic Institutions Inc., Washington, DC, 1995.

Peacock, S.M., and K. Wang, Seismic consequences of

warm versus cool subduction metamorphism: examples from southwest and northeast Japan, Science,

286, 937–939, 1999.

Ryan, W.B. F.R.S. Detrick, K. Becker, J. Bellingham, R.

Lukas, J. Lupton, L. Mullineaux, and J. Sipress, Illuminating the Hidden Planet: the Future of Seafloor Observatory Science, 135 pp., National Academy Press,

Washington, D.C., 2000.

Suyehiro, K., and Y. Fukao, Seafloor observatories: Japanese experience, Eos Trans. AGU, 79, F62–F63, 1998.

FIGURE 4: Possible NEPTUNE seismic network design based on a modified version of the node topology used for the conceptual design

and engineering feasibility study. Yellow symbols are used to depict primary nodes whose location have changed from Figure 1.
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The Cascadia Subduction Zone and Related Subduction Systems | 169

POLARIS: New opportunities in Cascadia

John F. Cassidy1, Michael G. Bostock2, Gail Atkinson3, David W. Eaton4, Isa Asudeh5, David B. Snyder5,

John Adams5, Ian J. Ferguson6, Alan G. Jones5 and Colin J. Thomson7

1 Pacific Geoscience Centre, Geological Survey of Canada, PO Box 6000, Sidney, British Columbia, V8L 4B2, Canada

2 Earth and Ocean Sciences, The University of British Columbia, 6339 Stores Road, Vancouver, V6T 1Z4, Canada

3 Department of Earth Sciences, Carleton University, 1125 Colonel By Drive, Ottawa, Ontario, K1S 5B6, Canada

4 Department of Earth Sciences, University of Western Ontario, London, Ontario, N6A 5B8, Canada

5 Geological Survey of Canada, 7 Observatory Crescent, Ottawa, Ontario, K1A 0Y3, Canada

6 Department of Geological Sciences, University of Manitoba, Winnipeg, Manitoba, R3T 2N2, Canada

7 Department of Geological Sciences, Queens University, Kingston, Ontario, K7L 3N6, Canada

cassidy@pgc.nrcan.gc.ca, bostock@geop.ubc.ca, gma@ccs.carleton.ca, deaton@uwo.ca, jadams@NRCan.gc.ca,

IJ_Ferguson@UManitoba.CA, thomson@geol.queensu.ca

POLARIS (Portable Observatories for Lithospheric Analysis and Research Investigating Seismicity) is a multi-institutional C$10 million project recently funded by the

Canadian Foundation for Innovation, provincial governments and universities, and private industry across

Canada. This project will provide new, state-of-the-art

portable geophysical observatories for research into

lithospheric structure, continental dynamics, and earthquake hazards in Canada. POLARIS infrastructure will

comprise 90 three-component broadband seismographs

and 30 magnetotelluric (MT) mobile field systems and

complementary satellite telemetry data acquisition. Over

the initial four year installation of this project, the seismograph network will be deployed as three subarrays of

30 instruments each. The MT instruments will be used

in shorter-term deployments at each of the three seismic subarrays for lithospheric imaging, and continuous

recording at selected elements for deep-mantle imaging.

The initial scientific objectives include:

1) three-dimensional detailed mapping of the asthenosphere and upper mantle of the Slave Province in

Canada’s north to assist the emerging diamond industry;

2) mapping lithospheric structure and earthquake hazards in the heavily populated area of southern Ontario

(identifying zones of crustal weakness, obtaining accurate earthquake parameters, and ground motion

attenuation studies).

3) mapping the structure and earthquake hazards over

the Cascadia subduction zone in southwest British

Columbia (subducting oceanic plate, site-response in

urban areas, identifying active crustal faults, and seismic attenuation studies).

The latter array will provide new opportunities into research involving slab seismicity, including focal mechanisms, attenuation studies, and detailed structural studies.

Further research initiatives that will be possible include:

testing rapid warning systems for ground shaking in the

urban areas of Canada; developing new fine-scale imaging

techniques using the scattered wavefield; and investigating geomagnetically induced currents that can cause

major disruptions to electrical and pipeline infrastructure.
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