Characteristics of Hawaiian Volcanoes

Professional Paper 1801

U.S. Department of the Interior

U.S. Geological Survey
Halema‘uma‘u lava lake and Kīlauea Caldera, with a snowcapped Mauna Loa in the background. Untitled painting by D. Howard

Hitchcock, 1894. Image courtesy of the National Park Service, Hawai‘i Volcanoes National Park, HAVO catalog number 452. 

COVER

Summit eruptive vent of Kīlauea Volcano and the southwest flank of Mauna Loa, as seen from Waldron Ledge. The small cluster 

of lights on the right side of the image is the Hawaiian Volcano Observatory and the Jaggar Museum. USGS photograph by M.P. 

Poland, February 4, 2011. 
Characteristics of Hawaiian Volcanoes

Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski, editors

U.S. Department of the Interior

U.S. Geological Survey

Professional Paper 1801
U.S. Department of the Interior

SALLY JEWELL, Secretary

U.S. Geological Survey

Suzette M. Kimball, Acting Director

U.S. Geological Survey, Reston, Virginia: 2014

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Poland, M.P., Takahashi, T.J., and Landowski, C.M., eds., 2014, Characteristics of Hawaiian volcanoes:

U.S. Geological Survey Professional Paper 1801, 428 p., http://dx.doi.org/10.3133/pp1801.

ISSN 1044-9612 (print) 

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http://dx.doi.org/10.3133/pp1801
iii

Foreword

The Hawaiian Islands and their volcanoes have featured prominently in the history of 

the United States Geological Survey (USGS) nearly back to the 1879 founding of the 

organization. In 1882, USGS Director John Wesley Powell sent Captain Clarence E. Dutton, 

an officer in the United States Army who was detailed to the USGS, to Hawai‘i—then still 

an independent kingdom—to study its volcanic geology in preparation for mapping in the 

Cascade Range. Dutton was an inspired choice for the assignment. He was already well 

known for his explorations in the western United States, thanks in large part to his vivid 

written accounts of the Grand Canyon region, and his observations of the volcanoes, land, 

and people of Hawai‘i after 4 months of field work (published as part of the “4th Annual 

Report of the U.S. Geological Survey” in 1884) are no less engaging. Dutton’s experience 

in Hawai‘i was a great aid to his subsequent assignment as the head of the USGS Division 

of Volcanic Geology, which mapped volcanoes throughout California, Oregon, Washington, 

Utah, Arizona, and New Mexico.

USGS work in Hawai‘i subsequently shifted toward water resources, especially as related 

to agricultural development. In 1909, USGS geologist Walter Mendenhall toured the islands 

and established a framework for systematic observations that were eventually assumed by 

what had become the Territory of Hawaii. In 1919, the Territory requested a comprehensive 

assessment of the geology and water resources of the entire island chain. One of the main 

participants in this work was USGS geologist Harold T. Stearns. Over the ensuing 30 years, 

Stearns published 12 comprehensive reports (Hawaii Division of Hydrography Bulletins) 

covering the characteristics of every major Hawaiian island (except Kauai, which was 

covered in 1960 in volume 13 by another longtime USGS geologist, Gordon Macdonald). 

The work of Stearns and his colleagues has stood the test of time and is still an important 

resource for geologists working throughout the State.

In 1924, the USGS took over operation of the Hawaiian Volcano Observatory (HVO), 

renewing its commitment to the study of Hawaiian geology. Founded in 1912 at the edge of 

the caldera of Kīlauea Volcano, HVO was the vision of Thomas A. Jaggar, Jr., a geologist 

from the Massachusetts Institute of Technology, whose studies of natural disasters around 

the world had convinced him that systematic, continuous observations of seismic and 

volcanic activity were needed to better understand—and potentially predict—earthquakes 

and volcanic eruptions. Jaggar summarized the aim of HVO by stating that “the work 

should be humanitarian” and have the goals of developing “prediction and methods of 

protecting life and property on the basis of sound scientific achievement.” These goals 

align well with those of the USGS, whose mission is to serve the Nation by providing 

reliable scientific information to describe and understand the Earth; minimize loss of life 

and property from natural disasters; manage natural resources; and enhance and protect 

our quality of life. In fact, Jaggar and the USGS expanded volcano monitoring in the 1920s, 

establishing an observatory at Lassen Peak, California, and a seismic monitoring station 

at Dutch Harbor, Alaska. These efforts and a planned expansion of nationwide volcano 

monitoring, with HVO at its hub, were derailed by the Great Depression and World War 

II, and the National Park Service took over operation of HVO during 1935–47. The USGS 

returned to administer HVO in 1947, continuing to oversee the observatory’s evolution 

from humble beginnings as a largely one-person operation into a world-class laboratory for 

volcano and earthquake research and a critical facility for natural hazards assessment and 

mitigation in Hawai‘i.
vi

In 2012, HVO celebrated the centennial of its founding. In the more than 100 years since 

Jaggar began making systematic observations of Hawaiian volcanism, HVO has been 

responsible for numerous innovations and scientific insights into natural hazards and Earth 

processes. For example, in the 1920s and 1930s, HVO scientists made the first forecasts 

of tsunami arrival times from distant earthquakes. The development of modern seismic 

networks was started, in large part, by the work of HVO scientist Jerry P. Eaton during the 

1950s. HVO has also served as a training ground for volcanologists from the United States 

and around the world, following the example of Dutton’s visit to Hawai‘i as a means of 

preparing him for mapping volcanoes elsewhere in the country. As awareness of volcanic 

hazards in the United States grew—especially following the reawakening and eruption of 

Mount St. Helens in 1980—four new observatories were established to monitor volcanic 

activity in Alaska, California, Yellowstone, and the Cascade Range of Washington and 

Oregon. HVO’s legacy as a training ground for volcanologists continues even to this day, 

with the observatory participating in the development of scientists from other countries as 

part of an international volcano-monitoring education program run by the Center for the 

Study of Active Volcanoes (a cooperative effort of HVO, the University of Hawai‘i at Hilo, 

and the University of Hawai‘i at Mānoa and founded by former HVO Scientist-in-Charge 

Robert Decker).

The legacy of USGS scientists in Hawai‘i goes far beyond scientific accomplishments. In 

fact, USGS employees who spent time in Hawai‘i have helped to shape the organization on 

a national level, and many have gone on to assume high-level positions within the USGS. 

This long history of symbiosis has been mutually beneficial—studies of Hawai‘i and its 

volcanoes have had a strong impact on the direction of USGS science, and USGS practices 

have shaped HVO’s evolution into the international center it is today.

The chapters in this volume were written by scientists with strong ties to the USGS, 

including many who spent portions of their careers at HVO studying how Hawaiian 

volcanoes work. “Characteristics of Hawaiian Volcanoes” establishes a benchmark for the 

current understanding of volcanism in Hawai‘i, and the articles herein build upon the elegant 

and pioneering work of Dutton, Jaggar, Stearns, and many other USGS and academic 

scientists. Each chapter synthesizes the lessons learned about a specific aspect of 

volcanism in Hawai‘i, based largely on continuous observation of eruptive activity (like that 

occurring now at Kīlauea Volcano) and on systematic research into volcanic and earthquake 

processes during HVO’s first 100 years. Researchers and students interested in basaltic 

volcanism should find the volume to be a valuable starting point for future investigations 

of Hawaiian volcanoes and an important reference for decades to come, as well as an 

informative and entertaining read.

Suzette M. Kimball

Acting Director

U.S. Geological Survey
v

Preface

“The history of these volcanoes is such as has been supplied by no other volcanic region”

“The records of such a region, whoever the reporter, are of great importance to science”

 —James Dwight Dana, 1890

The Hawaiian Islands have long been recognized as an exceptional natural laboratory for 

volcanology. Indeed, American geologist James Dana emphasized as much with the words quoted 

above from his 1890 treatise on volcanology, “Characteristics of Volcanoes.” Dana visited the 

Hawaiian Islands twice, in 1840 and 1887, and in the interim was kept apprised of the activity 

at Kīlauea and Mauna Loa through letters from local observers, such as Reverend Titus Coan 

(whose interpretations Dana famously, and sometimes erroneously, questioned in the pages of 

the American Journal of Science). In fact, having seen volcanoes around the world (especially 

those in Italy, which were at the time the focus of the volcanological community), Dana states with 

authority that “the two active craters of Hawaii [Kīlauea and Mauna Loa] should share equally 

with Vesuvius and Etna in the attention of investigators.” Much of Dana’s book is dedicated to 

describing Hawaiian volcanoes, and in the preface he provides a surprisingly long list of volcano 

“facts” that were determined from investigations in Hawai‘i rather than at other volcanoes. He also 

laid out a number of “points requiring elucidation,” which are listed in table 1 of the first chapter of 

this volume (by Tilling and others). “The geologist who is capable of investigating these subjects,” 

Dana writes, “will find other inquiries rising as his work goes forward.”

Looking back at the first century of the Hawaiian Volcano Observatory (HVO), it is difficult to 

dispute Dana’s wisdom. Many of the issues he cited have since been thoroughly explored, and the 

answers are now well known. For example, “the temperature of the liquid lava” and “the kinds of 

vapors or gases escaping from the vents or lakes” (two of Dana’s “points”) were some of the first 

problems addressed by Frank A. Perret, Thomas A. Jaggar, and their colleagues when HVO was 

established during 1911–12. Some of Dana’s questions, however, are still just as relevant today 

as they were in 1890. The “differences between the lavas of the five Hawaiian volcanoes,” and 

“the movement of the lavas in the great lava-columns, and the source or sources of the ascensive 

movement” remain compelling issues for modern volcanologists studying Hawai‘i.

In the current volume, “Characteristics of Hawaiian Volcanoes,” we describe the present state of 

the art in understanding how Hawaiian volcanoes work, building on Dana’s initial comprehensive 

examinations of volcanism in Hawai‘i and the studies of many others that have followed. As 

was true in Dana’s time, “much remains to be learned from the further study of the Hawaiian 

volcanoes.” Today, for example, major problems include:

•	 The composition (including volatiles) and depth of the magma source region and its variability 

over space and time

•	 The mechanism for distribution of magma among individual active volcanoes

•	 The connectivity or communication between adjacent volcanoes

•	 The characteristics of magma supply to individual volcanoes over time

•	 The compositional evolution and volume history of individual volcanoes over time, including 

the source and mechanism of rejuvenated-stage volcanism

•	 The inception and initial evolution (both structural and compositional) of Hawaiian volcanoes, 

especially with regard to the still-submerged Lō‘ihi
viii

• The cause of the paired Loa and Kea compositional trends

• The geometries and volumes of the magma plumbing systems at Kīlauea and Mauna Loa

• The mechanisms for volcanic flank instability, and the forecasting of huge landslides

• The causes of great earthquakes and the interactions between tectonic and magmatic activity

• The causes of both eruption onset and eruption termination

• The physics of lava flows, especially with regard to forecasting flow behavior

• The influence of bubble nucleation and coalescence, and of degassing and eventual 

outgassing, on magma convection and eruption style

• The characteristics of three-phase (liquid magma, exsolved gas, and crystals) fluid flow, and 

the manifestations of such flow in geophysical and geochemical time series

• The mechanisms of explosive basaltic eruptions (including the importance of magma-water 

interactions), caldera collapse, and possible cyclicity in explosive and effusive behavior

• The quantification of diverse hazards associated with Hawaiian volcanoes, and the best 

methods for communicating these hazards with the public

Many of these questions are relevant far beyond the confines of the Hawaiian Islands. For 

example, dual compositional trends, like the “Loa Range” and “Kea Range” recognized by Dana, 

are apparent at several Pacific hotspot chains, and the question of how (or if) adjacent, closely 

spaced volcanoes are connected and interact is relevant around the world. As the Hawaiian 

Volcano Observatory enters its second century, continued research making use of Hawai‘i 

as a natural laboratory for volcanology promises insights that will not just contribute to local 

knowledge, but advance the field as a whole.

The 10 chapters that make up this volume treat in detail various aspects of Hawaiian volcanism, 

from the evolution of the volcanoes that make up the island chain to the dynamics of effusive 

and explosive eruptions. This book is not intended to supersede “Volcanism in Hawaii”—the 

magnificent 62-chapter dual-volume USGS Professional Paper 1350, which was produced 

in conjunction with HVO’s 75th anniversary in 1987. Instead, “Characteristics of Hawaiian 

Volcanoes” provides new perspectives on important aspects of basaltic volcanism by synthesizing 

past studies with insights from recent data and modeling, with the ultimate goal of developing 

new models of basaltic volcanism in Hawai‘i and elsewhere. Our hope is that these contributions 

will update the foundation of understanding for Hawaiian volcanism, serving as a starting point for 

researchers as well as providing ideas and stimuli for new avenues of scientific investigation. 

Hawai‘i offers an unparalleled opportunity for students of volcanology to study a vast range of 

problems and processes, and the first 100 years of the Hawaiian Volcano Observatory have 

been an exceptional period of investigation, discovery, and understanding. The outstanding 

eruptive and intrusive activity at multiple volcanoes, long record of volcanism, and relatively 

easy access will ensure that Hawai‘i remains one of the world’s foremost volcano laboratories 

for the next 100 years and beyond. In Dana’s words, “There is terrible sublimity in the quiet work 

of the mighty forces, and also something alluring in the free ticket offered to all comers.”

Michael Poland

July 28, 2014

Hawai‘i Volcanoes National Park, Hawai‘i 
vii

Acknowledgments 

We owe a debt of gratitude to the numerous staff, associates, and volunteers of the Hawaiian 

Volcano Observatory, past and present, without whose efforts and dedication none of the results 

contained in the pages of this volume would have been possible. The authors of the chapters that 

follow would also like to express their gratitude to the many reviewers who helped to evaluate their 

manuscripts, including Steve Brantley, Tom Casadevall, Mike Clynne, Laszlo Kestay, Fred Klein, 

Cynthia Gardner, Lopaka Lee, Peter Lipman, Christina Neal, Matt Patrick, Christina Plattner, Scott 

Rowland, Willie Scott, Dave Sherrod, John Sinton, Tom Sisson, Don Swanson, Wes Thelen, Nathan 

Wood, and Tom Wright. Reference lists for all chapters were compiled, checked, and maintained by 

Taeko Jane Takahashi—an enormous task. Ben Gaddis’s curatorial expertise and vast knowledge of 

the HVO image archive was invaluable. The manuscripts were guided through the USGS approval 

process by Jim Kauahikaua, Tom Murray, Jane Ciener, Keith Kirk, and Mike Diggles. Editorial guidance 

was provided by Dave Sherrod, while Carolyn Donlin, Scott Starratt, Peter Stauffer, and Manny 

Nathenson managed the review, editing, and formatting process. Individual chapters were edited by 

George Havach, Katherine Jacques, Claire Landowski, Sarah Nagorsen, and Peter Stauffer, while the 

design and layout of the book was done by Jeanne DiLeo and Cory Hurd. 

ix

Contents

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

Preface ............................................................................................................................................................v

Acknowledgments .......................................................................................................................................... vii

Chapters

1. The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism

By Robert I. Tilling, James P. Kauahikaua, Steven R. Brantley, and Christina A. Neal .........................................1

2. The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory

By Paul G. Okubo, Jennifer S. Nakata, and Robert Y. Koyanagi.........................................................................67

3. Growth and Degradation of Hawaiian Volcanoes

By David A. Clague and David R. Sherrod ..........................................................................................................97

4. Instability of Hawaiian Volcanoes

By Roger P. Denlinger and Julia K. Morgan.......................................................................................................149

5. Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes

By Michael P. Poland, Asta Miklius, and Emily K. Montgomery-Brown .............................................................179

6. Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes

By Rosalind T. Helz, David A. Clague, Thomas W. Sisson, and Carl R. Thornber............................................237

7. One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory

By A. Jeff Sutton and Tamar Elias......................................................................................................................295

8. The Dynamics of Hawaiian-Style Eruptions: A Century of Study

By Margaret T. Mangan, Katharine V. Cashman, and Donald A. Swanson.......................................................323

9. A Century of Studying Effusive Eruptions in Hawai‘i

By Katharine V. Cashman and Margaret T. Mangan..........................................................................................357

10. Natural Hazards and Risk Reduction in Hawai‘i

By James P. Kauahikaua and Robert I. Tilling ...................................................................................................397
x

The Hawaiian Volcano Observatory Technology Station (left) and Instrument House (center) at the rim of Halema‘uma‘u Crater in 1913. 

USGS photograph, July 6, 1913 (photographer unknown).
Chapter 1

The Hawaiian Volcano Observatory—A Natural Laboratory 

for Studying Basaltic Volcanism

By Robert I. Tilling1, James P. Kauahikaua1, Steven R. Brantley1, and Christina A. Neal1

1

U.S. Geological Survey.

Abstract

In the beginning of the 20th century, geologist Thomas 

A. Jaggar, Jr., argued that, to fully understand volcanic and 

associated hazards, the expeditionary mode of studying 

eruptions only after they occurred was inadequate. Instead, 

he fervently advocated the use of permanent observatories 

to record and measure volcanic phenomena—at and below 

the surface—before, during, and after eruptions to obtain 

the basic scientific information needed to protect people 

and property from volcanic hazards. With the crucial 

early help of American volcanologist Frank Alvord Perret 

and the Hawaiian business community, the Hawaiian 

Volcano Observatory (HVO) was established in 1912, and 

Jaggar’s vision became reality. From its inception, HVO’s 

mission has centered on several goals: (1) measuring and 

documenting the seismic, eruptive, and geodetic processes 

of active Hawaiian volcanoes (principally Kīlauea and 

Mauna Loa); (2) geological mapping and dating of deposits 

to reconstruct volcanic histories, understand island 

evolution, and determine eruptive frequencies and volcanic 

hazards; (3) systematically collecting eruptive products, 

including gases, for laboratory analysis; and (4) widely 

disseminating observatory-acquired data and analysis, 

reports, and hazard warnings to the global scientific 

community, emergency-management authorities, news 

media, and the public. The long-term focus on these goals 

by HVO scientists, in collaboration with investigators from 

many other organizations, continues to fulfill Jaggar’s 

career-long vision of reducing risks from volcanic and 

earthquake hazards across the globe. 

A volcano observatory must see or measure the 

whole volcano inside and out with all of science 

to help.

 —Thomas A. Jaggar, Jr. (1941)

This chapter summarizes HVO’s history and some of 

the scientific achievements made possible by this permanent 

observatory over the past century as it grew from a small 

wooden structure with only a small staff and few instruments 

to a modern, well-staffed, world-class facility with state-ofthe-art monitoring networks that constantly track volcanic 

and earthquake activity. The many successes of HVO, from 

improving basic knowledge about basaltic volcanism to 

providing hands-on experience and training for hundreds of 

scientists and students and serving as the testing ground for 

new instruments and technologies, stem directly from the 

acquisition, integration, and analysis of multiple datasets 

that span many decades of observations of frequent eruptive 

activity. HVO’s history of the compilation, interpretation, 

and communication of long-term volcano monitoring and 

eruption data (for instance, seismic, geodetic, and petrologicgeochemical data and detailed eruption chronologies) is 

perhaps unparalleled in the world community of volcano 

observatories. The discussion and conclusions drawn in 

this chapter, which emphasize developments since the 75th 

anniversary of HVO in 1987, are general and retrospective 

and are intended to provide context for the more detailed, 

topically focused chapters of this volume.

Introduction

The eruption of Vesuvius in 79 C.E. prompted the first 

scientific expedition (by Pliny the Elder) to study volcanic 

phenomena, as well as the first written eyewitness account 

(by Pliny the Younger) of eruptive activity (Sigurdsson, 

2000). The new science of geology emerged in the 19th 

century, focusing on the deduction of past events from 

current Earth exposures—“the present is the key to the past.” 

This approach was also used for studying active geologic 

processes like volcanic eruptions: scientific studies of 

volcanoes were conducted during short-lived expeditions, 

generally undertaken in response to major eruptions (for 

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014
2 Characteristics of Hawaiian Volcanoes

instance, the 1815 Tambora and 1883 Krakatau eruptions in 

Indonesia) and done substantially after the event was over. 

Three large explosive eruptions in the Caribbean-Central 

American region in 1902—La Soufrière (Saint Vincent, 

West Indies), Montagne Pelée (Martinique, West Indies), and 

Santa María (Guatemala)—claimed more than 36,000 lives 

and showed the inadequacy of the deductive approach for 

protecting people and property from natural disasters. These 

catastrophic eruptions in the early 20th century set the stage 

for the emergence of the science of volcanology as we know 

it today. 

Dr. Thomas Augustus Jaggar, Jr., a 31-year-old 

geology instructor at Harvard University (and also a parttime employee of the U.S. Geological Survey [USGS]), 

was a member of the scientific expedition sent by the 

U.S. Government to investigate the volcanic disasters at 

La Soufrière and Montagne Pelée in 1902. High-speed, 

incandescent pyroclastic flows (nuées ardentes) from 

Montagne Pelée obliterated the city of St. Pierre and killed 

29,000 people in minutes, making it the deadliest eruption 

of the 20th century (Tanguy and others, 1998, table 1). The 

Pelée eruption’s power and deadly impacts left a profound 

impression on the young professor, and he decided to devote 

his career to studying active volcanoes (Apple, 1987). A 

half-century later, Jaggar reflected in his autobiography: 

“As I look back on the Martinique expedition, I know 

what a crucial point in my life it was. . . . I realized that the 

killing of thousands of persons by subterranean machinery 

totally unknown to geologists and then unexplainable was 

worthy of a life work” (Jaggar, 1956, p. 62). In reaching his 

life-changing decision, Jaggar was swayed by his strong 

conviction that the expeditionary approach in studying 

volcanoes was inadequate. Instead, he firmly believed that, 

to understand volcanoes fully and to mitigate effectively the 

impacts of their hazards, it is necessary to study and observe 

them continuously—before, during, and after eruptions. After 

meeting the renowned American volcanologist Frank Alvord 

Perret, who was already using this approach at Vesuvius in 

1906, Jaggar became even more convinced about advocating 

for the establishment of permanent Earth observatories. While 

at Harvard, and later as a professor at the Massachusetts 

Institute of Technology (MIT), Jaggar pursued his life’s goal 

to establish a permanent observatory at some place in the 

world to study volcanoes and earthquakes (see “Founding of 

the Hawaiian Volcano Observatory” section, below).

Scope and Purpose of This Chapter

In 1987, to commemorate the 75th anniversary of the 

founding of HVO, the U.S. Geological Survey published 

Professional Paper 1350 (Decker and others, 1987). This 

two-volume work still stands as the most comprehensive 

compilation of the many studies on Hawaiian volcanism 

by USGS and other scientists through the mid-1980s. The 

62 papers contained in Professional Paper 1350 (and the 

references cited therein) provide an invaluable database for 

understanding Hawaiian volcanism. It is beyond the scope 

of this volume to synthesize fully the abundance of scientific 

data, new interpretations, and insights that have accrued 

in the quarter century since that publication. Instead, the 

papers in this current volume are retrospective and focus 

on volcano monitoring and selected topical studies that 

refine and extend our ideas about how Hawaiian volcanoes 

work. Efforts summarized here have been led primarily by 

HVO researchers and other USGS scientists but were often 

completed in close collaboration with non-USGS colleagues 

from government, academic, and international institutions. 

Much of HVO’s contribution to science since 1987 has been 

the direct result of monitoring the continuing eruption of 

Kīlauea Volcano.

Specifically, this introductory chapter provides 

historical context for the subsequent chapters, which 

encompass these themes: the key role of permanent seismic 

and other geophysical networks in volcano monitoring 

(Okubo and others, chap. 2); evolution of oceanic shield 

volcanoes (Clague and Sherrod, chap. 3); flank stability 

of Hawaiian volcanoes (Denlinger and Morgan, chap. 4); 

magma supply, storage, and transport processes (Poland and 

others, chap. 5); petrologic insights into basaltic volcanism 

(Helz and others, chap. 6); chemistry of volatiles and gas 

emissions (Sutton and Elias, chap. 7); dynamics of Hawaiian 

eruptions (Mangan and others, chap. 8); effusive basaltic 

eruptions (Cashman and Mangan, chap. 9); and natural 

hazards associated with island volcanoes (Kauahikaua and 

Tilling, chap. 10). Interpretations provided by these topical 

studies are derived from, and constrained by, long-term 

data—visual, geophysical, and petrologic-geochemical—on 

the eruptive processes and products of Kīlauea and Mauna 

Loa obtained by HVO over many decades. The papers in 

this volume, we believe, reflect current HVO science and 

reinforce Jaggar’s vision that reduction of volcano risk 

requires the integration of systematic monitoring data 

and related research, comprehensive hazards assessments 

based on past and current eruptive activity, and effective 

communication of hazards information to authorities and 

the potentially affected populace. Thanks to the progress in 

the past 100 years, we now have many more scientific tools 

than were available in the early 20th century to improve our 

understanding of volcanic phenomena. Clearly, the legacy of 

Thomas Jaggar is alive and well.

The history of HVO is, in many ways, the history of 

basaltic volcanology. It is almost impossible to separate 

contributions by HVO and USGS scientists and student 

volunteers from those of our partners in academia and other 

institutions, but we chose to focus on the big ideas that came 

from work on Kīlauea that predominantly involved scientists 

and students from HVO and other USGS groups. The future of 

systematic scientific studies of Hawaiian volcanoes relies now, 

as during the past century, on continued government-academic 

and scientist-student partnerships.
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 3

Founding of the Hawaiian Volcano 

Observatory

After his work at Montagne Pelée in the Caribbean and 

Vesuvius in Italy, Thomas Jaggar led a scientific expedition, 

funded by Boston businessmen, to various volcanoes in the 

Aleutian Islands of Alaska in 1907. There, he witnessed the 

reactivation of Bogoslof volcano rising out of the sea but 

bemoaned the loss of data on the continuing eruption after the 

expedition had returned to the United States:

The remarkable processes of volcanism and earth 

movement in the Aleutian Islands deserve continuous, close study from an observatory erected for the 

purpose on Unalaska. The winter of 1907–8 has been 

wasted—lost to science, because no observers were 

stationed there. (Jaggar, 1908, p. 400).

The Messina earthquake, later in 1908, added to the human 

toll from natural disasters that Jaggar summarized as “100 

persons a day since January 1, 1901” (Jaggar, 1909). In his 1909 

publication about the earthquake, Jaggar put forth his master plan 

for 10 small observatories in “New York, Porto Rico [sic], Canal 

Zone, San Francisco, Alaska, Aleutian Islands, Philippines, Hawaii, 

Scotland, and Sicily.” The overall cost would be a $4.2 million 

endowment that would continue support for each observatory with 

$10,000 per year (Jaggar, 1909). The goals of these observatories 

were simple: (1) prediction of earthquakes, (2) prediction of 

volcanic eruptions, and (3) earthquake-proof engineering and 

construction in volcanic and seismic lands (Jaggar, 1909).

He also expressed deep admiration for the efforts of the 

Japanese in establishing geophysical monitoring (“. . . their 

island empire is girdled with observatories”) and spent 5 

weeks in Japan in 1909 with layovers in Honolulu both ways 

(Jaggar, 1910). During his return layover, Jaggar spoke to the 

Honolulu Chamber of Commerce about the unique possibilities 

for science afforded by the establishment of an observatory 

at the edge of Kīlauea Volcano. Lorrin Thurston, a wellconnected businessman and political figure, also spoke to the 

Chamber about the “purely commercial advantages of securing 

for Kilauea such an observatory. . . . From a purely business 

point of view it would pay Hawaii to subscribe the funds 

necessary for the maintenance of the observatory, irrespective 

of the great scientific benefit to accrue.” Jaggar asked for a 

commitment of $5,000 per year to locate an observatory at 

Kīlauea but received a promise of only half that amount before 

he left for Boston (Hawaiian Gazette, 1909). With this partial 

encouragement, Jaggar redoubled his efforts back in Boston 

to seek financial supporters—in New England as well as in 

Hawai‘i—to build the observatory, including the facilities to 

house instruments and records, a laboratory, and offices.

Jaggar and his associates were able to raise funds during 

1909–11 in Boston to purchase seismometers and temperaturemeasuring instruments, to support initial field studies at 

Kīlauea, and to construct a temporary small frame building 

(the “Technology Station”) on the rim of Halema‘uma‘u 

Crater (fig. 1A) for observations of the continuous lava-lake 

activity. Neither Jaggar nor any other MIT scientists were 

able to travel to Hawai‘i during 1910–11, however, and so 

the earliest observations and studies at Kīlauea fell to E.S. 

Shepherd (Geophysical Laboratory of the Carnegie Institution of 

Washington, D.C.) and Frank A. Perret (Apple, 1987). Doubtless, 

Jaggar worried that his delay would be during a critical time 

in the nascent observatory; thus, he wisely enlisted Perret—a 

prominent, volcano-savvy scientist already well known for 

his work at Vesuvius, Etna, and Stromboli volcanoes—to be 

his proxy. Jaggar considered Perret to be “the world’s greatest 

volcanologist” (Jaggar, 1956, p. xi).

A B

Figure 1. Photographs showing facilities of the Hawaiian Volcano Observatory (HVO) through the years. A, The “Technology Station” (circled) 

on the eastern rim of Halema‘uma‘u Crater, built by Frank A. Perret in 1911, was the first, though temporary, of a number of buildings that HVO 

has occupied since its founding (USGS photograph by Frank A. Perret). B, Aerial view of present-day HVO and Jaggar Museum (lower left corner) 

on the northwestern rim of the summit caldera of Kīlauea Volcano, with plume rising from vent in Halema‘uma‘u Crater. This vent opened in midMarch 2008 and has remained active through mid-2014 (USGS photograph taken in September 2008 by Michael P. Poland).
4 Characteristics of Hawaiian Volcanoes

Under Jaggar’s direction, Perret built the Technology 

Station and conducted an experiment to measure the 

temperature of Kīlauea’s active lava lake (see “Field 

Measurements of Lava Temperature” section, below). He also 

started nearly continuous observations and measurements of 

the lava lake then within Halema‘uma‘u Crater. At Thurston’s 

urging, he wrote weekly updates in The Pacific Commercial

Advertiser (later The Honolulu Advertiser), which happened to 

be owned by Thurston. In those updates, which began HVO’s 

long tradition of regular scientific communications and public 

outreach (see “Communication of Scientific Information and 

Public Outreach” section, below), Perret listed himself as 

“Director pro tem” of the Technology Station, the first building 

of the not yet formally established observatory. In the summer 

of 1911, Perret’s work, as summarized weekly in Thurston’s 

newspaper, greatly impressed and excited Thurston’s group of 

Honolulu financial backers, prompting the group—formally 

organized on October 5, 1911, as the “Hawaiian Volcano 

Research Association” (HVRA)—to renew their pledge of 

financial support for the permanent observatory at Kīlauea. 

Perret’s achievements thus provided a solid financial, as well 

as scientific, base upon which Jaggar soon built the formal 

observatory. HVRA’s funds, however, did not directly support 

HVO’s effort until mid-1912, when the 5-year contract 

between HVRA and MIT became official (Dvorak, 2011).

In January 1912, Jaggar (fig. 2) arrived to resume the 

continuous observations begun by Perret and to start erecting an 

observatory building with financial and material donations from 

Hilo businesses. The year 1912 has long been recognized as when 

HVO was founded. Though there was no formal ceremony or 

event to mark its “official” establishment, 1912 has long been 

recognized as the year when HVO was founded. In any case, the 

founding date must be some time between July 2, 1911, when 

Perret arrived to begin continuous observations, and July 1, 

1912, when Jaggar received his first paycheck as HVO Director 

from the HVRA (Dvorak, 2011; Hawaiian Volcano Observatory 

Staff, 2011). By mid-February 1912, construction was begun on 

what was to become the first of several permanent facilities of 

HVO, located near the present-day Volcano House Hotel on the 

northeastern rim of Kīlauea Caldera (Apple, 1987).

Why in Hawai‘i?

To fully appreciate Jaggar’s accomplishment in 

establishing HVO, we first must look back in time to the early 

20th century. Before 1912, only three volcano observatories 

existed in the world: (1) the Vesuvius Observatory (Reale 

Osservatorio Vesuviano [now a museum] on the flank of 

Vesuvius volcano in Italy), whose construction was completed 

in 1848; (2) the fledgling observatory established in mid1903 on the island of Martinique by the French in response 

to continuing eruptive activity at Montagne Pelée; and (3) 

the Asama Volcano Observatory, Japan, established in 1911 

(Suwa, 1980) by the renowned seismologist Fusakichi Omori, 

who later became a close colleague and friend of Jaggar. 

Given Hawai‘i’s geographic isolation in the middle of 

the Pacific Ocean, what prompted Jaggar to build a volcano 

observatory at Kīlauea rather than pursuing his scientific 

studies at one of the three existing observatories? Jaggar 

(1912, p. 2) had a number of compelling scientific, as well as 

practical, reasons, including these: (1) Kīlaueaʼs location is in 

“American” territory rather than in a foreign land, “and these 

volcanoes are famous . . . for their remarkably liquid lavas 

and nearly continuous activity”; (2) at other volcanoes the 

eruptions are more explosive and an observatory located close 

Figure 2. Photograph showing 

Thomas A. Jaggar, Jr., founder of 

the Hawaiian Volcano Observatory 

(HVO), tending to seismometers in 

1913 in the Whitney Laboratory of 

Seismology (photograph courtesy 

of the Bishop Museum, Honolulu). 

Inset (upper left corner) shows 

portrait photograph of Jaggar in 

1916 (USGS/HVO photograph).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 5

enough to the center of activity is in some danger. Kīlauea, 

while displaying great and varied activity, is relatively safe; (3) 

earthquakes are frequent and easily studied; and (4) “Kilauea 

is very accessible,” only 50 km by road from Hilo harbor, 

which, in turn, is only a one-day sail from Honolulu (the most 

developed city in the Territory of Hawaii). Jaggar obviously 

believed that these reasons more than compensated for the 

disadvantages of Hawai‘i’s geographic remoteness.

For Jaggar, another important motivation for locating 

an observatory in Hawai‘i was the powerful support from 

Thurston, who was well connected financially and was keen to 

promote Kīlauea as a tourist attraction. At the time, the volcano 

was already an emerging tourist destination, and Thurston 

was the major stockholder of the Volcano House Hotel on the 

rim of Kīlauea Caldera and the owner of the train from Hilo 

to the volcano. Thurston’s role as a tourism promoter is aptly 

described by Dvorak (2011, p. 35):

His grand plan was to make Kilauea one of the 

scheduled stops for the increasingly numerous 

passenger ships crossing the Pacific. If someone 

wanted to see the lava lake, that person had to ride 

his train; if anyone wanted to stay overnight, his hotel 

would provide the only available accommodations.

Indeed, according to Allen (2004, p. 196), “Lorrin A. 

Thurston created the foundation for Hawaii tourism.” 

Conceivably, funds raised by the HVRA to support HVO were 

given more for the promise of tourism enhancement, rather than 

for stated reasons of scientific advancement or reduction of risk. 

This notion, while not explicitly documented, may explain the 

final phrase of one of Jaggar’s major goals for HVO (Jaggar, 

1913, p. 4): “Keep and publish careful records, invite the whole 

world of science to co-operate, and interest the business man” 

[italics in original]. A skillful promoter himself, Jaggar thus 

successfully merged his scientific interests with the more 

commercial interests of the Honolulu businessmen.

Reflecting on the founding of HVO decades later, Jaggar 

comments in his book Volcanoes Declare War (1945, p. 148): 

 It is appropriate that Honolulu should have been 

the American community to establish first a volcano 

observatory. We are in the midst of the greatest 

ocean, surrounded by earthquakes and volcano lands. 

The place is like a central fire station, and there is 

some appeal to the imagination in possessing a world 

fire alarm center.

After alluding to several “disasters” (volcanic and 

earthquake) in the circum-Pacific regions, he states that these 

disasters “and a hundred others constitute an endless warfare, and 

what more fitting center for mobilization against it than the natural 

laboratory of the Island of Hawaii?” Jaggar clearly envisioned an 

observatory as needing to employ multiple scientific approaches 

and all available and emerging technologies in its studies. With 

its frequent eruptions, earthquakes, and tsunamis, the Island 

of Hawai‘i was the perfect locale for conducting continuous 

scientific observation to more fully understand eruptive and 

seismic phenomena and their associated hazards. HVO’s studies 

of earthquake, volcanic, and tsunami hazards and associated 

mitigation strategies, with some case histories, are treated in the 

chapter by Kauahikaua and Tilling (this volume, chap. 10).

Post-Jaggar History of HVO

Jaggar served as HVO director through periods when the 

observatory was administered by the Weather Bureau (1919–24), 

the U.S. Geological Survey (1924–35), and the National Park 

Service until his retirement in 1940. In 1947, administration 

of HVO permanently returned to the U.S. Geological Survey 

and, in 1948, the HVO operation was relocated to its present 

site at Uēkahuna Bluff on the caldera’s northwestern rim 

(fig.1B); its facilities were gradually expanded with addition 

of a geochemistry wing in the early 1960s and the construction 

in 1985–86 of a much larger adjoining building with a viewing 

tower. For detailed accounts of the pivotal roles played by 

Jaggar, Perret, Thurston, and the HVRA in the founding of HVO 

and of the observatory’s early history, see Macdonald (1953a), 

Apple (1987), Barnard (1991), Hawaiian Volcano Observatory 

Staff (2011), Dvorak (2011), Kauahikaua and Poland (2012), and 

Babb and others (2011). 

Developing, Testing, and Using VolcanoMonitoring Techniques

Since the growth and spread of volcano surveillance 

throughout the world, experience clearly has shown that 

seismic and geodetic monitoring techniques are the most 

diagnostic and useful tools to detect eruption precursors 

(Tilling, 1995; Scarpa and Tilling, 1996; McNutt, 2000; 

McNutt and others, 2000; Dzurisin, 2007; Segall, 2010). 

By the early 20th century, the common association between 

premonitory seismicity and ground deformation had been 

documented (for example, Omori, 1913, 1914; Wood, 1913, 

1915). As a specific example, after summarizing seismic and 

ground-tilt behavior at Kīlauea, The Volcano Letter of August 

14, 1930, states (Powers, 1930, p. 3), 

The conclusion drawn from all this evidence is that 

lava pressure is increasing under Halemaumau. It is 

impossible to say whether or not this will result 

in an eruption, [bold in original] but it can be said 

confidently that conditions look more favorable now 

than at any time in the past several months.

On November 11, 1930, an 18-day eruption began in 

Halema‘uma‘u.

Throughout the 20th century, HVO has served as a 

developing and testing ground for volcano-monitoring instruments 

and techniques, many of which have been further adapted for 

use at other volcanoes worldwide. In recent decades, advanced 

satellite-based volcano-monitoring techniques—Global Positioning 

System (GPS), interferometric synthetic aperture radar (InSAR), 
6 Characteristics of Hawaiian Volcanoes

and thermal imaging—have been successfully applied at Kīlauea 

and Mauna Loa, providing much more complete spatial and 

temporal time series of deformation patterns and lava-flow 

inundation than ever before. The data collected by HVO’s longterm ground-based monitoring program, however, have proved 

to be invaluable for checking and validating the results obtained 

from space-age monitoring techniques. In summarizing HVO 

deformation studies and techniques employed during 1913–2006, 

Decker and others (2008, p. 1–2) emphasized that “Many of 

the techniques are complementary; for example, using GPS and 

satellite measurements of benchmark positions provides ‘ground 

truth’ for InSAR (satellite radar interferometry) maps.” Below, we 

offer some examples of developments in instrumental and volcanomonitoring techniques during HVO’s history.

Seismic Monitoring

The use of seismic waves to detect unrest at volcanoes 

began in the mid-19th century. A Palmieri (electromagnetic) 

seismograph at the Vesuvius Observatory detected precursory 

seismic activity before the 1861, 1868, and 1872 eruptions at 

Mount Vesuvius (Giudicepietro and others, 2010). The first 

seismometer in Hawai‘i was installed on O‘ahu in 1899 (Klein 

and Wright, 2000), and instrumental recording of earthquakes on 

the Island of Hawai‘i was initiated with the completion in 1912 

of the Whitney Laboratory of Seismology (fig. 2)—a basement 

vault beneath HVO’s main building. The first seismometers 

at HVO were two instruments imported from Japan (shipped 

directly to Hawai‘i) and one from Germany (shipped from 

Strasburg via Boston); some of these were modified later to 

better record volcanic seismicity at Kīlauea. The data from the 

seismographs that were collected in the HVO vault were flawed 

for a variety of reasons (for instance, building vibrations, nearby 

cultural noise, wide fluctuations in vault temperature) but still 

provided useful information (Apple, 1987; Klein and Wright, 

2000). For example, these first instruments were sufficient in 

establishing that Hawaiian eruptions were preceded by increased 

seismicity and ground tilt changes (Wood, 1915).

During the early decades of seismic monitoring, HVO 

never operated more than five stations (two at Kīlauea’s summit, 

one at ~3,300 m elevation on the eastern flank of Mauna Loa, 

and two outlying ones in Kealakekua and Hilo). Moreover, these 

early instruments were heavy and unwieldy, had low sensitivity, 

and lacked the capability to transmit data to the observatory. 

It was not possible to determine accurate earthquake locations 

because of the inadequate density of seismometers, imprecise 

timing mechanisms, and lack of direct data transmission. 

Nonetheless, the early seismic recordings generally sufficed 

to estimate relative intensity and distance to origin and to 

discriminate whether an earthquake was associated with Kīlauea, 

Mauna Loa, or Hualālai or was a teleseism (Apple, 1987; 

Wright, 1989; Wright and others, 1992a). 

Seismic monitoring at HVO was upgraded substantially 

with the arrival in 1953 of seismologist Jerry P. Eaton, who 

introduced the smaller, more sensitive, electromagnetic 

seismometer and established the first telemetered seismic network 

at Kīlauea within a few years (Wright, 1989; Klein and Wright, 

2000; Okubo and others, this volume, chap.2). Signals from six 

seismometers were transmitted to the observatory via overland 

cables and recorded on smoke-drum seismographs. Data from 

this rudimentary network, combined with a crustal-velocity 

structure model also developed by Eaton, made possible routine 

determination of earthquake locations and magnitudes; Eaton’s 

modernization of the network enabled HVO to produce catalogs 

of reliably located earthquakes by the 1960s (Wright, 1989). 

Equally important, Eaton’s seismic network in Hawai‘i served as 

a prototype upon which a number of “modern” networks in other 

regions (for example, California) were based.

With expanded coverage, more sensitive instruments, and 

a more accurate seismic velocity model, the quality of the data 

catalog improved. By 1974, the HVO seismic network had 

grown to 34 seismic stations, and by 1979 all seismic data were 

processed by computer (Klein and others, 1987). The availability 

of high-quality data from the modern seismic network made it 

possible to extend the catalog of Hawaiian earthquakes back in 

time by estimating locations and magnitudes of reported historical 

events (Wyss and Koyanagi, 1992; Klein and Wright, 2000). The 

first digital seismometers were installed at Kīlauea as part of a 

joint United States-Japan seismic experiment in 1996 (McNutt 

and others, 1997); about 10 broadband seismometers remained 

operational at Kīlauea summit after the experiment but were 

not used in routine processing until 2007, when HVO’s data 

acquisition software was upgraded from Caltech-USGS Seismic 

Processing (CUSP) to Earthworm (Okubo and others, this volume, 

chap. 2). The American Recovery and Reinvestment Act funding 

of 2009 allowed HVO to fully upgrade its seismic network with 

more broadband seismometers and digital telemetry. As of this 

writing (mid-2014), HVO’s seismic network (fig. 3A) is among the 

densest volcano-monitoring networks in the world, consisting of 

57 stations over the five volcanoes of the island, 21 of which use 

broadband digital instruments.

HVO’s seismic-monitoring data constitute an integral 

component in chronological narratives and interpretations 

of all Hawaiian eruptions since the first instrument became 

operational. Okubo and others (this volume, chap. 2) review, 

in detail, the evolution of HVO’s seismic monitoring 

systems with time, highlighting the significant findings from 

progressively improving data that sharpen our understanding 

of how Hawaiian and other basaltic volcanoes work. 

Geodetic Monitoring

It is now well demonstrated that the surfaces of active 

volcanoes deform in response to inflation or deflation of 

subsurface magma reservoirs and hydrothermal systems (see, 

for example, Murray and others, 2000; Dzurisin, 2007). This 

phenomenon had been recognized but was poorly understood 

in the early 20th century; however, from its beginning in 

1912, HVO used geodetic measurements to track ground 

deformation. Decker and others (2008) provide a detailed 

account of the methodologies—including some developed, 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 7

156° 155°30' 155°

19°

19°30'

20°

19°

19°30'

20°

Mauna Kea

Kohala

Mauna

Loa

Hualālai

Kīlauea

HVO

156° 155°30' 155°

Tilt/strain network

Strainmeter

Electronic tiltmeter

Tilt survey array

Level lines

20 KILOMETERS

0 10 MILES

0

Seismic station

Mauna Kea

Kohala

Mauna

Loa

Hualālai

Kīlauea

Mauna Kea

Kohala

Mauna

Loa

Hualālai

Kīlauea

Continuously recording site

Survey benchmark

GPS network

Mauna Kea

Kohala

Mauna

Loa

Hualālai

Kīlauea

A B

C D

E

adapted, or refined by HVO—that have been employed for 

deformation studies on active Hawaiian volcanoes. Drawing 

from this summary, we comment below on the historical 

importance of some of the early techniques and measurements 

and then consider satellite-based geodesy.

Tilt

The earliest tilt measurements were made after the 

discovery that the seismographs in the basement of HVO’s 

first building (the “Whitney Vault”) were affected by 

deflection (relative to the vault floor) of the horizontal 

pendulums, apparently in response to deformation of 

Kīlauea’s summit. The deflection-induced offsets on the 

seismograms could be related to summit tilt (Apple, 1987). 

“Thus, the Hawaiian Volcano Observatory . . . inadvertently 

began to record tilt continuously . . . from 1913 to 1963” 

(Decker and others, 2008, p. 8). While crude, these 

inadvertent seismometric measurements well recorded the 

large tilt changes related to the 1924 explosive eruptions. 

Beginning in the 1950s, the quality of tilt measurements 

improved greatly with use of permanent and portable watertube tiltmeters (Eaton, 1959) and by the installation in 1966 

of a continuously recording mercury-capacitance tiltmeter 

in the basement of HVO’s facilities on Uēkahuna Bluff (the 

“Uēkahuna Vault”; Decker and others, 2008).

Additional continuously recording electronic tiltmeters, 

including electronic borehole tiltmeters, were later installed 

at Kīlauea and Mauna Loa (fig. 3C). Four analog borehole 

tiltmeters operating along the Kīlauea East Rift Zone 

(ERZ) documented dike propagation and the onset of the 

Pu‘u‘Ō‘ō eruption in January 1983 (Okamura and others, 

1988). The borehole tilt networks were expanded throughout 

the 1990s and into the 21st century, and during 2010–11, 

several advanced digital borehole tiltmeters were installed 

on Kīlauea and Mauna Loa (their broad frequency response 

allows them to record low-frequency seismic tremor and 

Figure 3. Maps and photograph showing 

location of selected Hawaiian Volcano 

Observatory (HVO) geophysical monitoring 

stations and level lines on the Island of Hawai‘i as 

of 2012. A, Seismic stations. B, Borehole tilt and 

scalar strain measurement stations. C, Global 

Positioning System (GPS) network. D, Level 

lines. E, An HVO technician servicing one of the 

monitoring stations (ALEP) on Mauna Loa, at 

which a seismometer and continuously recording 

GPS are co-located (USGS photograph by Kevan 

Kamibayashi). Parts A–D are modified from Tilling 

and others (2010). Not all stations provide realtime data to HVO; some are campaign sites (for 

example, blue circles in B and C).
8 Characteristics of Hawaiian Volcanoes

teleseisms and blurs the line between seismic and geodetic 

monitoring). Tilt measurements at Kīlauea’s summit since 

1912 (fig. 4) constitute the world’s longest duration and most 

comprehensive time-series dataset for such measurements. 

Electronic Distance Measurement (EDM)

Among the world’s volcano observatories, HVO was a 

pioneer in using theodolites, or electronic distance measurement 

(EDM), to routinely measure horizontal displacements at 

deforming volcanoes, beginning in 1964 (Decker and others, 1966). 

A major advantage of EDM over traditional triangulation is the 

relative ease in measuring three sides of a triangle (“trilateration”) 

to yield precise determination of horizontal displacement vectors. 

Trilateration surveys in the 1970s and 1980s were HVO’s 

mainstays for tracking horizontal distance changes related to 

eruptions, intrusions, and earthquakes (Decker and others, 1987, 

2008). HVO’s network of EDM benchmarks, later also used for 

GPS monitoring (fig. 3D), grew significantly through the early 

1990s. EDM data also conclusively showed that Kīlauea’s south 

flank was moving seaward several centimeters per year (see “Flank 

Instability” section, below). The EDM technique is now used 

only for training purposes, to give students and scientists, mostly 

from developing countries, background about ground-deformation 

monitoring (see discussion in “Cooperative Research and Work 

with Other Organizations” section, below).

Satellite-Based Geodesy

During the past quarter century, satellite-based techniques 

(space geodesy) have increasingly been used to measure 

ground deformation related to a wide variety of dynamic 

earth processes, including fault movement/rupture and strain 

accumulation and release at volcanoes. To date, the two 

techniques most widely used to detect and image deformation 

at active volcanoes are the Global Positioning System (GPS) 

and interferometric synthetic aperture radar (InSAR). (For 

good summaries of the principles and applications of these 

techniques, see Dzurisin, 2007; Lu and Dzurisin, 2014.)

Because repeat GPS measurements can yield both vertical 

and horizontal displacements, the GPS technique quickly 

became the geodetic-monitoring tool of choice at Hawaiian 

and other volcanoes. Beginning in 1996, in cooperation 

with investigators at the University of Hawai‘i, Stanford 

University, and other institutions, HVO established sites for 

continuous GPS measurement on Kīlauea, Mauna Loa, and 

Mauna Kea volcanoes. At present, HVO’s continuous GPS 

monitoring network consists of 60 receivers (fig. 3D). In the 

21st century, the combination of continuous and campaign GPS 

measurement, together with conventional geodetic methods, 

has provided greater time resolution for geodetic changes at 

Hawaiian volcanoes unattainable in the previous century. The 

comprehensive geodetic data now available make possible more 

1920 1940 1960 1980 2000 -400

-300

-200

-100

0

100

200

300

400

500

Tilt, in microradians

Whitney N-S

Whitney E-W

Uēkahuna N-S

Uēkahuna E-W

1930 1950 1970 1990 2010

Year

Figure 4. Graph showing fluctuations in summit tilt at Kīlauea Caldera during the period 1913–2011, as measured by the seismometric method 

in the Whitney Vault and by water-tube tiltmeters at Uēkahuna (see text for discussion). Tilt readings in both north-south and east-west directions 

are shown. The seismometric tilts have been converted to microradians (~0.00006 degree), the conventional measurement unit for tilt change. The 

seismometric record is offset from the water-tube tiltmeter record because they pertain to different geographical sites at Kīlauea’s summit.
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 9

detailed and better constrained 

models of summit and rift zone 

magma reservoir, transport, and 

eruption dynamics at Kīlauea and 

Mauna Loa (see, for example, 

Cervelli and Miklius, 2003; Miklius 

and others, 2005; Poland and others, 

2012; Wright and Klein, 2014; 

Poland and others, this volume, 

chap. 5). Moreover, continuous GPS 

measurements (fig. 5A) have made 

possible real-time tracking of ground 

deformation associated with magma 

movement, eruption dynamics, and 

the motion of Kīlauea’s unstable 

south flank (see “How Hawaiian 

Volcanoes Work” section, below). 

The potential of InSAR 

in volcano monitoring was 

first demonstrated by imaging 

the 1992–93 deflation at Etna 

Volcano, Italy (Massonnet and 

others, 1995). This technique 

is especially powerful in that it 

captures deformation of the entire 

radar-imaged ground area, rather 

than change at individual points, 

as measured by other monitoring 

techniques (for instance, GPS, 

EDM, tilt, and leveling). InSAR 

mapping was first tested at Kīlauea 

in 1994, but the interferograms 

contained large errors because 

of atmospheric effects related to 

Hawai‘i’s tropical environment, 

resulting in ambiguous interpretation 

(Rosen and others, 1996). With 

time, however, the InSAR technique 

improved as atmospheric artifacts 

were more easily recognized and 

new techniques developed to 

mitigate such artifacts. InSAR is 

now a versatile tool routinely used 

for mapping volcano deformation 

at Kīlauea (fig. 5B) and Mauna Loa 

(for example, Amelung and others, 

2007) and at volcanoes around the 

world (for example, Dzurisin and 

Lu, 2007, and examples summarized 

therein). The development in 2007 

of airborne InSAR (a radar pod 

attached to fixed-wing aircraft) 

eliminated constraints of orbit paths 

and satellite repeat passage, thereby 

providing much greater flexibility 

in the acquisition of data. Airborne 

0 2.83

Range change, in cm

0

0

2 4 KILOMETERS

1 2 MILES

B

155°20' 155°15' 155°10' 155°05' 155° 154°55'

19°10'

19°15'

19°20'

19°25'

19°30'

0

0

5 10 KILOMETERS

2.5 5 MILES

10 5 0 cm

A

East Rift Zone

Kīlauea

Caldera

Kīlauea

Caldera

East

Makaopuhi

Crater

Rift

Zone

Pu‘u ‘Ō‘ō

Figure 5. Map and InSAR (interferometric synthetic aperature radar) image showing horizontal 

and vertical ground deformation at Kīlauea Volcano. A, An example of geodetic monitoring results 

using continuously recording Global Positioning System (GPS) receivers of the Hawaiian Volcano 

Observatory (HVO), Stanford University, and the University of Hawai‘i. Vectors indicate the horizontal 

displacements produced by magma intrusion into Kīlauea’s upper East Rift Zone and associated brief 

eruption during June 17–19, 2007. Station locations are at the tails of vectors, and circles indicate 

95-percent confidence levels (modified by Asta Miklius from Montgomery-Brown and others, 2010, 

figure 3). B, An InSAR interferogram, derived from a pair of satellite radar images acquired by the 

European Envisat satellite 35 days apart for the same event in 2007. The patterns of fringes indicate 

subsidence of Kīlauea Caldera and a combination of uplift and subsidence near Makaopuhi Crater as 

magma from the summit reservoir intruded into the East Rift Zone (image by Michael Poland, USGS).
10 Characteristics of Hawaiian Volcanoes

InSAR data were collected in Hawai‘i in 2010 and 2011, and 

the results were invaluable in documenting subsurface dike 

processes associated with the March 2011 Kamoamoa fissure 

eruption (Lundgren and others, 2013).

Precise Gravity Monitoring

Measuring changes in the acceleration of gravity, 

coupled with precise leveling, is the only known way to 

measure changes in subsurface mass associated with magma 

movement. The first reliable measurements of gravity changes 

were made in the 1970s. Jachens and Eaton (1980) analyzed 

gravity data for Kīlauea summit stations measured before and 

after a major summit deflation produced by the November 

29, 1975, earthquake and interpreted the results to indicate 

a mass loss, probably from two sources in the south part of 

Kīlauea Caldera. Dzurisin and others (1980) exploited gravity 

measurements to conclude that the November 1975 earthquake 

of magnitude (M) 7.2 (Tilling and others, 1976) created void 

space in the summit area that completely filled with magma 

over the subsequent months, thereby setting the stage for two 

intrusions into the East Rift Zone in mid-1976.

Johnson (1992) interpreted gravity and leveling 

measurements at Kīlauea, specifically for the period 1984–86, 

and distinguished three variables that modulate deflation of 

the magma reservoir: volume and depth of magma transfer; 

pressure and volume of CO2 gas; and spreading of the summit 

area. Kauahikaua and Miklius (2003) interpreted gravity and 

leveling trends from 1983 through 2002 in terms of massstorage changes of Kīlauea’s magma reservoir. Johnson and 

others (2010) examined measurements from late 1975 to 

early 2008 over a broader network of measurement sites to 

document the refilling of void space inferred by Dzurisin and 

others (1980) beneath the summit created by the 1975 M7.2 

earthquake (Tilling and others, 1976). In addition, Johnson and 

others (2010) suggest the 2008 Kīlauea summit vent probably 

tapped the magma that had accumulated since 1975.

Continuous gravity measurements started in 2010 

at Uēkahuna Vault and on the caldera floor above the 

Halema‘uma‘u “Overlook vent” (as the 2008 summit eruptive 

vent has been informally named), and in 2013 at Pu‘u ‘Ō‘ō 

(Michael Poland, oral commun., 2013). These data have 

already documented a previously unknown oscillation with a 

period of 150 s that is almost certainly not seismic in nature. 

Carbone and Poland (2012) suggest that its origin may be 

linked to convective processes within the shallow magma 

reservoir. In addition, gravity changes detected during abrupt 

changes in lava level within the Overlook vent are consistent 

with the near-surface magma having a very low density, 

compatible with a gas-rich foam (Carbone and others, 2013).

Time-Lapse Photography

Localized deformation of the ground surface can also be 

captured in time-lapse photography at very fine spatial and 

temporal scales compared to any other geodetic monitoring 

technique. For example, photographic measurements made 

during 2011 at the Overlook vent within Halema‘uma‘u 

and at Pu‘u ‘Ō‘ō crater showed that the lava lake levels 

varied sympathetically, indicating that an efficient hydraulic 

connection linked Kīlauea’s simultaneous summit and East Rift 

Zone eruptive activity (Orr and Patrick, 2012; Patrick and Orr, 

2012a). Significantly, these photographically documented lava 

level changes “mirror trends in summit tilt and GPS line length” 

(Patrick and Orr, 2012a, p. 66). Tilling (1987) previously 

reported a similar correlation—but based on limited and 

imprecise data—between Kīlauea’s summit tilt and the levels 

of active lava lakes at Mauna Ulu and ‘Alae. However, with 

acquisition of digital, high-resolution time-lapse photographic 

data now possible, detailed measurements of fluctuations in lava 

level can monitor localized changes in magma pressurization 

(inflation versus deflation) in the volcanic plumbing system 

feeding eruptive vents (Patrick and Orr, 2012a; Patrick and 

others, 2014, figs. 6, 7, and 9; Orr, 2014). 

Volcanic Gas Monitoring

Sutton and Elias (this volume, chap. 7) summarize the 

history of volcanic gas studies at HVO during the past century. 

Here, we present a few selected highlights. 

Regular measurements of volcanic gases, especially 

SO2 and CO2, became part of HVO’s monitoring program 

in the late 1970s. Initially, monitoring of gas composition 

was accomplished using direct sampling near eruptive vents 

for gas chromatographic analysis at HVO (Greenland, 1984, 

1987a, 1987b; and references therein). It was during this time 

that remote-sensing techniques began to be used to monitor 

gas-emission rates—correlation spectrometry (COSPEC) for 

SO2 and infrared spectrometry for CO2 (Casadevall and others, 

1987). Since 1987, huge strides have been made in the field of 

remote-sensing measurements—ground-, plane-, and satellitebased—of volcanic gases (see, for instance, Carn and others, 

2003; Nadeau and Williams-Jones, 2008; and Carn, 2011). 

The COSPEC (fig. 6A) was the instrument used to make SO2

emission measurements at Kīlauea through September 2004, 

when—after a period of comparison and calibration with 

newer instruments—it was replaced by a miniature ultraviolet 

spectrometer (nicknamed FLYSPEC; fig. 6B), which is much 

smaller and more portable (Elias and others, 2006; Horton 

and others, 2006). Another regular component of HVO’s 

gas monitoring program uses the Fourier transform infrared 

(FTIR) spectrometer (McGee and Gerlach, 1998; McGee and 

others, 2005). The FTIR is capable of analyzing many other 

species of volcanic gases in addition to SO2 and CO2, thereby 

making possible estimates of the ratios of other gas species not 

directly measured by FLYSPEC.

Over the past two decades, near-real-time remote-sensing 

measurements of gas emission—of SO2 (regularly) and CO2

(infrequently)—have become one of the primary tools, along with 

seismic and geophysical techniques, in HVO’s volcano-monitoring 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 11

program for Kīlauea and Mauna Loa (see, for example, Elias 

and others, 1998; Sutton and others, 2001, 2003; Elias and 

Sutton, 2002, 2007, 2012). Indeed, the time-series data for SO2

emission rates at Kīlauea (fig. 7) acquired from 1979 to the present 

constitutes the longest duration dataset of its type for any volcano 

in the world; CO2 emission measurements at Kīlauea were added to 

the mix starting in 1995 and collected more frequently after 2004. 

Since 1958, atmospheric CO2 levels have been 

continuously monitored by the National Oceanic and 

Atmospheric Administration (NOAA) Mauna Loa Observatory 

(MLO) on the north slope of Mauna Loa (at 3,397 m elevation—

above the inversion layer). Estimates of volcanic CO2 emission 

can be obtained from analysis of the “excess” amounts above 

normal atmospheric levels captured during periods when wind 

directions bring emissions from known volcanic sources to 

the sensors (Ryan, 1995, 2001). For a detailed discussion of 

gas studies and their importance for HVO’s overall volcanomonitoring program, as well as the most recent developments, 

see Sutton and Elias (this volume, chap. 7).

HVO’s increased gas-measuring capability allows for more 

measurements (in different locations) to be made easily, and the 

availability of long-term data now permits identification of possibly 

significant changes in emission rate (in other words, greater than 

“background” variations) from long-measured sources. Beginning 

in 1983, the essentially continuous eruptive activity at Pu‘u ‘Ō‘ōKupaianaha has been accompanied by relatively high rates of 

SO2 emission, fluctuating between <500 and >3,000 metric tons 

of SO2 per day (fig. 7A). This relatively high, nonstop rate of gas 

emission at Kīlauea—from both the summit and the East Rift 

Zone—has produced a persistent “vog” (volcanic smog) that poses 

a significant volcanic hazard for Hawai‘i residents and visitors 

(Sutton and others, 1997). This problem worsened with the onset 

of the 2008-present Halema‘uma‘u eruption at Kīlauea’s summit, 

which greatly increased the SO2 emission rate (fig. 7B) from a 

second location (in addition to the East Rift Zone eruptive vents, 

like Pu‘u ‘Ō‘ō), exacerbating vog conditions in more communities, 

some much closer to the summit vent than Pu‘u ‘Ō‘ō (for detailed 

discussion, see Kauahikaua and Tilling, this volume, chap. 10). 

USB

control

cable

UV window

Motor controllers

for cell wheel

and scanner

Battery for

USB hub

Data

acquisition

computer

GPS

Spectrometer

Motorized SO2

calibration cells

COSPEC

FLYSPEC

A B

Figure 6. Photographs of spectrometers used by the Hawaiian Volcano Observatory (HVO). A, The correlation spectrometer (COSPEC), seen 

here in stationary operating mode at Pu‘u ‘Ō‘ō, was the workhorse instrument used by HVO to measure SO2 emission rates at Kīlauea through 

2004 (USGS photograph by J.D. Griggs). It has been replaced by the lighter, less cumbersome, and lower-cost FLYSPEC, a miniature ultraviolet 

spectrometer. B, The compactness of the latest model of the miniature ultraviolet spectrometer can be appreciated from this schematic (top) 

of the measurement system (Horton and others, 2006, figure 1). Road-based configuration (bottom) for running the FLYSPEC and COSPEC 

instruments side-by-side for experiments conducted during 2002–03 (Elias and Sutton, 2007, figure 2A). For detailed discussion of FLYSPEC 

measurements, see Horton and others (2006), Elias and others (2006), and Sutton and Elias (this volume, chap. 7).
12 Characteristics of Hawaiian Volcanoes

0

1,000

2,000

3,000

4,000

5,000

6,000

7,000

8,000

9,000

10,000

1997 1998 1999 2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010 2011 2012 2013

0

500

1,000

1,500

2,000

2,500

1979 1981 1983 1985 1987 1989 1991 1993 1995 1997 1999 2001 2003 2005 2007 2009 2011 2013

Summit eruption

April 1982 (<1 day)

Onset of

summit eruption

mid-March 2008

Onset of

summit eruption

mid-March 2008

SO2 emission rate, in metric tons per day

A

B

Summit

East Rift Zone

Summit

Year

Figure 7. Graphs of SO2 emission rates at Kīlauea during the period 1979–2013 (updated from Elias and others, 1998; Elias 

and Sutton, 2002, 2007, 2012). A, Comparison of emission rates for the Pu‘u ‘Ō‘ō-Kupaianaha eruption (East Rift Zone) and 

Kīlauea summit during the period 1997–2013. Spikes in rates reflect surges or new outbreaks during eruptive activity. B, Emission 

rates for Kīlauea summit for the period 1979–2013. During 1979–97, rates were higher and more variable than those during the 

1997–2007 period, before sharply increasing with the onset of continuous eruptive activity in mid-March 2008 (summary data 

plots by Tamar Elias, USGS).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 13

Petrologic-Geochemical Monitoring

When Gordon Macdonald joined the HVO staff in 

the early 1940s, he initiated what might be considered the 

forerunner of “petrologic-geochemical” monitoring of 

Hawaiian eruptions—the systematic sampling and analysis of 

eruptive products to examine temporal changes to composition 

as an eruption progresses. Beginning in 1952, every eruption 

at Kīlauea and Mauna Loa, ranging in duration from days to 

decades, has been “sampled at regular intervals to cover both 

the spatial and temporal distribution of eruptive products” 

(Wright and Helz, 1987, p. 628). Such systematic sampling 

(fig. 8) is especially important during prolonged eruptions, 

such as the ongoing Pu‘u ‘Ō‘ō-Kupaianaha eruption (1983 to 

present). Time-series compositional data complement those 

from seismic, geodetic, and gas monitoring to characterize 

preeruption conditions and processes within magma 

reservoirs, as well as syneruptive dynamics (for example, 

Thornber, 2003; Thornber and others, 2003). The availability 

of extensive time-series compositional datasets for Hawaiian 

lavas fostered the development of the breakthrough concepts 

of “olivine control,” “magma batches,” and “magma mixing” 

(see, for instance, Powers, 1955; Wright, 1971), which are 

now widely applied in igneous petrology, especially in studies 

of basaltic volcanism. The obvious synergy between HVO 

petrologic-geochemical monitoring and basic research is 

amply illustrated by the many advances in our petrologic 

understanding of basaltic volcanism, as reviewed by Helz and 

others (this volume, chap. 6). 

Geologic, Petrologic, and Geochemical 

Investigations

The first geological studies of Hawaiian volcanoes 

were made during 1840–41, as part of the U.S. Exploring 

Expedition, commanded by U.S. Navy Lieutenant Charles 

Wilkes (Wilkes, 1844). These early scientific observations 

were led by James Dwight Dana, a 27-year-old, Yale-educated 

natural scientist who had worked at Mount Vesuvius in 

1834. Dana is considered the first American volcanologist, 

and his expedition report “constitutes a virtual textbook of 

Hawaiian volcanology and geology” (Appleman, 1987, p. 

1607). Clarence E. Dutton—the first USGS geologist to work 

in Hawai‘i—spent 5 months in the Hawaiian Islands in 1882 

“for the purpose of studying the features and processes of a 

volcano in action, and thus obtaining the practical knowledge 

which is essential to the investigation of extinct volcanoes” 

(Dutton, 1884, p. xxvi) before starting a new mapping 

assignment in the Cascade Range of the Pacific Northwest. 

The work and findings of Dana (1849) and Dutton have 

guided subsequent geologic studies. In his memoir, Dana 

(1890) summarized concepts learned from study of Hawaiian 

volcanoes, listing topics needing further study in the preface 

of that volume (table 1). As is obvious from the now-abundant 

volcanologic literature, HVO and many other investigators 

continue to address many of these fundamental research 

questions—first emphasized by Dana more than a century 

ago—to better understand basaltic volcanism in Hawai‘i and, 

by extrapolation, elsewhere.

Figure 8. In the regular 

petrologic-geochemical 

monitoring of eruptive 

products, Hawaiian 

Volcano Observatory 

(HVO) scientists, using 

a steel cable with a 

hammerhead attached, 

collect a sample through 

a skylight of an active 

lava tube originating at 

Pu‘u ‘Ō‘ō vent. Molten 

lava adheres to the 

hammerhead dangling at 

edge of skylight above 

the flowing lava stream. 

(USGS photograph by 

J.D. Griggs, December 5, 

1990).
14 Characteristics of Hawaiian Volcanoes

Observations of Kīlauea Lava Lake Activity

Because of its frequent and sustained effusive eruptions 

and relative accessibility, Kīlauea Volcano has served as 

a prime natural laboratory for systematic studies of active 

basaltic lava lakes. Table 2 is a compilation of episodes of 

lava-lake activity at Kīlauea, lasting 30 or more days, during 

the period pre-1823 to the present.

Between Jaggar’s first visit to Kīlauea in 1909 with fellow 

MIT geologist R.A. Daly to early 1924, the continuously active 

lava lake at Halemaʻumaʻu Crater provided a natural and easily 

accessible focus for systematic visual observations of eruptive 

activity at Kīlauea. The accounts of the earliest observers were 

conveniently summarized in Brigham (1909) and Hitchcock 

(1911). Daly incorporated his 1909 lava-lake observations into a 

larger work on volcanism (Daly, 1911), concluding that (1) the 

circulation of lava in the shallow, saucer-shaped Halemaʻumaʻu 

lakes was driven by two-phase magmatic convection within a 

narrower conduit below; (2) lava fountains were the result of 

“explosive dilation” of entrained gas bubbles as magma rose 

within that conduit; (3) the lake surface was a “froth”; and 

(4) the predominant gas release was at lake edges, where lava 

pooled. Based on night-and-day observations in 1911, Perret 

(1913a,b,c,d) came to the conclusion that lava-lake circulation 

was driven by the sinking of cooled, degassed slabs at the 

eastern end of the lake into the same conduit from which fresh, 

vesiculating magma was rising before flowing westward along 

the bottom to emerge at the western end of the lake.

Jaggar’s much longer series of lava-lake observations 

led him to more refined ideas of the active lava lake and 

associated processes within the Halema‘uma‘u pit crater 

(Jaggar, 1920). He correctly concluded that gases are in 

solution in magma at depth (“hypomagma”) and only start 

to vesiculate in the magma as it rises (“pyromagma”). Once 

the lava loses its gas while in the lake, it becomes denser 

and more viscous (“epimagma”) and collects at the base and 

along the edges of the lake, as well as draining back into 

 “But much remains to be learned from further study of the Hawaiian volcanoes. Some of the points requiring elucidation are the following:

• the work in the summit-crater between its eruptions;

• the rate of flow of lava-streams and the extent of the tunnel-making in the flow;

• the maximum thickness of streams;

• the existence or not of fissures underneath a stream supplying lava;

• the temperature of the liquid lava;

• the constitution of the lava at the high temperatures existing beneath the surface;

• the depth at which vesiculation begins;

• the kinds of vapors or gases escaping from the vents or lakes;

• the solfataric action about the craters;

• the source of the flames observed within the area of a lava lake;

• the differences between the lavas of the five Hawaiian volcanoes . . . ;

• the difference in kind or texture of rock between the exterior of a mountain and its deep-seated interior or centre . . . ;

• the difference between Loa, Kea, and Haleakala in the existence below of hollow chambers resulting form lava discharges . . .

• the movements of the lavas in the great lava columns, and the source or sources of the ascensive movement.”

Table 1. Some aspects of Hawaiian volcanism that geologist James D. Dana thought still needed investigation in 1890 (excerpted from Dana, 

1890, p. vii–viii).

Table 2. Eruptions at Kīlauea Volcano over the past ~200 years.

Eruption site Year 

start

Duration 

(days)1 Selected references

Halema‘uma‘u Pre-1823 >35,000

(intermittent)

Jaggar (1947) 

Bevens and others (1988)

Halema‘uma‘u 1934 33 Jaggar (1947)

No eruptive activity at Kīlauea during 1935–51

Halema‘uma‘u 1952 136 Macdonald (1955)

Kīlauea Iki 1959 36 Richter and Eaton (1960)

Richter and others (1970)

Halema‘uma‘u 1967 251 Kinoshita and others (1969)

Mauna Ulu 1969 875 Swanson and others (1979)

Mauna Ulu 1972 455 Tilling (1987)

Tilling and others (1987a)

Mauna Ulu 1973 222 Tilling and others (1987a)

Pauahi 1979 30 Tilling and others (1987a)

Pu‘u ‘ŌʻōKupaianaha1983 >11,500 

(intermittent, 

ongoing)

Wolfe and others (1987)

Wolfe and others (1988)

Heliker and others (2003a)

Orr and others (2012)

Halema‘uma‘u >2,400 

(ongoing)

Hawaiian Volcano 

Observatory Staff (2008)

Patrick and others (2013)

1

Updated from Peterson and Moore (1987, table 7.3).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 15

the hypomagma at depth. In addition, Jaggar hypothesized 

that deformation around and within the lake occurred by 

pressurization of the epimagma rather than the solid rock 

surrounding the crater and lake. Jaggar made many direct 

measurements of the lava lakes (Jaggar, 1917a,b), confirming 

that they were shallow (only ~14–15 m deep) and hotter at 

the bottom than at the surface (confirming Perret’s circulation 

model). Lake depths were determined by insertion of steel 

pipes, into which Seger cones were lowered to estimate 

temperatures (see Apple, 1987, fig. 61.5; also see discussion 

in “Field Measurements of Lave Temperature” section, 

below). The time series of lava-lake elevations compiled by 

Jaggar was later used by Shimozuru (1975) to demonstrate a 

semidiurnal oscillation correlated with Earth tides.

The 1969–74 Mauna Ulu eruptions produced the first 

sustained lava-lake activity in a rift zone of Kīlauea during 

historical time. During the 1969–71 activity, nearly continuous 

lava lakes exhibited “gas pistoning”—an episodic release of gas 

accumulated beneath the lake’s crust—that was first described 

by Swanson and others (1979). During 1972–74, fluctuations 

of the surface height of the Mauna Ulu lava lakes correlated 

with variations in Kīlauea’s summit tilt, indicating an efficient 

hydraulic linkage between the summit magma reservoir and the 

active lava lakes (Tilling, 1987). Later, gas-piston behavior was 

also commonly observed within pits in the floor of Pu‘u ‘Ō‘ō 

crater (for example, Orr and Rea, 2012). Although not a lake per 

se, a perched lava channel that formed at Pu‘u ‘Ō‘ō in mid-2007 

also exhibited gas-piston-like gas release cycles (Patrick and 

others, 2011a) and seeps of much denser, more pasty lavas from 

its base (“epimagma” in Jaggar’s terms).

The return of prolonged lava-lake activity to 

Halema‘uma‘u in 2008 and to Pu‘u ‘Ō‘ō in 2011 has provided 

new opportunities to quantify and refine the observations made 

throughout HVO’s history. Webcam records of active lavalake activity have allowed categorization and quantification of 

typical lake behaviors, for example, hours-long rise/fall events 

during which lava levels gradually rise while gas emissions and 

seismic tremor levels drop, followed, in turn, by a quick lavalevel drop and resumption of gas emissions and tremor; summit 

lava lake levels that track summit tilt records; and variation 

of lake-circulation patterns associated with the location of 

spattering sinks (as first described by Daly, 1911, and Perret, 

1913a,b). Continuous gravity measurements during lava lake 

level changes suggest that at least the upper few hundred meters 

of the magma column is a foam with a density of about 1,000 

kg/m3 (Carbone and others, 2013).

Because Kīlauea lava lakes are located within a dense 

seismic network, recent lake observations are closely tied 

to unique seismic signatures. For example, the continued 

stoping of the conduit in which the Halema‘uma‘u lava 

lake sits frequently produces rock falls directly into the 

lava lake. These rock falls, in turn, commonly produce a 

characteristic composite seismic signature that starts with a 

high-frequency signal (presumably the rock breakage) that 

grades into long-period (LP) frequencies (interaction with 

the lava lake) and, ultimately, into a very-long-period (VLP) 

signal that can last for several minutes (possibly related to 

deep pressure oscillations within the magma column; Patrick 

and others, 2011a). Moreover, the rock falls into active lava 

lakes are known to trigger explosive events (Orr and others, 

2013) and initiate rise/fall events. 

Origin and Fractionation of Hawaiian Magmas

The earliest geochemical investigations at Hawaiian 

volcanoes involved the sampling and analysis of volcanic 

gases—all at Kīlauea except for two samples from Mauna 

Loa—by A.L. Day, E.S. Shepherd, and T.A. Jaggar during 

the period 1912–19 (Day and Shepherd, 1913; Shepherd, 

1919, 1921). Some of these early collections of gas from 

Halema‘uma‘u are still considered among the best in the world 

in terms of sample purity and analytical precision (Gerlach, 

1980; Greenland, 1987b). After these notable investigations, 

the study of eruptive volcanic gases then languished for 

decades, only to be resumed sporadically in the 1960s, mostly 

centered on specific short-lived eruptions or intrusions, 

including the beginning of the Pu‘u ‘Ō‘ō eruption of Kīlauea 

and the 1984 eruption of Mauna Loa (Greenland, 1987a). 

For informative summaries of gas studies through the mid1980s, the interested reader is referred to Greenland (1987a,b, 

and references therein). Beginning in the 1990s, systematic 

studies of the composition and rate of gas emission—using 

continuously recording optical or multispectral remote sensing 

of gas species—became an integral component of HVO’s 

current volcano-monitoring and research program (see Sutton 

and Elias, this volume, chap. 7).

Other than HVO’s volcanic gas studies during 

its early decades, however, “systematic collection and 

characterization of samples from eruptions, either 

megascopically, or by petrographic and chemical analysis” 

were lacking for much of the early 20th century (Wright, 

1989, p. xix). It was not until the early 1940s, with the 

arrival of Gordon Macdonald, that regular collection of lava 

and tephra samples for laboratory analysis was inaugurated 

at HVO; he was the first to make a comprehensive petrologic 

and geochemical study of Hawaiian lavas (Macdonald, 

1949a,b). In the 1950s Howard A. Powers introduced the 

now commonly used magnesia-variation diagrams, olivinecontrol lines, and the concepts of magma batches and magma 

mixing in analyzing chemical variations in erupted lava 

(Powers, 1955). Illustrative examples of plots of times-series 

compositional data for MgO and other oxides (bulk lava 

or glass) are used in Helz and others (this volume, chap. 6, 

their figures 3, 6, and 11). Perhaps unique for investigations 

of basaltic volcanism, shallow magma crystallization and 

fractionation processes have been documented directly by 

petrologic-geochemical studies of samples collected by 

drilling into passive lava lakes (see “Drilling Studies of 

Passive Historical Lava Lakes,” below, and Helz and others, 

this volume, chap. 6, for additional discussion). Figure 9 

petrographically demonstrates the progressive crystallization 
16 Characteristics of Hawaiian Volcanoes

with decreasing temperature of the still-molten portion of the 

solidifying 1965 lava lake in Makaopuhi Crater (Wright and 

Okamura, 1977).

Powers was the first to recognize that the historical 

lavas of Kīlauea and Mauna Loa are petrographically and 

chemically distinct. Building on Powers’s pioneering work, 

Wright (1971) produced a definitive compilation of the 

composition of Kīlauea and Mauna Loa lavas using all 

chemical and petrographic analyses available at the time. 

He proved that Mauna Loa lava compositions showed no 

correlation with time of eruption, nor with vent location, 

whereas Kīlauea lavas could be compositionally grouped 

according to eruption age and vent site (summit vs. rift zones). 

The work of Wright and his associates during the 1970s (for 

instance, Wright, 1971; Wright and Fiske, 1971; Wright and 

others, 1975; Wright and Tilling, 1980) set the stage for many 

subsequent petrologic-geochemical studies germane to the 

origin and fractionation of Hawaiian basalt (see, for example, 

Garcia and Wolfe, 1988; Garcia and others, 1989, 1992, 1998, 

2000). For comprehensive reviews of petrologic-geochemical 

studies of Hawaiian lava, the interested reader is referred to 

the many summary works (for example, Wright and Helz, 

1987; Rhodes and Lockwood, 1995; Tilling and others, 1987b; 

Helz and others, this volume, chap. 6, and references therein). 

Geological Mapping of Hawaiian Volcanoes

As noted by Wright (1989), except for partial maps of 

the lava flows of the 1840 Kīlauea eruption (Wilkes, 1844) 

and maps of some other historical eruptions produced by 

the Territorial Government Survey Office (Brigham, 1909; 

Hitchcock, 1911), early studies of Hawaiian volcanism 

lacked accurate and complete geologic maps. Beginning in 

the late 1920s, geological reports and accompanying maps 

for many of the Hawaiian Islands were produced by USGS 

geologists Harold T. Stearns and his associates, particularly 

Gordon Macdonald (see Sherrod and others, 2007, and 

references therein). The geologic map for the Ka‘ū District 

(Stearns and Clark, 1930) inaugurated the era of systematic 

geologic mapping in Hawai‘i, and Stearns and Macdonald 

(1946) compiled the first geologic map of the entire Island 

of Hawai‘i. Moreover, published accounts of eruptions at 

Kīlauea and Mauna Loa since the 1950s include reasonably 

good maps of erupted lava flows (for example, Swanson and 

others, 1979; Tilling and others, 1987a; Wolfe and others, 

1987; Lockwood and others, 1987; Heliker and others, 2003b). 

Equally and perhaps more important, the newer mapping also 

delineated deposits of prehistoric eruptions, thereby allowing 

longer term reconstruction of eruptive history and hazards 

assessment (for instance, Peterson, 1967; Walker, 1969; 

Lipman and Swenson, 1984; Holcomb, 1987; Lockwood 

and Lipman, 1987; Lockwood and others, 1988; Moore and 

Clague, 1991; Buchanan-Banks, 1993; Neal and Lockwood, 

2003). Assignments of absolute or relative ages to prehistoric 

lavas were made increasingly possible by careful mapping of 

1,170 °C

1,065 °C

760 °C

Decreasing temperatures

Figure 9. Photomicrographs (field of view ~1.2 mm) of samples 

collected by drilling through the crust of the cooling 1965 lava 

lake in Makaopuhi Crater (Wright and Okamura, 1977). The 

lava-lake drilling operations at Makaopuhi were similar to those 

conducted in 1975 for Kīlauea Iki lava lake (see fig. 20). The 

samples, which represent the still-molten portion of the lava lake 

at the time of sampling and at the in-hole temperatures indicated, 

show increasing crystallinity and sequential appearance of mineral 

phases with decreasing temperature. The progressive darkening of 

the glassy matrix reflects the presence of submicroscopic Fe-Tioxide grains and imperfect quenching (USGS photomicrographs by 

Richard S. Fiske; image from Tilling and others, 2010).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 17

stratigraphic relationships and refinements in radiometric and 

paleomagnetic dating methods during the latter part of the 

20th century (discussed below).

In the mid-1980s, HVO launched the Big Island Map 

Project (BIMP) to update the geologic map of the Island of 

Hawai‘i, based on maps generated from 1975 to 1995 by more 

than 20 geologists from the USGS and various universities. 

The new compilation (Wolfe and Morris, 1996a; Trusdell 

and others, 2006) represented a quantum advance over the 

1946 map in its detailed portrayal of the distribution and 

ages of prehistoric, as well as historical, eruptions (fig. 10). 

In addition, this compilation also includes location maps of 

all radiocarbon dating sites and samples collected for majorelement chemical analysis (Wolfe and Morris, 1996b). The new 

mapping confirmed the geologic youthfulness of Mauna Loa 

and Kīlauea volcanoes inferred by earlier investigators. About 

90 percent of Mauna Loa’s surface is covered by lava younger 

than ~4,000 years old; Kīlauea’s surface is even younger, 

with ~90 percent plated with lavas younger than 1,100 years 

(Holcomb, 1987; Lockwood and Lipman, 1987; Wolfe and 

Morris, 1996a). The original geologic mapping by Stearns and 

A B

0

0

20 40 KILOMETERS

10 20 MILES

colleagues, updated in places by various geologists, including 

new Haleakalā mapping by USGS geologist Dave Sherrod, 

and the BIMP data over the intervening years, was registered 

against modern topographic maps, compiled and published as a 

state geologic map in modern geographic information systems 

(GIS) formats by Sherrod and others (2007). More detailed 

geologic mapping of Mauna Loa volcano, to be published at 

a 1:50,000 scale in five sheets, is currently underway (F.A. 

Trusdell, oral commun., 2012).

Radiometric and Paleomagnetic Dating of Lava Flows

Geological mapping at Hawaiian volcanoes benefited 

greatly from the advent and development of radiometric 

and paleomagnetic dating methods in the 20th century. 

Potassium-argon age determinations were crucial in 

establishing the age progression of the Emperor Seamount–

Hawaiian Ridge volcanic chain (for example, Clague and 

Dalrymple, 1987; Langenheim and Clague, 1987; Clague and 

Sherrod, this volume, chap. 3). For dating the lavas erupted 

Figure 10. Geologic maps of the Island of Hawai‘i. A, The first geologic map for the entire island (published scale 1:125,000), by Stearns and 

Macdonald (1946), who based their compilation on all then-existing geologic information. Although remarkable and useful in its time, this map 

lacked geologic definition and age data for the prehistoric volcanic deposits that underlie most of the island. B, The geologic map (published scale 

1:100,000) compiled in 1996 by Wolfe and Morris (1996a), was the culmination of the Big Island Map Project (see text). This newer map clearly 

shows in much greater detail the distribution of the prehistoric deposits, reflecting the improved knowledge gained from mapping and dating 

studies conducted since 1946 (images from Babb and others, 2011).
18 Characteristics of Hawaiian Volcanoes

at Hawai‘i’s active volcanoes, however, radiocarbon (14C) 

and paleomagnetic dating techniques have proven the most 

useful. During the first 75 years of HVOʼs history, many 

hundreds of radiocarbon dates on carbon-bearing samples 

(see Rubin and Suess, 1956; Rubin and others, 1987) were 

used in the preparation of Hawaiian geologic maps and the 

Wolfe and Morris compilation (1996a,b). Since 1987, many 

more radiocarbon dates have been obtained and used in the 

preparation of updated geologic maps and refined eruption 

frequencies for Hawaiian volcanoes (for instance, Sherrod and 

others, 2006, 2007; Trusdell and Lockwood, in press).

Oftentimes, finding carbon-bearing materials in the field 

suitable for dating is a challenge. From practical experience 

gained during the mapping of Mauna Loa, HVO scientists 

learned to identify the most favorable field settings to find 

datable carbon (see Lockwood and Lipman, 1980), which 

has benefited investigators studying some other basaltic 

volcanoes (for example, in the Canary Islands and the 

Azores). Paleomagnetic dating of lavas, by matching the 

preserved magnetic direction in lava with the record of the 

secular variation in the Earth’s magnetic field, has been used 

with good success in Hawai‘i (for instance, Holcomb and 

others, 1986; Holcomb, 1987; Lockwood and Lipman, 1987). 

This method is premised on determination of the directions 

in remnant magnetization of well-dated (generally by 

radiocarbon) Hawaiian samples that span a range of geologic 

time. Lavas of unknown age can be dated or time-bracketed 

by comparing their magnetization directions with the history 

of secular variation of the magnetic field (for example, 

Holcomb and others, 1986; Hagstrum and Champion, 1995; 

Clague and others, 1999; Sherrod and others, 2007). Dozens 

of new radiocarbon ages, combined with paleomagnetic 

data, were obtained to augment field studies in recent 

reconstructions of the eruptive history of geologically recent 

explosive activity at Kīlauea (for instance, Fiske and others, 

2009; Swanson and others, 2012a).

How Hawaiian Volcanoes Work

From century-long observations of Kīlauea’s changing 

eruptive activity between the summit and rift zones and 

decades-long monitoring data and topical research, we have 

a general conceptual model of the subsurface magmatic 

systems that sustain Hawaiian volcanism (see, for example, 

Decker, 1987). Although specifics about the deep (>20 km) 

source regions are still poorly resolved, this model depicts 

magma generation and ascent in the mantle, its transport 

to and storage within one or more shallow reservoirs, and 

ultimately its eruption at the surface. In this section, we 

touch upon some selected attributes and operative processes 

of Hawaiian volcanic plumbing systems; the other chapters 

in this volume treat in greater detail, and for differing time 

scales (geologic vs. historical), how Hawaiian volcanoes 

have worked in the past and are working at the present time.

Shallow Magma Reservoirs in Shield Volcanoes

The first inferences about Kīlauea magma chambers came 

from an analysis of leveling data acquired before and after 

the explosive eruptions of May 1924. These data documented 

dramatic subsidence of the caldera floor associated with the 

retreating lava column and related decreases in a subsurface 

pressure source at a depth of 3.5 km below the south part of 

Kīlauea Caldera. “The changes of the pressure of the spherical 

sources may correspond perhaps to the decrease of the hydrostatic pressure of magma reservoirs below the area and may 

have been caused by the intrusion or the extrusion of magma 

from the reservoirs” (Mogi, 1958). Later, using data collected 

by HVO and other scientists through the 1950s, Eaton and 

Murata (1960, fig. 5) proposed the first dynamic model for the 

magmatic system of Kīlauea. This model, based primarily on 

seismic and other geophysical data, involved the following 

elements: (1) a magma-source region deeper than 60 km; (2) 

poorly defined pathways (“permanently open conduits”) for 

magma to rise into the crust; and (3) collection and storage 

of magma in a shallow (<5 km) summit reservoir for a finite 

time, from which magma later is erupted at the summit or 

intruded and possibly erupted along a rift zone. The basic 

features of this now half-century-old “classic” model (fig. 

11) still apply, but decades of additional data and subsequent 

analyses (for example, Dzurisin and others, 1984; Ryan, 

1987; Delaney and others, 1993; Tilling and Dvorak, 1993; 

Dvorak and Dzurisin, 1997; Cervelli and Miklius, 2003; 

Wright and Klein, 2006, 2014; Poland and others, 2012, this 

volume, chap. 5) have sharpened the model and introduced 

some important refinements. Variations of the Eaton-Murata 

model include the concept that Kīlauea’s summit region may 

contain more than one shallow reservoir beneath the surface 

(for example, Fiske and Kinoshita, 1969), that secondary 

shallow reservoirs operate within the rift zones, and that 

magma pathways between the summit and rift zones may 

extend to greater depths than previously inferred. Poland and 

others (this volume, chap. 5, fig. 10) introduce a new model 

consistent with all that has been learned from recent decades 

of monitoring and modeling studies. 

On the basis of differences in lava chemistry and patterns 

of eruptive behavior, nearly all post-1950 studies favor the view 

that the volcanic systems of Kīlauea and Mauna Loa apparently 

operate independently, even though they may tap, at depth, the 

same magma-source region. A study by Miklius and Cervelli 

(2003), however, suggests that, during 2001–02, a “crustal-level 

interaction”—though not necessarily a physical connection—

may have linked these two volcanoes, and Gonnermann and 

others (2012) suggest a zone of porous melt accumulation in the 

upper mantle beneath both volcanoes that can transmit pressure 

variations between the two systems. Because the subsurface 

configurations of sustaining volcanic systems figure directly or 

indirectly in many studies of Kīlauea and Mauna Loa eruptions 

(for example, Decker and others, 1987; Heliker and others, 2003b; 

Poland and others, 2012, this volume, chap. 5; and references cited 

therein), we do not discuss them further here. 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 19

Long-Term Inflation-Deflation Cycles

As is now well known, the premonitory indicators of a 

Hawaiian eruption typically include heightened seismicity, 

inflationary ground deformation, and increased gas emission. As 

the shallow summit reservoir experiences magma influx and (or) 

hydrothermal pressurization, the volcano undergoes swelling or 

inflation. The inflation, in turn, deforms the volcano’s surface, 

and this can be tracked in real time by geodetic monitoring, 

particularly variation in summit tilt. For Hawaiian volcanoes, 

preeruption inflation is generally gradual and can last for weeks 

to years. Once an eruption begins, however, the summit reservoir 

typically undergoes rapid contraction or deflation as pressure is 

suddenly relieved. During deflation, changes in tilt and in vertical 

and horizontal displacements of benchmarks are opposite to 

those during inflation. Kīlauea’s behavior during and between 

eruptions is remarkably regular, generally characterized by the 

so-called inflation-deflation cycle, first recognized in 1930 and 

perhaps earlier (Jaggar, 1930). These cycles are observed for nearly 

every Hawaiian eruption or intrusion regardless of size (fig. 4). 

The largest tilt change (nearly 300 microradians), to date, for an 

inflation-deflation cycle was recorded for the Kīlauea Iki eruption 

in 1959; however, similar cycles can involve much smaller tilt 

changes (<20 microradians), as during episodes 1–20 (1983–84) of 

the ongoing Pu‘u ‘Ō‘ō activity (Wolfe and others, 1988, figure 1.2).

Effusive Eruptions

Except for the 1924 explosive activity and minor ash 

emissions since 2008 at Kīlauea’s summit, only effusive eruptions 

have occurred during HVO’s first 100 years of existence. 

Consequently, volcano monitoring and research efforts on 

Hawaiian volcanoes have focused overwhelmingly on eruptive 

processes and products of nonexplosive historical eruptions, and a 

vast body of scientific literature on effusive volcanism in Hawai‘i 

has accrued accordingly. These extensive data have prompted the 

worldwide use of the adjective “Hawaiian” to describe effusive or 

“gentle” eruptions wherever they happen—in other words, assigned 

with rankings of 0 to 1 in the Volcanic Explosivity Index (VEI) 

of Newhall and Self (1982). Because “Hawaiian-style” eruptive 

activity has received much scientific attention (for example, Decker 

and others, 1987; Cashman and Mangan, this volume, chap. 9, 

and references therein), there is no need for us to elaborate here. 

In contrast, however, recent studies indicate that, before 1800, 

explosive eruptions at Kīlauea were common and powerful.

0

0

5

10

20

30

40

Kīlauea

Caldera

Southwest Rift Zone

East Rift Zone

Halema‘uma‘u Crater

Hawaiian

Volcano Observatory

Mauna

Ulu Pu‘u ‘Ō‘ō

Shallow magma

reservoir

Pacific Ocean

Depth,

in kilometers

Figure 11. Cartoon of the 

“classic” model of Kīlauea’s 

volcanic plumbing system, 

showing magma input from 

depth, the shallow reservoir, 

and the volcano’s two rift zones 

(highly simplified from Ryan, 

1987, figure 52.29, and as used 

in Tilling and others, 2010).
20 Characteristics of Hawaiian Volcanoes

Explosive Eruptions

Recognizing pyroclastic deposits in addition to lavaflows 

at Kīlauea and elsewhere in Hawai‘i, early observers and 

investigators reasoned that Hawaiian volcanoes have erupted 

explosively in the past (for instance, Ellis, 1825; Emerson, 1902; 

Hitchcock, 1911; Perret, 1913e; Powers, 1916; Stone, 1926; 

Wentworth, 1938; Powers, 1948). Indeed, Hitchcock (1911, p.167) 

concluded that “Kilauea . . . has not always been the tame creature 

of today” (as quoted in Swanson and others, 2012a, p.8). A map 

showing the generalized distribution of the known ash deposits of 

Hawai‘i was published by Wentworth (1938, fig. 6); however, to 

date, other than in the Kīlauea region, pyroclastic deposits have 

not been well studied. The extensive and deeply weathered ash 

deposits (“Pāhala Ash”) southwest of Kīlauea were originally 

thought to have issued from a buried caldera in the Pāhala area. 

It is now thought that these deposits may have been produced 

from explosive activity of Kīlauea Volcano (Sherrod and others, 

2007). Beginning in the 1980s, studies of the products of explosive 

Kīlauea eruptions have shown that Kīlauea is much more 

explosive, energetic, and hazardous than anyone but the early 

geologists considered (Decker and Christiansen, 1984; McPhie 

and others, 1990; Mastin, 1997; Mastin and others, 1999, 2004; 

Fiske and others, 2009; Swanson and others, 2011a,b, 2012a,b). 

The May 1924 eruptions at Halema‘uma‘u (fig. 12A)—

Kīlauea’s only significant explosive activity during recorded 

history—occurred after the withdrawal 3 months earlier of the 

lava lake, and they were well documented by HVO (Jaggar, 

1924; Jaggar and Finch, 1924). The cause for this eruption is 

interpreted to be the interaction of groundwater with Kīlauea’s 

shallow magma reservoir (see, for example, Decker and 

Christiansen, 1984; Dvorak, 1992). Yet this eruption was much, 

much smaller than the phreatomagmatic eruption in 1790 C.E. 

(McPhie and others, 1990; Mastin, 1997; Mastin and others, 

1999, 2004). Base surges associated with this powerful explosive 

event caused many fatalities, which have been estimated decades 

after the eruption to range from 80 to >5,000 (Swanson and 

Christiansen, 1973; Swanson and others, 2011b). Even the 

minimum-fatalities estimate would make Kīlauea historically 

the deadliest of all U.S. volcanoes. Recent detailed geologic and 

dating studies reveal that the 1790 event was only one of many 

explosive eruptions throughout a 300-year period between about 

1500 and 1800 C.E. (Fiske and others, 2009; Swanson and others, 

2011a,b, 2012a,b). These ongoing studies have led to a new 

interpretation of Kīlauea’s long-term volcanic behavior: during 

the past 2,500 years of its eruptive history, centuries of effusive 

eruptions have alternated with centuries of dominantly explosive 

ones when there exists a deep summit caldera (Swanson and 

others, 2011a,b). Moreover, the new data about pre-1924 

pyroclastic deposits also have resulted in two findings that entail 

implications for volcanic hazards (see Kauahikaua and Tilling, 

this volume, chap. 10): Kīlauea erupts explosively about as often 

as does Mount St. Helens in the Cascade Range of Washington; 

and at least six of the pre-1924 explosive eruptions were highly 

energetic, sending centimeter-size lithic clasts to jetstream heights 

(Fiske and others, 2009; Swanson and others, 2011b). 

Flank Instability

Repeat regional triangulation measurements during the late 

19th and early 20th centuries (for example, Wilson, 1927, 1935; 

Wingate, 1933) hinted that some benchmarks on Kīlauea Volcano 

were not fixed. Although the instability of Kīlauea’s south flank 

was recognized by Moore and Krivoy (1964), the actual model of 

seaward displacement of the south flank by rift dilation was first 

presented by Fiske and Kinoshita (1969). Since then, past and 

contemporaneous volcano flank movement has been confirmed by 

further geologic studies and abundant geodetic measurements (see, 

for example, Owen and others, 1995, 2000).

B

A

Figure 12. Photographs evidencing explosive behavior by Kīlauea 

Volcano. A, A crowd of visitors posing in front of the May 24,1924, 

explosive eruption plume from Halema‘uma‘u Crater (photograph 

courtesy of the Bishop Museum, Honolulu). Such explosive activity 

occurs when magma in the shallow summit reservoir interacts with 

groundwater. A person who ventured too close to take a picture was 

killed by a falling volcanic block. B, Pyroclastic deposits are well exposed 

in this cliff (~12 m high) southwest of Keanakāko‘i Crater, Kīlauea 

summit. These strata are part of a phreatomagmatic sequence (the 

“Keanakāko‘i Tephra” of Swanson and others, 2012a) produced by 

explosive activity much more powerful than the 1924 explosive eruption 

(see text; USGS photograph by Donald A. Swanson).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 21

Tracking Flank Movement

As HVO’s ground-deformation networks expanded and 

became more sophisticated, beginning in the 1970s, EDM and later 

GPS data demonstrated that this persistent seaward displacement 

occurs at average rates of several centimeters per year (fig. 13). 

The first comprehensive study of volcano flank movement by 

Swanson and others (1976), made possible by the development of a 

two-color, long-distance laser EDM for making much longer shots 

between benchmarks, concluded that the driving force for flank 

displacements was intrusion of magma into the rift zones. During 

catastrophic flank motion events, such as the M7.7 earthquake in 

November 1975 (Tilling and others, 1976, Nettles and Ekström, 

2004), the sudden seaward flank displacements are much greater 

than those measured during “normal” volcanic activity (Lipman 

and others, 1985). Since then, many other studies—involving both 

ground- and space-based geodetic techniques—have generally 

reinforced these findings but differ in details of the conceptual 

models for flank displacements (for example, Lipman and others, 

1988, 1990, 2002; Delaney and others, 1990, 1993; Cayol and 

others, 2000; Denlinger and Morgan, this volume, chap. 4; Poland 

and others, this volume, chap. 5). 

PUOC

155°20' 155°15' 155°10' 155°05' 155° 154°55'

19°10'

19°15'

19°20'

19°25'

19°30'

Kīlauea Caldera

East Rift Zone

0

0

5 10 KILOMETERS

10 5 0 CENTIMETERS 2.5 5 MILES

Slow Slip Events

First reported in 2001 for the Cascadia subduction zone 

(Dragert and others, 2001), aseismic slip events are now recognized 

as common in subduction regimes around the world (see Beroza 

and Ide, 2011). Cervelli and others (2002a, p. 1015) recognized 

very similar sudden aseismic slip of Kīlauea’s south flank in 2000, 

based on data from the continuous GPS network—the first such 

“recording of a silent earthquake in a volcanic environment.” 

The 2000 slow slip event was equivalent to an earthquake with a 

moment magnitude of 5.7 but without the seismic shaking. Since 

this initial discovery, several more “slow slip” events on Kīlauea’s 

south flank have been recognized, including a family of periodic 

events that occur every 2.2 years, give or take a month or two 

(Brooks and others, 2006; Poland and others, 2010). Each event is 

also associated with characteristic seismicity within the south flank 

(Segall and others, 2006a,b; Wolfe and others, 2007; MontgomeryBrown and others, 2009). A slow slip event in mid-2007 was 

thought to be triggered by a rift-zone dike intrusion by Brooks 

and others (2008, p. 1177), who suggested that “both extrinsic 

(intrusion-triggering) and intrinsic (secular fault creep) fault 

processes” can produce slow slip events at Kīlauea. 

Figure 13. Map of Kīlauea showing continuously recording Global Positioning System (GPS) measurements 

obtained by the Hawaiian Volcano Observatory (HVO) geodetic networks during 2003–06. Station locations are at 

the tails of vectors, and circles indicate 95-percent confidence levels. The vectors show the horizontal movement 

of the ground surface during that period. They indicate net inflation of the summit region, as well as document the 

persistent, ongoing seaward displacement of Kīlauea’s mobile south flank at rates of several centimeters per year 

(image by Michael Poland and Asta Miklius, USGS).
22 Characteristics of Hawaiian Volcanoes

Geophysical Investigations of Subsurface 

Volcano Structure

The International Geophysical Year (1957–58) and the 

prospect of siting a deep research drill hole north of Maui spurred 

geophysical work on all the islands by HVO and University 

of Hawai‘i at Mānoa scientists in the early 1960s. The drilling 

project, dubbed the “Mohole,” was scrapped by lack of funding, 

but the scientific gains from the geophysical work were profound. 

Using a recently expanded and updated seismic network at 

Kīlauea Volcano, Eaton and Murata (1960) determined a basic 

seismic velocity structure for the Island of Hawai‘i from P-wave 

arrival times while Kinoshita and others (1963) mapped the 

gravity variations caused by the island’s substructure. The 

results of these studies, together with follow-up seismic studies 

(summarized by Hill and Zucca, 1987), revealed that the weight of 

the island’s volcanoes depresses the oceanic crust below, causing 

it to flex. The internal structure of the volcanoes was consistent 

with a model of dense cores beneath summits and rift zones. 

More recently, detailed gravity mapping (Kauahikaua and others, 

2000) and seismic tomography (Okubo and others, 1997) revealed 

discernible variations within those dense cores.

In 1978, the USGS conducted several airborne magnetic and 

very-low-frequency (VLF) electromagnetic resistivity surveys 

over the Island of Hawai‘i (Flanigan and Long, 1987). These 

surveys indicated that Kīlaueaʼs East Rift Zone is characterized 

by a highly magnetized shallow linear zone superposed over a 

deeper, broad unmagnetized body. Hildenbrand and others (1993) 

reinterpreted the magnetic data in terms of magnetic properties 

of basalts and concluded that the deep, unmagnetized body was 

hydrothermally altered rock, and that variations in the highly 

magnetized shallow zone defined cooling intrusions. 

Hawaiian Volcano

Observatory

All of these geophysical data, combined with the results 

of resistivity soundings and self-potential surveys, defined the 

hydrothermal systems of Kīlauea Volcano (Kauahikaua, 1993). 

High-level groundwater within the summit region and convection 

within the aquifer in the vicinity of thermal sources in the summit 

and rift zones produce large self-potential electrical anomalies 

at the surface (Zablocki, 1978). Bedrosian and Kauahikaua 

(2010) recently revisited the use of electromagnetic resistivity 

and self-potential monitoring to geophysically characterize the 

Halema‘uma‘u Overlook eruptive vent area.

Evolution of Volcanic Landforms

Hawaiian eruptions produce stunning examples of 

volcanic landforms and associated features, including 

shields, cones, craters, flow fields, lava lakes, lava tubes, lava 

deltas, and earthquake faults. Since the beginning of HVO, 

scientists have documented well—and often in real time—the 

processes and products resulting in the formation of many 

volcanic landforms. These are recorded and published in 

various venues, ranging from HVO’s early serial publications 

(compiled in Fiske and others, 1987; Bevens and others, 

1988) and annotated bibliographies and pictorial histories 

(Macdonald, 1947; Wright and Takahashi, 1989; Wright 

and others, 1992) to the general scientific literature. In the 

discussion below, we highlight a few selected landforms 

whose construction and evolution were well documented by 

HVO staff. Knowledge gained from studies of these historical 

examples can constrain interpretations regarding the origins 

of similar landforms created in the geologic past in Hawai‘i 

and elsewhere, or on other planetary bodies.

Figure 14. Photographs showing the 

eruption within Halema‘uma‘u Crater 

that began in mid-March 2008. Right: 

Aerial view (looking south) in November 

2008 of the eruption plume from the 

vent within Halema‘uma‘u Crater; the 

Hawaiian Volcano Observatory (HVO) on 

the rim of Kīlauea Caldera is circled. This 

summit activity has continued into 2014. 

Left: Nighttime view in January 2010 of 

the active lava pond in the progressively 

enlarging vent (here ~120 m across). 

USGS photographs by Tim Orr.
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 23

Halema‘uma‘u Crater

The formation of Halema‘uma‘u pit crater within Kīlauea’s 

summit caldera predated the start of scientific observations in the 

19th century. Nonetheless, changes with time in Halema‘uma‘u’s 

configuration and lava-lake activity within it were noted in the 

accounts of the early explorers and settlers. It was recognized that 

the crater occupied the highest area of the caldera floor, which 

represented the surface of the summit of a broad, nearly flat shield.

In response to rapid magma withdrawal preceding the 1924 

explosions, the crater floor dropped about 200 m below the rim; after 

the explosions, it was observed that the rubble-filled crater floor had 

dropped another 200 m (Jaggar and Finch, 1924). The 1924 events 

doubled the width of Halema‘uma‘u to more than 900 m. Lavas 

from post-1924 eruptions at Halema‘uma‘u have refilled much of 

the crater, such that its floor was raised to the current level of about 

85 m below the present-day rim. The Overlook eruptive vent of the 

2008–present summit activity, which now has become the longestlived summit vent since 1924 (Hawaiian Volcano Observatory 

Staff, 2008), is centered within a smaller pit within Halema‘uma‘u 

(fig. 14). The volume of material lost by collapse during the 1924 

explosive eruption was more than 250 times the volume of lithic 

tephra ejected (Jaggar, 1924). A similar relation characterized the 

mid-March 2008 start of the Halema‘uma‘u Overlook eruption, 

in which the estimated volume of the collapse also far exceeded 

the volume of ejecta (Houghton and others, 2011; Swanson and 

others, 2011b). Does this relationship, for two comparatively small 

historical explosive eruptions, also hold for much larger explosive 

events that have occurred in the geologic past?

Mauna Ulu and Kupaianaha Lava Shield Vents

Before the 1983–present Pu‘u ‘Ō‘ō-Kupaianaha eruption, 

the 1969–74 eruption at Mauna Ulu (Hawaiian for “growing 

mountain”) had been the longest-lived historical flank 

eruption—a distinction now held by continuing activity at Pu‘u 

‘Ō‘ō. Episodic overflows from the Mauna Ulu lava lake built 

a lava shield (fig. 15) that reached a maximum height about 

121m above the pre-1969 surface; moreover, overflows from 

the adjacent ‘Alae lava lake constructed a smaller shield, whose 

summit attained a height nearly 90 m above the surrounding 

surface (Holcomb and others, 1974; Swanson and others, 1979; 

Tilling and others, 1987a). Several other lava shields are found 

at Kīlauea along its two rift zones, including Kanenuiohamo 

(dated by radiocarbon analysis at 750–400 years before present; 

Sherrod and others, 2007), Heiheiahulu (~1750 C.E.; Trusdell 

and Moore, 2006), and Maunaiki (1919–20). It was not until 

the 1969–74 Mauna Ulu eruption, however, that modern 

scientific observations could be made to document, in detail, 

the beginning and growth history of a volcanic shield. The 

observations of the processes and durations involved in the 

development of Mauna Ulu and ‘Alae (Swanson and others, 

1979; Tilling and others, 1987a) could be compared with studies 

of a later shield (built during 1986–92), when eruptive activity 

shifted 3 km downrift from the Pu‘u ‘Ō‘ō vent to another site of 

outbreak. Repeated overflows from the new active lava pond fed 

by this new vent (later named Kupaianaha) constructed a shield 

that attained a maximum height of ~58 m above the preeruption 

surface (Heliker and Mattox, 2003).

Mauna Ulu

Pu‘u Huluhulu

‘Alae

Figure 15. Aerial photograph looking west on September 8,1972, of the two volcanic shields (Mauna Ulu in background, and the 

smaller ‘Alae shield in foreground) that developed during the 1969–74 eruption on Kīlauea’s upper East Rift Zone. Both were built 

by repeated overflows from active lava lakes, as can be seen happening here from the perched lava pond at ‘Alae. The liquid lavapond surface of ‘Alae (about 200 m across) indicates scale (USGS photograph by Robert I. Tilling).
24 Characteristics of Hawaiian Volcanoes

The Complex Cone at Pu‘u ‘Ō‘ō

The 1983–present Pu‘u ‘Ō‘ōKupaianaha eruption—the longest lived flank 

eruption at Kīlauea in more than 500 years—

has provided an unprecedented opportunity 

for making detailed observations of the 

evolution of a complex basaltic cone (in other 

words, one not predominantly composed 

of cinder). Growth and collapse of the cone 

during its first 20 years of eruption are 

summarized by Heliker and others (2003a). 

Built from the combined accumulations of 

fountain-fed lava flows, agglutinated spatter, 

and rootless flows, the Pu‘u ‘Ō‘ō cone quickly 

attained its maximum height (255 m) by mid1986. When the Kupaianaha vent formed in 

July 1986, cone growth ceased at Pu‘u ‘Ō‘ō. 

In July 1987, Pu‘u ‘Ō‘ō abruptly collapsed 

to form a steep, narrow, 100-m-deep crater, 

which deepened to 180 m by December 1988 

and was enlarged by piecemeal rock falls 

and collapses of the crater walls. An active 

lava pond was observed intermittently as 

Kupaianaha erupted through February 1992 

(Heliker and others, 2003a). 

With the demise of the Kupaianaha 

vent, the eruption returned to Pu‘u ‘Ō‘ō, 

where, over the next 20 years, the activity 

was dominated by a wide variety of 

processes, including (1) crater floor and wall 

collapses and subsequent filling with lava; 

(2) construction of small (“mini”) shields 

by flank vents that undermined the west and 

south flanks, leading to further cone collapse; 

and (3) episodic crater overflows (Heliker 

and others, 2003a; Poland and others, 2008). 

In 1997, another major collapse took place 

at the summit of the cone (fig. 16), and, by 

late 2002, the cone’s highest point was only 

187m above the preeruption surface. Many of 

the crater floor collapses correspond in time 

to intrusions up or down the rift zone. By 

2011, the central crater was greatly elongated 

along the rift axis, the summit had lowered 

by more than 84 m from its highest elevation, 

and much of the east, west, and south flanks 

of the cone were buried by lava flows. As 

of this writing (mid-2014), the story of Pu‘u 

‘Ō‘ō continues to unfold (Orr and others, 

2012); it is likely that the ongoing eruption at 

Pu‘u ‘Ō‘ō will continue to modify the cone’s 

configuration by similar processes, perhaps 

ultimately leading to its total collapse and 

complete burial by new overflows.

Lava shield

Pu‘u ‘Ō‘ō

1992

Lava shield

West gap

1999

Lava shield

2012

A

B

C

Figure 16. Photographs of the volcanic landforms built at the Pu‘u ‘Ō‘ō vent as they 

evolved over time. A, View of the Pu‘u ‘Ō‘ō cone in June 1992 at its maximum height 

(~255m). An unnamed lava shield that had developed against the cone’s west flank during 

the previous 4 months is visible in the middle ground. People (circled) indicate scale of 

the landforms (USGS photograph by Tari Mattox). B, Same view of the Pu‘u ‘Ō‘ō cone in 

1999 after a major collapse of its summit and west flank, forming the prominent “west gap” 

(USGS photograph by Christina Heliker). C, Same view in 2012 after additional piecemeal 

smaller collapses, continued shield growth, and accumulation of episodic crater overflows. 

The cone has become a much less imposing landmark (USGS photograph by Tim Orr).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 25

Lava Tubes

Lava tubes are common volcanic landforms on Hawaiian 

volcanoes (fig. 17), but the processes of their development were 

not studied “live” until the 1969–74 Mauna Ulu eruption (Swanson 

and others, 1979; Tilling and others, 1987a). Sustained lava 

flows produced during this eruption allowed, for the first time 

historically, systematic and repeated field observations of lava-tube 

formation and evolution. Peterson and others (1994) concluded, 

from observations at Mauna Ulu, that lava tubes can form by four 

principal processes, and that their development was favored by low 

to moderate volume rates of flow for sustained periods of time. 

The subsequent, and even longer-duration, Pu‘u ‘Ō‘ō-Kupaianaha 

eruption (1983–present), provided HVO and other scientists 

unprecedented opportunities to observe active lava tubes and to 

extend and refine the early findings from studies during Mauna Ulu 

(see Heliker and others, 2003b, and references therein). During the 

40-plus years since the Mauna Ulu eruption, the key role played 

by lava tubes in effusive basaltic volcanism has been thoroughly 

investigated in Hawai‘i and documented in numerous summary 

works (for example, Peterson and Swanson, 1974; Swanson and 

others, 1979; Greeley, 1987; Tilling and others, 1987a; Peterson and 

others, 1994; Greeley and others, 1998; Heliker and others, 2003a; 

Helz and others, 2003; Kauahikaua and others, 2003; Cashman and 

Mangan, this volume, chap. 9; and references therein). Later (see 

“Development of Lava Flow Fields” section, below) we focus on 

the importance of lava tubes in the development of lava flow fields.

Perched Lava Ponds and Channels

Perched lava ponds and channels share a similar 

origin—episodic overflows of their banks raise the surface 

of flowing lava progressively higher than the surrounding 

ground (for example, Wilson and Parfitt, 1993). The 

formation and activity of a striking perched channel were 

documented in detail by Patrick and others (2011a). This 

channel formed in 2007 on relatively flat ground within 

1.4km of a new vent downrift of Pu‘u ‘Ō‘ō, and repeated 

lava overflows from it raised the lava surface nearly 

45m above the surrounding terrain at an average rate of 

about 0.3m per day. Integrated monitoring data and field 

observations detailed a process never before seen for a lava 

channel: cyclic lava spattering at intervals of 40–100 minutes 

along the channel margins typically led to sudden lava-level 

drops of ~1 m, accompanied by seismic tremor bursts with 

peak frequencies of 4–5 Hz. This pattern was interpreted 

by Patrick and others (2011a) to reflect a gas pistoning 

process driven by gas accumulation and release beneath 

the lava crust of the active channel. The recent studies at 

Kīlauea (for example, Hawaiian Volcano Observatory Staff, 

2007a,b; Patrick and others, 2011a; Patrick and Orr; 2012b) 

arguably represent the most comprehensive documentation 

to date—replete with detailed chronologies and illustrative 

photographs and diagrams—of the formation of perched lava 

ponds and channels and their evolution.

Figure 17. Photograph of the interior of Thurston Lava Tube (Hawai‘i Volcanoes National Park), looking toward an exit (photograph 

courtesy of Peter Mouginis-Mark, University of Hawai‘i at Mānoa). The tube cavity here is about 4 m high. This is an easily 

accessible, and heavily visited, example of an inactive lava tube; compare with the active lava tube shown being sampled in figure 8.
26 Characteristics of Hawaiian Volcanoes

Rootless Shields and Shatter Rings

The Pu‘u ‘Ō‘ō-Kupaianaha eruption has afforded 

multiple opportunities to study the formation of some 

smaller scale landforms, such as hornitos, elongate tumuli, 

rootless shields, and shatter rings (Kauahikaua and others, 

1998, 2003; Orr, 2011a; Patrick and Orr, 2012b). Unlike 

large lava shields (for example, Mauna Ulu and Kupaianaha) 

that are directly linked to a deep-seated (>1 km) vent, a 

rootless shield is fed by, and built above, an active lava tube. 

It forms at a breakout along the tube, generally associated 

with resumption of active flow within the tube following an 

eruptive pause (Kauahikaua and others, 2003). A perched 

lava pond quickly develops above the breakout point, and 

overflows from that in all directions construct a shieldlike structure. First seen in 1999, such features have been 

observed several times since then, most notably in 2002, 

2003, and during 2007–11, and can build as high as 20m 

above the surrounding surface. The largest one, to date 

(formed in 1999), had a diameter of more than 500 m and 

was topped by a flat, lava-ponded surface 175 m across 

(Kauahikaua and others, 2003). After becoming inactive, 

rootless shields collapse and reveal hollow interiors, 

suggesting that they contained shallow reservoirs filled with 

lava before draining. This hypothesis was confirmed several 

times when the side of some actively growing shields during 

2007–08 collapsed and produced fast-moving lava flows as 

they drained (Patrick and Orr, 2012b).

First noticed in the 1990s (Kauahikaua and others, 

1998, 2003), shatter rings also can form over active lava 

tubes carrying variable amounts of lava that alternately push 

the tube roof up when the tube is full and then let it down as 

the lava flux decreases. The repeated up-and-down flexing 

of the tube roof eventually fractures the roof in an oval 

shape defined by a ring of rubble with a fractured pāhoehoe 

surface in the middle (Kauahikaua, 2003, figs. 6–8). Actively 

deforming shatter rings sound like rocks constantly grinding 

against each other. Lava in the tube can break out to the 

surface through the rubble rings, typically while the roof is 

being pushed upward. Orr (2011a) documents the shatterring process and provides a comprehensive list of other 

flow fields in the world that have shatter rings. He suggests 

that recognition of shatter rings on other planets would be 

evidence that lava tube systems were once active there.

Lava Deltas

The prolonged Mauna Ulu and Pu‘u ‘Ō‘ō-Kupaianaha 

eruptions sent many tube-fed pāhoehoe lava flows into the 

ocean, thereby affording HVO scientists ample opportunity 

to track the detailed growth and collapse of lava deltas and 

to study accompanying explosive hydrovolcanic activity 

(for example, Peterson, 1976; Mattox and Mangan, 1997; 

Kauahikaua and others, 2003; Sansone and Smith, 2006). 

One of the main factors that promotes or limits delta growth 

is the steepness of the hydroclast-mantled submarine slope, 

which determines its ability to support the seaward extension 

of the leading edge of the overlying lava flows. Tube-fed lava 

flows have entered the ocean about 70 percent of the time 

during 1986–2010; the two longest duration ocean entries 

each lasted 22 months (Laeapuki in May 2005–March 2007 

and Waikupanaha in March 2008–January 2010; Tilling 

and others, 2010). Upon entry into the ocean, the hot lava is 

shattered by coming into contact with cold seawater, can be 

reworked by surf action, and is easily erodible. If lava entry is 

sustained, however, the fragmental debris can accumulate in 

the littoral zone, forming a bulwark of new land—enlarging 

and extending the shoreline seaward—over which additional 

lava advances to form lava deltas. 

Recent studies have documented the rapid changes 

that lava deltas can undergo after their formation. As an 

example, the east Laeapuki lava delta had grown to about 

44 acres in size by mid-November 2005 and then abruptly 

collapsed into the sea on November 28 to produce the 

largest lava-delta collapse observed to date (the collapse 

was captured on time-lapse video; see Orr, 2011b). With 

continued lava entry at the same location, the delta quickly 

rebuilt (fig. 18) to reach 64 acres in size by March 2007. 

Then, within a few months, the delta began a series of 

collapses, such that by June 2010, none of the originally 

huge east Laeapuki lava delta remained (Tilling and others, 

2010). We present this example to emphasize that, without 

regular direct observations, the complete history of the 

dynamic but short-lived evolution of this lava delta would 

have been difficult, if not impossible, to reconstruct from 

the now-available field evidence. Current studies of lava 

deltas by HVO provide potentially useful information (for 

example, vent effusion rate, mode of lava transport to ocean, 

duration of ocean entry) for deciphering the formation and 

histories of prehistoric lava deltas in Hawai‘i or at other 

island volcanoes (for instance, in the Canaries or the Azores, 

Stromboli, Iceland, and the Aleutians). 

Volcano Experiments and Topical 

Studies

The initiation of regular observation and measurement 

of volcanoes, while rooted in the early 20th century 

movement to establish volcano observatories, also has 

facilitated research that has transformed volcanology into 

the multidisciplinary science that it is today. With its high 

eruption frequency and relatively easy and safe accessibility 

to an ample supply of molten or freshly erupted solidified 

lava, Kīlauea is a natural laboratory for conducting 

experimental or specialized studies. Apple (1987) gives 

a highly readable summary about some of the earliest 

volcano “experiments”—mostly involving temperature 

measurements—conducted by Jaggar and his collaborators. 

Because of HVO’s isolated island location, these early 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 27

experiments necessarily depended on materials, equipment, 

and techniques that were available, or could be readily 

adapted, in Hawai‘i. In this section, we will revisit some 

topically driven studies undertaken during the past century 

with the aim of better understanding basaltic volcanism. The 

scientific literature contains results of these topically focused 

studies conducted by HVO and other scientists during the 

past century; here, we present only selected highlights.

Field Measurements of Lava Temperature

Even before HVO was formally established, a field 

experiment was undertaken in 1911 by Frank Perret and 

E.S. Shepherd to measure directly the temperature of the 

active lava lake at Halema‘uma‘u. The technique involved 

lowering a “special resistance thermometer” (a platinum 

thermocouple in an iron casing) into the molten lava lake 

by means of a specially designed cable-trolley system 

built across the lava lake (fig. 19A). After many logistical 

difficulties with the cable system and the loss of two 

thermometers, a reading of 1,010 °C was registered from 

the still-working third and last thermometer immersed 

about 0.6 m under the lava-lake surface (see interesting 

accounts by Perret, 1913f; Jaggar, 1917a; and Apple, 1987). 

Subsequent studies have shown that this temperature is 

too low for liquid basalt, reflecting chemical reactions of 

volcanic gases with the thermocouple elements that affected 

calibration. Nonetheless, the Perret-Shepherd experiment 

represented a historic scientific achievement in its time—

the first in-place temperature measurement and sampling 

(fig. 19B) of liquid lava ever made at any volcano. 

Temperature estimates of liquid lava were also made 

by thrusting steel pipes down to different depths below 

the surface of the active lake. At the end of the pipe was a 

cylinder containing six Seger cones, which are mixtures of 

clay and salt with variable known melting points used by 

ceramicists as temperature indicators. The pipe was inserted 

to a known depth and held in place for 6 minutes and then 

withdrawn—with pasty lava adhered to the part of the pipe 

that penetrated into still-molten material. Then, each of the 

Seger cones was examined to ascertain which had melted 

and which had not, thereby determining the thermal gradient 

within the examined depth interval (~1.5 m), assuming 

thermal equilibrium was attained (Jaggar, 1917a; Apple, 

1987, fig. 61.5). The thermal gradients obtained by HVO 

in 1917 doubtless were imprecise by today’s standards, 

but they nevertheless constituted the first and only attempt 

ever made to determine the shallow thermal gradient for an 

active lava lake. 

Because of the logistical difficulties attendant upon 

these earliest measurements of the temperature of liquid 

lava, HVO later relied on more robust thermocouples, 

optical pyrometers, and other hand-held instruments for field 

measurements of lava temperature. As geothermometric 

techniques became better calibrated, they have been 

increasingly used to determine lava temperatures of glassy 

water-quenched samples collected during the course of 

an eruption (for example, Helz and others, 1995, this 

volume, chap. 6), obviating the need for direct temperature 

measurements in the field.

In 1922, HVO conducted the first scientific drilling 

studies in solidified lava, using both churn-drill and rotarydrill rigs and imaginative adaptations of automobiles 

to haul drilling water and supplies to drill sites. Four 

holes were completed during this experiment, ranging in 

A

B

Figure 18. Oblique aerial photographs of lava-delta growth where 

lava enters the sea. A, Looking towards the west along Kīlauea’s south 

coast in early December 2005, showing the arcuate, newly exposed 

sea cliff four days after the collapse of the actively growing lava delta 

(approximated by dashed line) at east Laeapuki (USGS photograph by 

James Kauahikaua). B, With continuing lava entry into the sea, by late 

April 2006 the lava delta was rebuilt and larger in size than before the 

collapse (USGS photograph by Tim Orr).
28 Characteristics of Hawaiian Volcanoes

depth from 7 to 15 m, for temperature measurements in 

solidified lava at several sites at Kīlauea’s summit. None 

of the measured bottom-hole temperatures exceeded 97 °C 

(Jaggar, 1922). Later, Jaggar planned to have a summit 

network of shallow (3 m) boreholes, intended for periodic 

remeasurements to collect long-term data to track temporal 

changes in the temperature of the bedrock. After drilling 30 

holes for this network, however, the effort was abandoned 

in 1928, presumably because the results were inconclusive 

(Apple, 1987). While these early drilling experiments 

and the resulting temperature measurements produced 

few lasting scientific results, they provided an important 

conceptual framework for subsequent modern scientific 

drilling studies at Kīlauea. 

Drilling Studies of Passive Historical Lava Lakes

Beginning in 1960, HVO conducted pioneering drilling 

studies of three historical passive lava lakes: Kīlauea Iki 

(erupted in 1959), ‘Alae (erupted in 1963), and the west pit of 

Makaopuhi (erupted in 1965). Of these, only the Kīlauea Iki 

lava lake has not been buried by younger lavas. 

Five holes were drilled into the solidifying crust 

of Kīlauea Iki between April 1960 and December 1962 

(Rawson, 1960; Ault and others, 1961; Richter and Moore, 

1966). On July 25, 1960, the hole jointly drilled by the 

Lawrence Radiation Laboratory and HVO was the first to 

penetrate the crust, marking the first-ever drilling experiment 

specifically designed to directly measure temperature 

and sample the molten lava and associated volcanic gas. 

Owing to the crude and improvised equipment used, the 

first drillings at Kīlauea Iki encountered problems in the 

high-temperature environment. The early lessons learned 

regarding drilling techniques, however, were later applied to 

successful subsequent drilling studies at ‘Alae, Makaopuhi, 

and Kīlauea Iki (fig. 20). Because it was the largest and 

has remained accessible for study, the Kīlauea Iki lava lake 

was drilled repeatedly (in 1967, 1975, 1976, 1979, 1981, 

and 1988) to document its prolonged cooling history (Helz 

and Taggart, 2010; Helz and others, this volume, chap. 6). 

On the basis of its thermal history, Kīlauea Iki probably 

fully crystallized by 1995. The samples and temperature 

data collected from these drilling studies have provided 

significant insights into the cooling, crystallization, and 

differentiation of basaltic magma (for example, Peck and 

others, 1966; Richter and Moore, 1966; Wright and others, 

1976; Wright and Okamura, 1977; Peck, 1978; Wright and 

Helz, 1987; Helz, 1987a,b, 2009; Jellinek and Kerr, 2001; 

Helz and others, this volume, chap. 6).

In-Place Measurements of Molten Lava Properties

Boreholes drilled into cooling lava lakes provided 

entryways for instruments and other experimental devices for 

making measurements from the solidified surface at ambient 

temperature, down to the still-molten lava at magmatic 

temperature at the bottom of the hole. As examples, we 

briefly comment on two unique experiments: one conducted 

in the 1965 lava lake within Makaopuhi Crater (Shaw 

and others, 1968) to measure lava viscosity and another 

in the 1959 Kīlauea Iki lava lake to measure magnetic 

susceptibility (Zablocki and Tilling, 1976).

Directly measuring viscosity involved the insertion 

of a stainless steel spindle into the melt via a casing that 

was forced through the crust, which was ~4.4 m thick at 

the time of the measurements. The spindle was rotated by 

application of known torque using a system of weights 

and pulleys (Shaw and others, 1968, figs. 2 and 3), and the 

number of revolutions of the rotating spindle was measured 

by means of a stopwatch. The entire operation to obtain a 

A

B

Figure 19. Photographs recording early adventures in studying and 

sampling molten lava at Kilauea. A, The cable-trolley system used 

by Frank Perret in 1911 for temperature measurement and sampling 

of lava from the active Halema‘uma‘u lava lake. (Inset sketch and 

photograph are from Perret, 1913f, figs. 1 and 2, respectively.) B, 

Perret (right) and an assistant carry a sample of quenched molten lava 

collected by means of a pot and chain attached to the spanning cable 

of the apparatus shown in A (taken from Perret, 1913f, fig. 3). Inset 

(lower right corner) is a portrait of Perret in 1909 (photograph from 

Library of Congress).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 29

measurement—drilling, emplacement of the spindle, time 

required for readings, and withdrawal of the spindle—needed 

to be completed in one day, because if the hole was left open 

overnight, it became impossible to insert the spindle the 

next day due to the “hardening” of the crust-melt interface 

zone, and if the spindle was left too long in the hole, it 

would become encased in lava and could not be recovered. 

The simple but functional viscometer invented for use at 

Makaopuhi made possible the first-ever field measurement 

through a borehole of the viscosity of tholeiitic basaltic 

melt. The field data, combined with laboratory data to higher 

temperatures, indicated a Newtonian viscosity of about 4,000 

poises for the equilibrium phase at 1,130 °C.

During the February–March 1975 drilling at Kīlauea 

Iki, two of the three holes were used to measure magnetic 

susceptibility (Zablocki and Tilling, 1976). As with the field 

measurements of viscosity at Makaopuhi, the situation at 

Kīlauea Iki also required some special adaptations. Because 

of the high temperatures (>900 °C) in the holes, the in-hole 

sensing element had to be thermally insulated to maintain 

a constant 100 °C; this was achieved by placing the sensor 

in a water-filled, open-ended ceramic tube wrapped in thin 

asbestos sheets. Moreover, because the depths of the 1975 drill 

holes (~44 m) were much deeper than those in the pre-1975 

lava-lake drillings, it was necessary to devise a special flexible 

thermocouple system that was manually lowered and raised 

in the hole by means of a sheave to measure the temperature 

profile. For details about both the magnetic susceptibility 

and the thermocouple probe, see Zablocki and Tilling 

(1976, especially fig. 2). Measurements made in June 1975 

showed that the minimum temperature at which magnetic 

susceptibility drops to zero is about 540°C—a temperature 

that accords well with the range of Curie temperatures 

obtained in laboratory heating experiments of basaltic core 

samples. This particular special study at Kīlauea Iki represents 

“the first in-situ, ‘real time,’ determination of the apparent 

Curie temperatures of cooling basaltic lava.” (Zablocki and 

Tilling, 1976, p. 487).

The slowly cooling molten lava body within Kīlauea 

Iki Crater also provided an ideal setting for geophysical 

studies to determine its electrical properties. Anderson 

(1987) reported on the electrical structure of the crust 

above the molten lava, and Frischknecht (1967) used 

surface electromagnetic soundings to determine molten 

lava conductivity to be about 2 ohm-meters. Various 

electromagnetic and magnetic techniques were used to map 

the edges of the molten lava and document the contraction of 

the cooling edge with time (Smith and others, 1977).

A B

Figure 20. Photographs 

of drilling operations at 

Kīlauea Iki lava lake in 

1975. A, General view of 

the drill rig. It and other 

heavy equipment were 

lowered into the crater by 

helicopter. B, The top of 

drill hole 75–2 and view 

looking down the drill 

hole (~ 6 cm in diameter) 

from the surface to hole 

bottom (~40 m), where 

still-molten 1959 lava 

glows at ~1,110 °C. (USGS 

photographs by Robin T. 

Holcomb).
30 Characteristics of Hawaiian Volcanoes

Underwater Observations of Active Lava Flows

Fluid lava erupted or flowing under water can produce 

“pillow lavas,” which have been studied and mapped in many 

submarine volcanic environments, but their actual formation 

had never been directly observed in real time anywhere in 

the world before 1970. During lava entries into the ocean 

associated with long-lived flank eruptions, such as the Mauna 

Ulu (1969–74) and Pu‘u ‘Ō‘ō-Kupaianaha (1983–present) 

activity, scuba-diving observers were able to watch and film 

(fig. 21) the formation of pillow lava at the submerged fronts 

of actively advancing flows (for example, Moore and others, 

1973; Moore, 1975; Tepley, 1975; Tribble, 1991; Sansone 

and Smith, 2006). These real-time observations of pillowlava formation have generally confirmed the origin and 

emplacement mechanisms inferred from studies of ancient 

pillow lavas in Hawai‘i and elsewhere. Recent submarine 

studies (Takahashi and others, 2002) confirm Moore and 

Fiske’s (1969) assertion that the bulk of the below-sea part 

of a Hawaiian volcano is composed of pillow lavas. In 

addition to pillow-lava formation, Tribble (1991) also, for 

the first time, witnessed live the development of submarine 

channelized flows and submarine slope failures.

Lava-Seawater Interactions and Ocean-Entry Gas 

Plumes

The decades-long duration of the 1983-to-present 

eruption of Kīlauea has afforded ample opportunity to 

study lava-seawater interactions. Direct observations of 

the hydrovolcanic explosive activity produced upon lava 

entry into the ocean have improved our understanding of 

the processes that produce black sand beaches and littoral 

cones, common features along Hawai‘i’s coasts (for instance, 

Moore and Ault, 1965; Jurado-Chichay and others, 1996). 

Mattox and Mangan (1997) documented the conditions under 

which lava entered the ocean explosively and found that it 

required relatively high rates of lava flux into the water. 

When molten lava enters the ocean, the ensuing 

reaction with chlorides in the seawater can produce highly 

acidic steam plumes, sometimes called “laze” (lava haze). 

The plumes have been found to contain hazardously high 

concentrations of hydrochloric acid (pH 1.5–2), after Hawai‘i 

Volcanoes National Park personnel complained about their 

frequent exposure to laze while on duty in the lava-entry 

areas (Sutton and others, 1997; Johnson and others, 2000; 

Sansone and others, 2002; Edmonds and Gerlach, 2006).

Figure 21. Photograph of scuba divers filming an underwater lava flow during the Pu‘u ‘Ō‘ō-Kupaianaha 

eruption—the first such filming was done earlier, during the 1969–74 Mauna Ulu eruption. In this image, the crust 

of chilled lava is breaking open to reveal the advancing bright-orange molten lava inside (copyrighted photograph 

courtesy of Sharkbait World Pictures, Kailua-Kona, Hawai‘i).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 31

Use of Digital Cameras, Webcams, and Video to 

Precisely Document Eruptive Processes

Throughout HVO’s history, sketches, photographs, and 

movies were the mainstays for documenting eruptive activity at 

Kīlauea and Mauna Loa, and time-lapse photography was first 

used to record lava-lake activity during the 1969–74 Mauna 

Ulu eruption. Beginning in 2004, webcams, time-lapse digital 

photography, and forward-looking infrared (FLIR) images have 

come into wide and increasing use (see, for example, Orr and 

Hoblitt, 2008; Hoblitt and others, 2008, Orr and Patrick, 2012; 

Orr and Rea, 2012; Patrick and Orr, 2012a; Patrick and others, 

2014; Orr, 2014). These techniques have allowed volcanic 

processes to be “seen” and compared in near real-time with 

geophysical monitoring data as never before. Additionally, as 

emphasized by Patrick and others (2014), thermal cameras 

sometimes can also “see” through thick volcanic fume, thereby 

providing views of eruptive activity not possible by means of 

visual webcams and the naked eye. Figure 22 compares visual, 

FLIR, and composite images of lava flows. An added bonus 

is that direct views of Hawaiian eruptive activity can now be 

shared with the world via the Internet. For example, from the 

HVO Web page, it is possible to access the live views from 

webcams (including one at Mauna Loa’s summit), as well as 

to view time-lapse digital photography of different types of 

eruptive activity at various sites (Orr, 2011b). In part from 

analysis of these new data streams, Patrick and others (2011b) 

used detailed timing of Halema‘uma‘u’s degassing pulses 

related to tremor bursts and small explosive events at or near 

the top of the lava column to show that they produced verylong-period signals at about 1-km depth.

Development of Lava Flow Fields

During the prolonged Mauna Ulu (1969–74) and Pu‘u 

‘Ō‘ō-Kupaianaha (1983–present) eruptions, extensive lava flow 

fields developed by the emplacement of multiple fluid flows—

fed via long-duration surface channels or lava tubes—traveling 

great distances from the source vents to the coastal plains. The 

mechanisms and dynamics of lava flow emplacement are treated 

thoroughly by Kauahikaua and others (2003) and by Cashman 

and Mangan (this volume, chap. 9, and references therein), but 

we reemphasize below some processes well studied in Hawai‘i. 

Systematic observations of flow dynamics, combined with 

detailed mapping of flow advances during the growth of lavaflow fields at Kīlauea,

provide comprehensive case histories for 

comparing and contrasting the development of complex lavaflow fields—prehistoric and historical—at other volcanoes (for 

example, Calvari and Pinkerton, 1998; Guest and others, 2012; 

Solana, 2012; Branca and others, 2013).

Pāhoehoe-‘A‘ā Transition

The distinctive differences in appearance between 

pāhoehoe and ‘a‘ā lavas naturally drew the attention of 

early observers of Hawaiian eruptions. They recognized 

that these two contrasting types of lava could occur during 

the same eruption, and even along a single active flow; 

thus, their origin could not be explained by any inherent 

differences in the magma before eruption. As annotated by 

Wright and Takahashi (1989, p. 6), the earliest published 

account of the origin of ‘a‘ā was by W.D. Alexander (1859), 

who, while observing the 1859 eruption of Mauna Loa, 

likened ‘a‘ā formation to “graining” of sugar. Alexander’s 

notion of “sugaring” was later followed up scientifically by 

Jaggar (1947) and Macdonald (1953b), who hypothesized 

that ‘a‘ā formation involved volcanic degassing and change 

of viscosity.

Scientific interest in the pāhoehoe-‘a‘ā transition 

was rekindled, beginning with the 1969–74 Mauna Ulu 

eruption. During this eruption, many long lava flows over 

gentle and steep slopes afforded abundant opportunities to 

observe, in real time, transitions between the two types. 

Peterson and Tilling (1980) proposed a semiquantitative 

model relating viscosity and rate of shear that featured a 

“transition threshold zone” (TTZ) separating the pāhoehoe 

and ‘a‘ā flow regimes. During flow, once a discrete 

infinitesimal element of lava crosses the TTZ, pāhoehoe 

changes to ‘a‘ā (Peterson and Tilling, 1980, fig. 9). The 

original Peterson-Tilling model has been substantially 

modified and quantified by many subsequent studies (for 

example, Kilburn, 1981, 1993, 2000; Lipman and Banks, 

1987; Wolfe and others, 1988; Cashman and others, 1999). 

Some of the most recent studies (for instance, Hon and 

others, 2003; Kauahikaua and others, 2003) argue that 

the pāhoehoe-‘a‘ā transition is not always irreversible, as 

originally contended by Peterson and Tilling (1980); ‘a‘āpāhoehoe transitions do indeed occur, as shown during the 

Pu‘u ‘Ō‘ō-Kupaianaha eruption. Considerable scientific 

debate continues about the rheological nature and styles 

of transitions between ‘a‘ā and pāhoehoe during a single 

active flow (see Cashman and Mangan, this volume, 

chap. 9). Most eruptions at Kīlauea and Mauna Loa produce 

‘a‘ā flows initially but quickly transition to pāhoehoe 

flows for the remainder of activity. The ratio of duration in 

these phases historically has ranged from 1:7 (Mauna Ulu 

eruption) to 1:18 (1859 Mauna Loa eruption). Using these 

historical ratios, staff member Jack Lockwood speculated 

in an internal HVO communication in early 1987: “If 

historical ‘a‘ā/pāhoehoe chronological ratios of previous 

long-lived Hawaiian flank eruptions persist, the sustained 

pahoehoe production eruption which began in July 1986 

(“Phase 48”) could last from 24 to 63 years! This is, of 

course, outrageous; we know of no historical example 

of such long-lived activity—but is it really impossible?” 

(Lockwood, 1992, p. 6). Lockwood’s line of reasoning was 

met with considerable skepticism at the time, but now—

in hindsight—seems remarkably prescient and entirely 

plausible as we enter the 31st year of near-constant lava 

effusion on Kīlauea’s East Rift Zone (Orr and others, 

2012).
32 Characteristics of Hawaiian Volcanoes

County lava flow

viewing area

B

A

Figure 22. Images illustrating the use of forward-looking infrared (FLIR) imaging to monitor active lava flows. A, Comparison of conventional 

photographic image (left) of active lava flow from Pu‘u ‘Ō‘ō with an FLIR image of the same flow (right), with the yellow-orange colors showing 

more clearly its hottest (most active) parts (USGS images, March 14, 2008). B, Example of composite conventional photograph and FLIR image 

of active lava flow from Pu‘u ‘Ō‘ō at ocean entry; such composite images have been popular with emergency-management officials and the 

general public (USGS image, July 30, 2010).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 33

Lava-Tube Systems

With sustained effusive eruptive activity, well-integrated 

lava-tube systems commonly develop during the emplacement 

of lava flow fields. These systems thermally insulate lava 

flowing within the tubes, thereby allowing the lava to travel 

great distances and ultimately enter the ocean. The ability of 

fluid lava to flow great distances in this way contributes to the 

low-angle slopes that characterize Hawaiian and other shield 

volcanoes. HVO studies of lava-tube behavior have largely 

settled an oft-debated question: can the molten lava flowing 

through tubes erode down through its base? The answer is “yes,” 

by thermal and (or) mechanical processes, and the erosion rate 

has been observed to be as high as 10 cm/day for several months 

(Kauahikaua and others, 1998a, 2003). The key role played 

by the tubes in facilitating long-distance lava transport and in 

emplacing flow fields has been thoroughly studied in Hawai‘i, 

and we refer the interested reader to the numerous summary 

works (for example, Peterson and Swanson, 1974; Swanson and 

others, 1979; Greeley, 1987; Tilling and others, 1987a; Peterson 

and others, 1994; Heliker and others, 2003b; Helz and others, 

2003; Kauahikaua and others, 2003; Cashman and Mangan, this 

volume, chap. 9; and references therein). 

Inflation of Pāhoehoe Flows

Inflation features in pāhoehoe flows were first recognized 

during the 1919–20 Maunaiki eruption and described by Finch 

and Powers (1920, p. 41) as “A large schollendome [equant 

tumulus] . . . had been built up . . . occupied by recently 

solidified lava that had risen from below as was shown by its 

flat topped filling.” Prominent tumuli were also recognized on 

the floor of Kīlauea Caldera but were attributed to the action of 

gas pressure. This endogenous mechanism in the development 

of pāhoehoe flow fields was recognized by Holcomb (1987) 

while mapping recent Kīlauea lava flows; he labeled them 

“inflated.” The dominant mode of emplacement of pāhoehoe 

flows entering the coastal community of Kalapana on 

Kīlauea’s south flank in 1990 also involved endogenous 

growth or “inflation.” Molten lava that flows or spreads into 

solidifying crust at a flow front lifts its surface at decreasing 

exponential rates (for example, Walker, 1991; Hon and 

others, 1994; Cashman and Kauahikaua, 1997; Kauahikaua 

and others, 1998, 2003; Hoblitt and others, 2012). Once this 

mechanism was understood, the hazards posed by pāhoehoe 

flows could be better estimated using the rate of increasing 

height of inflating lava flows rather than only the proximity 

of infrastructure to recently active lava flows. The extensive 

documentation of the flow-inflation process at Kīlauea has 

contributed directly to studies of the processes and duration 

of the emplacement of continental flood basalts (for example, 

Columbia River Basalt, Deccan Traps), other hot-spot shield 

volcanoes (for example, Iceland), and the submarine basaltic 

flow fields that make up the ocean floor (for instance, Hon and 

others, 1994; Self and others, 1998; Thordason and Self, 1998; 

Sheth, 2006; and references therein).

Deep Scientific Drilling Studies

Beginning in the 1970s, several deep (>1 km) holes 

were drilled on the Island of Hawai‘i to learn more about the 

deeper parts of volcanic edifices and magmatic/hydrothermal 

systems not accessible to direct surface-based studies. These 

drilling studies, which were funded by the National Science 

Foundation (NSF), were led by investigators in academia, 

but HVO and other USGS scientists participated, directly or 

indirectly, in the acquisition of in-hole data, analysis of core 

samples, and interpretations of the findings.

Drilling into Kīlauea Caldera

In collaboration with HVO, the Colorado School of 

Mines secured NSF funding for the purpose of drilling a 

research borehole into the inferred hydrothermal convection 

cell above the shallow magma reservoir beneath the 

summit of Kīlauea Volcano. A preliminary time-domain 

electromagnetic survey of the summit caldera and areas 

over the summit magma reservoir identified a likely drilling 

target in the form of a shallow high-conductivity anomaly 

in the southern portion of the caldera (Jackson and Keller, 

1972). Drilling during April–July 1973 reached a depth of 

1,262 m and encountered elevated temperatures (maximum 

137 °C at hole bottom) but no magma (Zablocki and others, 

1974; Keller and others, 1979). The drill hole did confirm 

the existence of a shallow groundwater table about 500 m 

below the ground surface (700 m above sea level). The 

complicated nature of the hole’s thermal profile (Zablocki 

and others, 1974, fig. 3) reflects hydrothermal circulation, 

with a largely convective regime in the depth interval 500–

950 m and a largely conductive regime beneath that.

The summit drill hole was left open to allow further 

studies of the water table, including the episodic collection 

of water samples during 1973–76 (McMurtry and others, 

1977; Tilling and Jones, 1995, 1996), in 1991 (Janik 

and others, 1994), and during1998–2002 (Hurwitz and 

others, 2002, 2003; Hurwitz and Johnston, 2003) for 

laboratory analysis. These samples were unique in that they 

constituted the only analyzed samples of thermal water 

from directly above Kīlauea’s summit magma reservoir. 

Tilling and Jones (1995, 1996) were the first to report on 

the water chemistry of these samples and noted temporal 

changes in composition related to possible rainfall dilution 

and to increased partial pressure of CO2 related to volcanic 

degassing accompanying the December 31, 1974, eruption. 

The well was cleaned out in 1998 and the sampling 

resumed during 1998–2002. Analysis of waters showed 

continued temporal compositional variations (fig. 23), as 

well as a change in water level interpreted as response 

to a nearby magma intrusion. For detailed discussions of 

the temporal changes in water chemistry, well level, and 

borehole temperatures through 2002, see Hurwitz and 

others (2002, 2003) and Hurwitz and Johnston (2003).
34 Characteristics of Hawaiian Volcanoes

Hawai‘i Scientific Drilling Project (HSDP)

The primary justification for this NSF-funded project 

was to better characterize and understand the mantle plume 

inferred to sustain the Hawaiian hot spot within the context of 

plate tectonics (Stolper and others, 1996a; for details about the 

HSDP, see http://hawaii.icdp-online.org). As emphasized by 

Stolper and others (1996b, p. 11593), “Hawaii was a natural 

target since as the best studied volcanic construct on Earth, it 

is the archetype of ocean island volcanism and provides the 

best possible scientific framework for a major project of this 

sort.” We fully concur with this rationale. Two holes have been 

drilled as part of the HSDP—a 1,052-m-deep “pilot hole” 

(HSDP1) completed in 1993, and the HSDP2 hole, drilled to 

a depth of 3,110 m in 1999 and later, during 2004–07, to a 

depth of 3,508 m. In all, 4,600 m of rock core were collected 

(mostly from Mauna Kea volcano), spanning in geologic age 

from perhaps 700 ka to ~200 ka. Taking into account plate 

motion during this interval, the HSDP cores represent “the 

first systematic cross-sectional sampling of a deep mantle 

plume” (Stolper and others, 2009, p. 13). In-hole observations 

and measurements made during the drilling, together with 

petrologic, geochemical-isotopic, and geochronological 

studies of the core samples, have yielded unprecedented 

data on the internal structure, hydrogeologic regime, and 

evolution of a large ocean-island volcano. HSDP studies 

(still continuing) have contributed significantly to knowledge 

about intraplate magmatism, mantle plume dynamics, volcano 

growth and subsidence, and other key aspects of Hawai‘i’s 

geologic evolution (see, for example, Beeson and others, 

1996; DePaolo and Stolper, 1996; DePaolo and others, 1996, 

2001; Stolper and others, 1996a,b, 2009; Garcia and others, 

2007; Jourdan and others, 2012).

Geothermal Development in Hawai‘i

National interest in developing alternative energy 

sources, including geothermal, was spurred by the 1973 oil 

crisis that adversely impacted daily life in the United States. 

Although geothermal exploration in Hawai‘i had begun as 

early as 1961 (Macdonald, 1973), it greatly accelerated in the 

1970s with the inception of the Hawai‘i Geothermal Project 

(HGP) in 1973 (Thomas, 1990; Moore and Kauahikaua, 

1993). The early geothermal studies in Hawai‘i were led by 

the University of Hawai‘i at Mānoa (UHM), but many of them 

were done in collaboration with HVO scientists. The statewide 

inventory of geothermal resources (Thomas, 1984, 1986) 

incorporated several USGS- and UHM-led studies defining 

potential resources in Hawai‘i. The high-temperature resources 

were generally within the rift zones of active volcanoes on 

the Island of Hawai‘i, while the lower temperature resources 

were distributed on the older islands. In 1976, as part of the 

HGP, a public-private partnership developed Hawai‘iʼs first 

geothermal well (HGP-A), in Kīlauea’s lower East Rift Zone 

(Thomas, 1987, figure 56.1). With the installation in 1981 of 

a wellhead generator on this well, the partnership operated an 

experimental 3-megawatt power plant during 1982–90, before 

being replaced in 1990 by the commercial Puna Geothermal 

Venture (Boyd and others, 2002). Additional exploratory wells 

have been drilled in Puna, and at present about 30 megawatts 

of electricity (MWe) are being produced and fed into the local 

utility grid (Ormat Technologies, Inc., 2012). 

To better understand the geochemical and structural 

conditions of Kīlauea’s East Rift Zone (ERZ), and to monitor 

changes in the hydrothermal system, three cored scientific 

observation holes (SOH 1, SOH 2, and SOH 4), ranging in 

depth from 1.7 to 2.0 km, were drilled between 1989 and 

Mg

Ca Na+K

Basalt 1991–1993

1976

1975

1974

1973

1998–2002

Sea

water

HGP-A

KW00-WH

50% 50%

50%

Figure 23. Triangular diagram showing the cation 

composition of water samples from Kīlauea. Plotted 

are relative proportions of calcium (Ca), magnesium 

(Mg), and sodium (Na) and potassium (K). Circles and 

ovals are waters collected from the deep drill hole at 

the summit of Kīlauea between 1973 and 2002 (from 

Hurwitz and others, 2003, fig. 5). Solid dot (KW00-WH) 

is water from cistern at wellhead, solid squares represent 

seawater and water from geothermal well HGP-A on 

Kīlauea’s lower East Rift Zone (Thomas, 1987), and 

the hexagon (basalt) is bulk average of Kīlauea summit 

basalt. Significance of these compositional and temporal 

variations is discussed by Tilling and Jones (1996) and 

by Hurwitz and others (2003).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 35

1991 into the lower ERZ (Olson and Deymonaz, 1992; 

Olson and others, 1990). Studies of core from these holes 

have characterized the hydrothermal alteration mineralogy 

of the ERZ geothermal area (Bargar and others, 1996) and 

contributed to an improved understanding of the magmatic 

history of the ERZ (Quane and others, 2000). The routine 

commercial drilling of an injection well in 2005 in the Puna 

geothermal field penetrated dacitic melt (~1,050 °C) at a 

depth of 2,488 m, marking one of the very rare instances 

when molten material has been encountered accidentally in 

geothermal drilling (Teplow and others, 2009).

The potential of the geothermal resource at Kīlauea 

could be as great as 500 to 700 MWe (Olson and others, 

1990)—but can the full exploitation of this significant 

resource be balanced with the high threat of volcanic hazards 

in Hawai‘i (Hawaiian Volcano Observatory Staff, 2012a)? 

With the initiation of exploitation of geothermal resources, 

HVO scientists have cautioned that, in developing the hightemperature resources, it is necessary to recognize the inherent 

hazards posed by frequent eruptions in Hawai‘i, especially 

on Kīlauea’s ERZ (see, for example, Moore and Kauahikaua, 

1993; Moore and others, 1993; Kauahikaua and others, 1994).

Cooperative Research with Other 

Organizations

Because they are frequently active, relatively accessible, 

and generally safely approachable, Hawaiian volcanoes 

always have been attractive research targets, and consequently 

their study has produced thousands of publications. Many 

of the scientific studies involve only HVO and allied USGS 

researchers, but the overwhelming majority of the publications 

are products of collaborative work between HVO/USGS 

personnel and scientists of other organizations—universities 

and other research centers, national and international. This 

collaboration would come as no surprise to Jaggar, who, from 

the very beginning, advocated and supported collaborative 

scientific research: “A volcano observatory must see or 

measure the whole volcano inside and out with all of science 

to help” (Jaggar, 1941). By “all,” Jaggar clearly had in mind 

all fields of science, and all scientists whose expertise and 

work would contribute to an improved understanding of 

how volcanoes and earthquakes work. While Jaggar was a 

visionary scientific thinker, he also had common sense and 

realized that HVO would need all the help it could get to fulfill 

its scientific vision. It is far beyond the scope of this paper 

to detail HVO’s rich history of collaborative research during 

the past 100 years. Nonetheless, in the discussion to follow, 

we highlight a few examples (mostly since the 1950s) of joint 

research with scientists of other organizations.

During HVO’s early years, HVO collaborated with, 

and greatly benefited from, Fusakichi Omori and other 

Japanese colleagues in setting up the Whitney Laboratory 

of Seismology and the initiation of seismic monitoring at 

Kīlauea. Later, under the Japan-United States Cooperative 

Science Programme, a team of Japanese scientists lived and 

worked at HVO for about 8 months during 1963–64, during 

which time two eruptions occurred on Kīlaueaʼs East Rift 

Zone (Aloi Crater, August 1963; Nāpau Crater, October 

1963). By coincidence, also in August 1963, a Japanese 

training vessel (Kagoshima Maru) was visiting Hawai‘i, and 

an informal agreement was made with the ship’s captain to 

conduct a bathymetric survey of Papa‘u Seamount offshore of 

Kīlauea’s south flank (James G. Moore, oral commun., 2012). 

During the survey, two HVO staff (James Moore and Harold 

Krivoy) worked aboard the ship to assist with a depth recorder, 

and Dallas Peck and other HVO staffers installed and oversaw 

three transit stations on shore to track the ship’s position every 

10 minutes. This target-of-opportunity study resulted in the 

first offshore studies of Kīlauea’s south flank; the bathymetric 

map obtained (Moore and Peck, 1965), together with seafloor 

mapping by the U.S. Navy around the Hawaiian Islands in the 

late 1950s (the Pioneer/Rehoboth surveys), contributed to the 

discovery of Hawaiian submarine landslides, first described 

by Moore (1964). Following joint studies of the Aloi and 

Nāpau eruptions and the bathymetric survey collaboration 

with Japanese scientists—formal and informal—continued 

for years afterward, including a visit in 1965 by several HVO 

staff to Japanese volcano observatories at Asama, Aso, and 

Sakurajima. During this visit, Taal Volcano (Philippines) 

began to erupt, and James Moore and Kazuaki Nakamura 

traveled to the Philippines to assist local colleagues making 

key observations, including the documentation of pyroclastic 

base-surge phenomena (Moore and others, 1966).

This bilateral program with Japan in the mid-1960s 

exemplifies HVO’s collaborative research (for example, 

Minakami, 1965; Peck and Minakami, 1968; Aramaki 

and Moore, 1969). Since then, HVO has participated in 

cooperative studies with researchers from many scientific 

institutions (including, but not limited to, the University of 

Hawai‘i at Mānoa, University of Oregon, Stanford University; 

Monterey Bay Aquarium Research Institute; Smithsonian 

Institution; and University of Massachusetts). These joint 

studies have focused on many aspects of Hawaiian volcanism, 

including petrologic-geochemical dating studies of eruptive 

products; lava-tube processes and lava-flow emplacement 

(for example, Cashman and others, 1999; Kauahikaua and 

others, 1998, 2003); Kīlauea’s pre-20th century explosive 

eruptive activity (for instance, Fiske and others, 2009; 

Swanson and others, 2011a,b; 2012a,b); Mauna Loa Volcano 

(Rhodes and Lockwood, 1995; and references therein); slow 

slip dynamics of Kīlauea’s south flank (for example, Cervelli 

and others, 2002a; Wolfe and others, 2007; Brooks and 

others, 2008; Montgomery and others, 2009, 2010; Poland 

and others, 2010); and seismic tomography of the Hawaiian 

hot spot (for instance, Wolfe and others, 2009, 2011). Using 

data from the broadband seismic network around Kīlauea’s 

summit, HVO and USGS investigators, working with national 

and international collaborators, have made significant 

contributions to volcano seismology. A sampling of such 
36 Characteristics of Hawaiian Volcanoes

studies include modeling sources of shallow tremor at Kīlauea 

(Goldstein and Chouet, 1994); seismic refinement of the 

three-dimensional velocity structure of the Kīlauea Caldera 

(Dawson and others, 1999); and analyses of long-period (LP) 

and very-long-period signals to characterize seismic source 

regions and hydrothermal systems beneath Kīlauea Caldera 

(Ohminato and others, 1998; Almendros and others, 2001, 

2002; Saccorotti and others, 2001; Kumagai and others, 2005). 

This collaboration continues with studies of the seismicity of 

the 2008 summit lava lake (Chouet and others, 2010; Dawson 

and others, 2010).

Another very productive collaborative research took 

place during 1998–99 between scientists of the USGS, 

University of Hawai‘i, Monterey Bay Aquarium Research 

Institute, and several Japanese universities, under the auspices 

of the Japanese Marine Science and Technology Center 

(JAMSTEC). A number of cruises using the latest techniques 

in submarine geological studies, including the RV Kaiko (a 

remotely operated vehicle or ROV) and the then-deepest 

diving manned submersible (the Shinkai 6500) examined 

the deep submarine flanks of several Hawaiian volcanoes. 

The results were published in a monograph of the American 

Water

depth,

in meters

157° 156° 155° 154°

18°

19°

20°

0

500

1,000

1,500

2,000

2,500

3,000

3,500

4,000

4,500

5,000

5,500

6,000

0

0

50 KILOMETERS

25 50 MILES

Figure 24. Map showing the bathymetry of the sea bottom around the Island of Hawai‘i. Contour interval is 200 m. This bathymetry became 

much better mapped with high-resolution data obtained during the 1998–99 cooperative research between Japanese and U.S. scientists (see text). 

This image is modified from one of many illustrations in the CD–ROM accompanying the monograph by Takahashi and others (2002).
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 37

Geophysical Union (Takahashi and others, 2002), which 

includes 27 papers germane to the deep underwater geology 

around the Hawaiian Ridge. In addition, this volume is 

accompanied by a CD-ROM that contains high-resolution 

images of bathymetry (fig. 24), spreadsheets of the analytical 

data for samples collected, and other useful compilations.

HVO has also participated in other cooperative efforts 

that did not necessarily result in scientific publications. For 

example, in 1965, HVO conducted a field seminar for a group 

of National Aeronautic and Space Administration (NASA) 

astronauts slated for upcoming Moon landings of NASA’s 

Apollo Program. Hawai‘i provided an excellent place for 

such astronaut training because of its well-displayed volcanic 

features that could be compared and related to similar features 

on the lunar surface. Because of its frequent and, sometimes, 

prolonged eruptions, Kīlauea constitutes an accessible 

global resource for learning or teaching about volcanism. 

Consequently, HVO has interacted informally with many 

students, educators, and visiting scientists, seeking to collect 

samples, exchange ideas, or to test hypotheses or teaching 

curricula. Over the decades, many volcanologists from other 

countries have interned among HVO staff to learn monitoring 

techniques. In this regard, to take best advantage of HVO’s 

staff and resources, former Scientist-in-Charge Robert W. 

(Bob) Decker founded the Center for the Study of Active 

Volcanism (CSAV; http://www.uhh.hawaii.edu/~csav/) in 

1989 at the University of Hawai‘i at Hilo. CSAV is a training 

and outreach program and operates an international training 

course, intended primarily for scientists and science-support 

personnel in countries that have active or potentially active 

volcanoes. Present and past HVO staff members serve as 

instructors, and HVO provides some logistical support.

Communication of Scientific Information 

and Public Outreach

One of HVO’s main goals was—and remains—to 

“keep and publish careful records” (Jaggar, 1913, p. 4) of 

the observations of eruptive and associated activities at 

Hawaiian volcanoes. From 1912 through 1955, in addition 

to publications in peer-reviewed scientific journals, this was 

accomplished regularly by the observatory’s several serial 

publications (Special Reports, Weekly Reports, Monthly 

Bulletins, and Volcano Letters). These early publications 

presented chronological narratives of the visible eruptive 

activity, as well as associated seismic, geodetic, and other data. 

Some also contained discussions of volcanic and earthquake 

activity elsewhere in the world. After 1955, however, 

chronological narratives and results of volcano-monitoring and 

topical studies—geologic, geophysical, and geochemical—

of Hawaiian volcanism have been presented in USGS 

publications, journal articles, and books, now numbering 

many thousands. Suffice it to say, the volcanic behavior of 

Kīlauea and Mauna Loa since the early 20th century has been 

extraordinarily well documented. 

For all the science accomplishments at HVO in the 

past century, the support and survival of the observatory 

also required communicating and sharing information and 

interpretations of Hawai‘i’s eruptions and restless activity 

with fellow scientists and the public. Jaggar was well aware 

of the importance of keeping the general public informed; 

thus, he regularly spoke publicly and wrote many articles 

for newspapers about the activities at Kīlauea and Mauna 

Loa. Indeed, the all-inclusive communication of scientific 

information was a central theme of HVO long before the term 

“public outreach” was coined.

Print Publications

Jaggar (1912, p. 3) advocated that “Results obtained 

in connection with all subjects of investigation should be 

promptly published in the form of bulletins and memoirs.” 

Even before it occupied its first permanent building in 1912, 

the new observatory made public its scientific findings on a 

regular basis. A weekly update on Kīlauea Volcano activity 

was started by Frank Perret, who wrote six weekly articles 

issued between August 15 and September 17, 1911, for 

The Pacific Commercial Advertiser, published by Lorrin 

A. Thurston. When Jaggar arrived the following January, 

he immediately resumed the weekly updates, with the first 

issued on January 25, 1912. Weekly Reports and Bulletins 

became Monthly Bulletins that were published by the 

Hawai‘i Volcano Research Association, with the last issue 

covering the month of July 1929. The Volcano Letter, a 

monthly publication that included both scientific articles and 

updates of volcanic and earthquake activity, started being 

issued in late 1925 and continued until mid-1955. All of 

these early serial publications of HVO are readily accessible 

in several comprehensive volumes of collected works (Fiske 

and others, 1987; Bevens and others, 1988; Wright and 

Takahashi, 1989; Wright and others, 1992). 

Once it became a permanent part of the USGS in 1947, 

HVO began publishing more technical summaries with annual 

reports from 1948 to 1955, quarterly reports from 1956 to 

1958, and annual seismic summaries that included volcanic 

details starting in 1956 and continuing today as an annual 

seismic catalog. After the cessation in 1955 of The Volcano 

Letter, however, there was no regular popular medium through 

which even local residents could keep up with Hawaiian 

volcanic and earthquake activity, except for the frequent 

press releases directly related to increased or new activity and 

possible imminent hazards. As a result, the USGS, including 

HVO, later began to publish Fact Sheets—involving one or 

two pages of well-illustrated, nontechnical discussion—on 

popular topics (http://hvo.wr.usgs.gov/products/#factsheets).

While Bob Decker was a professor of geology at 

Dartmouth University, he wrote and submitted a weekly 

“Volcano Watching” article to the Hawaii Tribune-Herald
38 Characteristics of Hawaiian Volcanoes

during the year 1975. These articles were collated and 

formed the basis for a general-interest book about Hawaiian 

volcanism (Decker and Decker, 1980). Decker later served 

as HVO Scientist-in-Charge (SIC) during 1979–84. This 

weekly newspaper series was reinstituted in its present 

form (now called “Volcano Watch”) in November 1991 

by David A. Clague, HVO SIC during 1990–96, and 

continues to be written by HVO scientists and published 

in local newspapers and posted on the Internet (Hawaiian 

Volcano Observatory Staff, 2000). Each Volcano Watch 

includes an article about a geologic topic of relevance to 

volcano enthusiasts and Hawai‘i residents, as well as an 

update on current volcanic activity and descriptions of felt 

earthquakes. Through early 2013, more than 1,000 Volcano 

Watch articles have been written and shared with local 

residents via newspapers and with a broader audience via 

HVO’s Web site.

While these popular publications were being produced, 

HVO staff continued to keep their “careful records” in the 

form of Record Books that included daily observations 

and photos from 1912 to 1955, as well as internal weekly, 

monthly, bimonthly, and quarterly reports starting in 1967 

that were meant to document important details. While not 

official publications, these serve as resources for further 

research and as starting materials for formal publications.

HVO Web Site

HVO’s first Web site was initiated by several 

HVO staff in 1997 and consisted of infrequent updates 

on volcanic activity with some volcano background 

information. Within a year, a new Web site was launched 

to include near-monthly updates of Kīlauea’s eruptions, 

including images, graphics, and maps, and background 

information on Hawai‘i’s active volcanoes and earthquakes. 

The site continues to evolve with daily updates (more 

frequently, as needed, with rapidly changing activity), realtime monitoring data and webcams of erupting vents and 

flow fields, time-lapse digital photography of remarkable 

activity, background volcano information, types of hazards, 

photo field tours, and more. The HVO Web site continues to 

get several hundred thousand hits per day, but that number 

dramatically increases up to several million per day when 

activity picks up or a new eruption commences.

Public Education and Outreach

From the beginning, HVO has endeavored to routinely 

and effectively communicate the results of its work to 

the scientific community, Federal and local government 

agencies, news media, and the general public. As 

discussed previously, HVO uses many existing avenues 

of communication—print and electronic—to disseminate 

scientific information, and the observatory anticipates 

greater use of social media outlets in the future.

Volcano Awareness Month

In addition to ongoing efforts throughout the year, 

Hawai‘iʼs volcanoes received a boost in attention during the 

month of January, beginning in 2010, when Hawai‘i County 

Mayor Billy Kenoi declared it “Volcano Awareness Month.” 

The motivation for this specific designation is to enhance public 

awareness of Hawaiian volcanoes and associated hazards in 

particular and volcanic phenomena in general. Each January, 

many lectures, movies, excursions, and other volcano-related 

activities, in which HVO and affiliated scientists are heavily 

involved, are scheduled for the education and enjoyment of the 

public. (For a representative listing of activities, see Hawaiian 

Volcano Observatory Staff, 2012b). In connection with 

Volcano Awareness Month, an eight-page newspaper insert was 

produced in 2010 that included short articles on current activity 

and other topics of relevance to Hawai‘i residents who live on 

active volcanoes. On January 21, 2012, as part of the year-long 

celebration of its centennial, HVO held an open house that drew 

more than 1,400 visitors and at which was premiered a new 

general-interest booklet that recounts the story of HVO and of 

its century of scientific accomplishments and public service 

(Babb and others, 2011).

Is Public Outreach Working?

Given the considerable outreach efforts on the part 

of HVO, how aware are Hawai‘i residents about volcanic 

eruptions and hazards? A recent survey reported that residents 

of the Kona area (west Hawai‘i) exposed to lava flows from 

Mauna Loa and Hualālai “have received little or no specific 

information about how to react to future volcanic eruptions or 

warnings, and short-term preparedness levels are low” (Gregg 

and others, 2004, p. 531). In contrast, residents on the east 

side of the Island of Hawai‘i, who are frequently exposed to 

volcanic activity, are better informed and prepared to respond 

to a volcanic emergency (western Hawai‘i residents have not 

witnessed any volcanic activity since the 1950 eruption of 

Mauna Loa). Although greater efforts can be made to increase 

public awareness of volcano hazards in western Hawai‘i, 

the studies of Gregg and others (2004, 2008) may simply 

reflect basic human behavior, as demonstrated in nearly all 

volcanic regions of the world: the people directly affected 

by frequent eruptions naturally become more knowledgeable 

and, hence, better prepared to cope with future eruptive 

activity. Another contributing factor is that detailed guidance 

or plans of action about how to react, or what to do, during a 

volcanic crisis traditionally have not been given until a crisis 

begins to unfold. In recent decades, increased attention has 

been paid to studies of how people in various countries and 

cultures perceive volcanic risk and how they respond during 

volcanic crises (for example, Gaillard and Dibben, 2008). 

The findings of such studies vary in emphasis and specificity, 

but they all share a common thread: scientists are not, and 

cannot be, solely responsible for enhancing public awareness 

of volcanic hazards. Rather, meeting this challenge requires a 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 39

close partnership between scientists, government authorities, 

and other stakeholders to develop and implement effective 

risk-reduction strategies, tailored to the local jurisdiction. 

Specifically regarding Hawai‘i, HVO must maintain its present 

level of islandwide public outreach but, at the same time, 

strive to work more closely with State and County government 

officials to find ways to change the current perception of 

volcanic risk held by western Hawai‘i residents. Those 

residents need to be better prepared, should Hualālai reawaken 

or an eruptive outbreak occur on Mauna Loa’s western flank.

Some Notable Developments Since 

HVO’s 75th Anniversary

In commemoration of HVO’s 75th anniversary, the 

USGS published the massive, two-volume Professional 

Paper 1350 (Decker and others, 1987), which contained 62 

papers that presented results from diverse studies conducted 

by HVO and affiliated scientists through the mid-1980s. 

This comprehensive volume may be likened to a time 

capsule of the state-of-the-art knowledge about Hawaiian 

volcanism, acquired using the monitoring techniques and 

research disciplines available to HVO and, indeed, to the 

global volcanologic community at the time. In writing the 

introductory chapter to this current volume, we have chosen 

to use for our timeframe the entire first 100 years of HVO’s 

tracking of eruptions and earthquakes (Babb and others, 2011; 

Kauahikaua and Poland, 2012)—taking a largely retrospective 

look at the work of the observatory and its accomplishments 

that are not, or barely, discussed in the excellent historical 

summary by Apple (1987). Nonetheless, given the post1987 explosive growth in instrumentation, technology, and 

computer-based data processing and modelling, remote 

sensing, satellite-based geochemical and geodetic monitoring, 

and basic knowledge of physical volcanology, we believe 

it is instructive to highlight some major developments 

during the past quarter century at Kīlauea and Mauna Loa 

volcanoes, and in HVO’s capabilities to study them. Some 

aspects of these developments (summarized in table 3) have 

been discussed in the preceding sections of this chapter, but 

some are noted here for the first time. In any case, it is clear 

from table 3 that many powerful tools and techniques used 

in studying volcanic phenomena (for instance, GIS, GPS, 

InSAR, remote gas measurements, broadband seismometers, 

digital borehole tiltmeters and strainmeters)—which we now 

take for granted—have only become available and applied 

in recent decades. As examples, the first GPS measurements 

were not made in Hawai‘i until 1987 (Dvorak and others, 

1994), and data from digital broadband seismometers were 

unavailable before 1987 and hence not considered in the 

comprehensive summary of Kīlauea’s seismicity by Klein 

and others (1987). Such seismic data nowadays are necessary 

prerequisites for quantitative seismological studies of 

volcanoes (for example, Chouet, 1996, 2003).

Looking to the Future

With HVO’s first 100 years now past, what might we expect 

in the future? Unless plate-tectonic dynamics inexplicably change 

or cease to operate—a highly unlikely eventuality—Hawaiian 

volcanoes, especially Kīlauea, will continue to erupt frequently. 

Given the new interpretation of Kīlauea’s history of centurieslong alternation between effusive and explosive volcanic 

behavior (see “Explosive Eruptions” section, above), however, it 

is an open question whether future activity will be a continuation 

of the current mode of dominantly effusive eruptions or whether 

the volcano will enter a long period of explosive eruptions, as 

during 1500–1800 C.E. In any case, we can expect HVO and 

affiliated scientists to continue to use and improve existing 

approaches and tools in studying future eruptions, whatever their 

nature. We can also expect to see greater use of new, emerging 

monitoring techniques and improved means to share and 

communicate scientific information, some of which we touch 

upon in the discussion below.

Lava Flux Monitoring

The rate at which lava is erupted is a critical measure of 

an eruption’s status. The most widely used technique at HVO 

to estimate lava production is to repeatedly map the lava-flow 

area and estimate thicknesses across the area, from which an 

erupted volume can be calculated. Because reliable estimation 

of lava flux rates is such an important monitoring goal, a few 

new techniques have been developed, and the most promising 

ones are briefly summarized below.

Electromagnetic (EM) techniques have always offered 

great promise for volcano studies. Because hot, molten magma 

has a high electrical conductivity compared to cold, solidified 

lava, the differences between hot and cold domains should 

be easily detectable and mappable. In practice, however, the 

results of EM surveys in structural studies have been mixed, 

because groundwater-saturated lavas have moderately high 

conductivities compared to dry lavas (see, for example, 

Jackson and Keller, 1972). Nonetheless, shallow-depth EM 

techniques have proved to be very useful in mapping the 

molten interiors of active lava flows and lava tubes (for 

example, Kauahikaua and others, 1998). 

Under specific circumstances, the flux of lava flowing 

through a tube can be measured with the very low frequency 

(VLF) technique, which uses the electromagnetic fields 

created by remote powerful radio transmitters. If an open 

skylight is available so that a flow velocity for the lava can 

be measured, a VLF profile obtained at that location can be 

interpreted as a cross-sectional area of the molten lava in the 

tube which, when multiplied by the lava velocity, yields the 

lava flux through the tube. Monitoring the decreasing lava flux 

through Pu‘u ‘Ō‘ō’s master tube in the early 1990s led to a 

forecast that its feeding vent(s) would shut down soon; within 

a month of the forecast, the vent actually did become inactive 

(Kauahikaua and others, 1996).
40 Characteristics of Hawaiian Volcanoes

Table 3. Some major developments or events at Kīlauea and Mauna Loa volcanoes and at the Hawaiian Volcano Observatory since its 75th 

anniversary (Diamond Jubilee) in 1987 (compiled from diverse sources).

Development(s) or event(s) Relevance / significance / importance Selected references (and references therein)

Continuation of the Pu‘u ‘ŌʻōKupaianaha eruption into the 21st 

century.

The longest duration, nearly continuous, riftzone eruption at Kīlauea in more than 600 years, 

affording unprecedented opportunities to study a 

wide variety of processes and products of effusive 

basaltic volcanism.

Wolfe (1988)

Heliker and others (2003b)

Orr and others (2012)

Growth and collapse of the complex 

cone at Pu‘u ‘Ōʻō 1983–2002. 

The evolution of this prominent volcanic landform—

one of the most recent in the United States—is 

exceptionally well documented, both visually and 

instrumentally.

Heliker and others (2003a)

Detailed documentation of eruption 

dynamics in the development of 

extensive lava-flow fields, 1983–

present, extending and refining the 

findings from previous studies during 

the 1969–74 Mauna Ulu eruption.

The prolonged and still-continuing Pu‘u ‘ŌʻōKupaianaha eruption has provided unprecedented 

opportunities to make detailed studies of the key 

processes involved in the development of extensive 

lava-flow fields. Such processes include formation 

of complex lava-tube systems, pāhoehoe-‘a‘ā 

transitions, inflation of pāhoehoe flows and sheets, 

construction of perched lava ponds and channels, 

and building and collapse of lava deltas.

Heliker and others (2003b)

Kauahikaua and others (2003)

Greatly increased magma supply to 

Kīlauea during 2003–07.

Long inferred to vary little over decadal time 

scales, the magma supply rate to Kīlauea at least 

doubled during 2003–07 relative to previous rate 

estimates. This finding was made possible by a 

combination of detailed geodetic, geochemical, 

and gas-emission data.

Dvorak and Dzurisin (1993)

Poland and others (2012)

Poland and others (this volume, chap. 5)

Since mid-March 2008, renewed 

eruptive activity and operation of an 

active lava lake at Kīlauea summit.

The opening of the new vent within Halema‘uma‘u 

represents the first summit eruptive activity since 

1982. The ongoing lava-lake activity marks the 

longest duration summit eruption since 1924 and 

the first in which two vents—at the summit and rift 

zone—were simultaneously active for many years. 

The summit vent opened following the increased 

magma supply 2003–07.

Hawaiian Volcano Observatory Staff (2008)

Patrick and others (2013)

Long-term net subsidence of Kīlauea 

summit with the onset of the Pu‘u 

‘Ōʻō-Kupaianaha eruption in 1983, 

interrupted by a conspicuous inflation 

period during 2003–07.

This long-term summit subsidence was the longest 

in duration since continuous measurement of tilt 

began in 1912. The location of maximum subsidence 

(~2 km south of Halema‘uma‘u) dropped by more 

than 1.5 m between 1983 and 2003.

Tilling and others (2010)

Poland and others (2012)

(see also fig. 4, this chapter)

Changing magma supply rate to 

Mauna Loa? Reversal of overall 

Mauna Loa summit deflation trend 

since 1994 with renewed inflation 

beginning in May 2002. But then 

inflation decreased.

The rate of renewed inflation reached a maximum 

during 2004 before declining, perhaps suggesting 

that Mauna Loa’s magmatic system was also 

affected by the mantle-driven surge during 2003–07 

in magma supply to Kīlauea. Can this coincidence 

be taken as indirect evidence that the two systems 

somehow interact?

Miklius and others (2002)

Hawaiian Volcano Observatory Staff (2003)

Miklius and Cervelli (2003)

Gonnermann and others (2012)

Poland and others (2012)

The series of Japan-U.S. cooperative 

deep-water research cruises during 

1998–99 produced much new 

bathymetric and geologic data 

about the deep submerged flanks 

of Hawaiian volcanoes, especially 

Kīlauea.

Scientific information from these cooperative studies 

has added a submarine perspective in understanding 

the emplacement of pāhoehoe flows, morphology of 

rift zones, ancestral growth of Kīlauea, volcano flank 

stability, and other aspects of Hawaiian volcanism.

Takahashi and others (2002)

Clague and Sherrod (this volume, chap. 3)
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 41

Development(s) or event(s) Relevance / significance / importance Selected references (and references therein)

Recent discovery that Kīlauea’s 

volcanic behavior has been 

frequently, and sometimes 

energetically, explosive in the recent 

geologic past.

The new interpretation that Kīlauea can alternate 

in volcanic behavior between centuries-long 

dominantly effusive eruptions (19th century–present) 

and centuries-long explosive eruptions (1500–1800 

C.E.) shatters the long-held thinking that it is a 

“benign” volcano. This discovery has potentially 

serious hazard consequences for the summit area.

Fiske and others (2009)

Swanson and others (2011a,b)

Swanson and others (2012a,b)

Regular measurement of emission 

rate of SO2 at Kīlauea summit since 

1979 and the continuous monitoring 

of SO2 emission at both summit and 

East Rift Zone since 1997.

In addition to contributing to an improved 

understanding of Hawaiian volcanism, HVO’s 

continuous monitoring of SO2 emissions at 

Kīlauea—comprising the longest duration dataset of 

its kind—is critical for the assessment of the health 

hazards posed by vog (volcanic smog) to Hawai‘i’s 

residents and visitors. 

Sutton and others (1997)

Elias and Sutton (2012)

Sutton and Elias (this volume, chap. 7)

Kauahikaua and Tilling (this volume, chap. 10)

Miniaturization of instrumentation 

and improvements in gas-emission 

measurements using remote-sensing 

techniques.

The availability of lightweight, easily portable, and 

low-cost instruments and continuously recording 

systems have revolutionized volcanic gas monitoring 

of volcanoes in Hawai‘i and elsewhere.

Elias and others (2006)

Horton and others (2006)

Beginning in the early 1990s, 

installation of a subset of digital 

broadband seismometers as part of 

HVO’s telemetered seismic network.

Data from the broadband network have enabled 

tomographic studies of Kīlauea’s shallow magmatic 

system, precise determinations of the locations and 

mechanisms of long-period (LP) and very-longperiod (VLP) events, and correlations with observed 

rock falls and related degassing bursts in the active 

vent in Halema‘uma‘u Crater.

Dawson and others (1998, 1999, 2010)

Chouet and others (2010)

Patrick and others (2011b)

Orr and others (2013)

Okubo and others (this volume, chap. 2)

Precise determination of lava flow 

paths and refined frequency of 

inundation by lava of specific areas.

Geologic and dating studies have enabled more 

detailed characterization of the frequency of 

inundation by lava flows in specific areas downslope 

from potential eruptive vents. Such information 

has important implications for lava-flow hazards 

assessments. Specific pathways are predicted using 

digital elevation models (DEM) once a vent location 

or lava flow front is known. Although improved 

geologic and age data are useful for assessing 

hazards probabilities, they do not determine 

“precise” flow paths.

Kauahikaua (2007)

Trusdell and others (2002)

Kauahikaua and Tilling (this volume, chap. 10)

Use of digital time-lapse visual and 

thermal cameras in systematically 

documenting eruptive processes at 

erupting vents, along lava channels 

and flows, and where lava enters the 

ocean.

The level of detail in near-real time with Global 

Positioning System (GPS) timing affords detailed 

documentation of eruptive processes and their 

resultant products that can be compared precisely 

with geophysical data.

Orr and Hoblitt (2008)

Patrick and others (2011a,b)

Orr (2014)

Transition from “classical” to 

space geodesy in the monitoring of 

Hawaiian volcanoes.

Beginning in the late 1980s, HVO has been the 

proving ground for the testing and use of geodetic 

techniques for ground-deformation monitoring in 

real or near-real time. This has made possible the 

documentation of short-term processes not detected 

by previous infrequent campaign-style techniques.

Decker and others (2008)

Dzurisin (2007)

Poland and others (this volume, chap. 5)

Table 3.—Continued.
42 Characteristics of Hawaiian Volcanoes

Lava flux also can be estimated through knowledge of 

the sulfur dioxide emission rate, assuming that a fixed relation 

exists between the mass of SO2 emitted and the mass of lava. 

Applying this approach to Pu‘u ‘Ō‘ō activity during the period 

1997–2002, Sutton and others (2003) compared total volume 

estimates derived from SO2 emission rates with those from VLF 

measurements and found that they agreed within 10 percent.

Another promising tool may be the detection of lava and 

the estimation of eruption rate through the thermal radiance 

of the flowing lava itself. Harris and others (1998) first used 

Landsat satellite data to estimate instantaneous lava effusion 

rates at Kīlauea to test this idea. Wright and others (2001) 

provided a slightly different view, measuring average effusion 

rates by quantifying the area of lava flows. Thermal detection 

and tracking of changes during eruption was highly useful in 

establishing the correct timeline of a remote fissure eruption 

in 1997 (Harris and others, 1997). The most recent innovation 

in applying thermal radiance techniques is the use of handheld 

thermal cameras (for example, at Piton de la Fournaise; 

Coppola and others, 2010); however, this approach has not yet 

been tested in Hawai‘i. In 2010, HVO began using stationary 

infrared cameras to monitor Kīlauea’s eruptive vents, and 

work is in progress toward creating software alarms that 

trigger on detection of high radiant temperatures (M. Patrick, 

oral commun., 2011; Patrick and others, 2014).

Mapping Active Lava Flows

Much of the work that HVO’s geologists do during 

any eruptive crisis is mapping lava flows for the purpose of 

documenting their progress, especially the rate and direction 

of their advance, and estimating their discharge. Through 

the 1990s, flow mapping was done by traditional geologic 

mapping methods (aerial photographs) and surveying 

techniques (for instance, Wolfe and others, 1988). After GPS 

satellite transmission scrambling was turned off in May 2000, 

mapping accuracy of handheld GPS receivers improved to less 

than 10 m horizontally, but it was still necessary to carry the 

GPS receiver along the flow contacts and transfer the data into 

a GIS software system from which maps could be made. In the 

recent mapping and observations of active lava flows, the use 

of digital time-lapse photography, video, and thermal images 

has been highly instructive (for example, Patrick and others, 

2011a,b; Patrick and Orr, 2012a; Orr and others, 2013).

 Techniques using a pair of synthetic aperture radar 

(SAR) images are best known for their excellent resolution 

of ground deformation by interferometry (InSAR; see, for 

example, Dzurisin and Lu, 2007; Lu and Dzurisin, 2014, and 

references therein). Changes between SAR image acquisitions, 

however, can also delineate areas of incoherence, which 

denote regions of changing ground properties—for example, 

with resurfacing of the ground by new lava. This technique 

can be used to document recent lava flow activity in areas 

previously covered by older lava. Zebker and others (1996) 

demonstrated that such SAR incoherence maps do an excellent 

job of tracking active lava flows, and possibly effusion rates 

(given field-measured flow thicknesses), at Kīlauea. More 

recent work (Dietterich and others, 2012) further improves the 

technique and better demonstrates its ability to track active 

flows. This new spaced-based technique will never totally 

replace traditional ground mapping because of the inherent 

delay (latency) in round-trip transmission of satellite data, but 

SAR mapping may be very useful at active volcanoes where 

ground access is difficult and (or) dangerous.

Ambient Noise Seismic Tomography and 

Monitoring

Traditional seismology is based on the recording, 

processing, and interpretation of seismic signals—

earthquakes, tremors, and teleseisms are among the most 

common types of seismic energy that travels through the 

Earth. But these signals are mixed with seismic “noise.” One 

of the main sources of ambient seismic noise is ocean waves, 

and techniques have been developed recently to use this noise 

as a source for monitoring subsurface activity. Brenguier and 

others (2007, 2008, 2011) and Duputel and others (2009) have 

used ambient seismic noise to successfully monitor Piton 

de la Fournaise, a frequently active shield volcano on the 

French island of La Réunion in the Indian Ocean. As of this 

writing (mid-2014), seismologists at HVO and the University 

of Hawai‘i are developing the tools to use this emerging 

technique as another way to continuously monitor the seismic 

properties of the subsurface. 

The year 2012 marked the start of a 5-year program of 

scientist exchanges with the Observatoire Volcanologique du 

Piton de la Fournaise, and one of the goals of this program 

will be the transfer of modern monitoring techniques 

in both directions. Such an exchange should further the 

development of ambient noise monitoring techniques at both 

observatories.

Unmanned Aircraft Systems (UAS)

Until recently, unmanned aircraft systems (UAS)—

commonly called drones—have been used almost exclusively 

for military and espionage purposes. In recent years, 

however, UAS have been increasingly deployed for some 

specific civilian applications (for example, remote-sensing 

studies, fighting of wildland fires, tracking hurricanes, 

surveillance of pipelines) where there is a need for rapid, 

low-cost reconnaissance of large areas at no health or safety 

risk to personnel (see, for instance, Merlin, 2009). The 

USGS has used UAS in various biological or environmental 

projects (see http://uas.usgs.gov/) and even collaborated with 

Advanced Ceramics Research to use UAS for monitoring 

lava dome growth at Mount St. Helens, Wahsington, 

in 2004 (Patterson and others, 2005; Smith and others, 

2009). To date, UAS have not been employed at Hawaiian 

volcanoes; although HVO scientists are eager to find and 
The Hawaiian Volcano Observatory—A Natural Laboratory for Studying Basaltic Volcanism 43

test possible applications for UAS platforms in volcano 

monitoring and research, at the moment we are only in the 

early planning stages until uncertainties about available 

sensors are resolved. HVO has been involved, however, in a 

few innovative collaborative projects using meteorological 

balloons in connection with studies of volcanic gas 

distribution on a regional scale (for example, Donovan, 

2008). In addition, airborne SAR acquisitions of Kīlauea 

have been made by NASA’s Jet Propulsion Laboratory 

(Lundgren and others, 2013) using an instrument that is 

intended to one day fly on board UAS.

Greater Use of Social Media in Communicating 

Hazards Information

Early in its history, HVO’s hazard assessments and 

related information were first made public by direct telephone 

communication and further disseminated by newspaper 

publication. From the 1980s until 2005, direct faxes to the 

emergency-management officials were also used to spread 

the word. The next major change in communication strategy 

was HVO’s adoption of the communication potential of the 

Internet in 1997. HVO’s original Web site hosted limited 

material, mostly hazards-related information releases, monthly 

updates on volcanic and earthquake activity, listings of 

new publications, and the like. The Web site content is now 

much more comprehensive, posting real- and near-real-time 

monitoring data, up-to-date lava flow maps, locations of current 

earthquakes, photographs, and many links to other information 

about Hawaiian volcanoes. Emergency managers (for example, 

Hawai‘i Volcanoes National Park, County of Hawai‘i, and 

State of Hawaii) continue to be updated directly via phone calls 

and e-mails. In fact, any person interested can subscribe to the 

USGS Volcano Notification Service to automatically receive 

e-mail volcano alerts from HVO or any other USGS volcano 

observatory by signing up at http://volcanoes.usgs.gov/vns/.

With the advent of social media outlets, such as Twitter 

and Facebook, anyone can sign up and post text, photographs, 

and video that can be shared with a specified set of fellow 

users ranging from everyone to only your closest friends. 

An added advantage is the ability to be alerted automatically 

to changes in the Facebook or Twitter offerings of any of 

your friends. No longer do interested people need to keep 

checking for changes in someone’s status or postings—they 

can be alerted when these changes occur. The implications 

for hazards communication are obvious. An entity such as 

HVO could have relatively static content available for those 

who seek it, as well as content that is regularly updated with 

automatic alerts going to anyone who signs up for them. Once 

a user has found the site and worked through the information 

of interest at that time, he or she can be automatically alerted 

when that information changes.

For effective communication of hazards information, 

every available communication medium should be used—the 

message must get to the places where the public is listening. 

No longer can we expect people who may be affected by 

natural disasters to search for the communication method(s) 

we are using. We must deliver our message in every 

way conceivable to most effectively disseminate hazards 

information in a timely manner. As an agency, the USGS has 

been using social media to communicate with stakeholders 

since 2007, when the first podcast was launched; however, 

many of the primary social media outlets that are used today 

were only fully employed by the USGS in 2010. These primary 

resources—like Facebook, Twitter, YouTube, and Flickr—are 

information-sharing Web sites that people visit to learn about 

timely news, view imagery, watch videos, and interact with 

USGS social media ambassadors through comment strings. 

USGS also uses subscription-based data feeds to push content 

to individuals who have requested to receive specific updates. 

To communicate programmatic news and public updates, 

the USGS Volcano Hazards Program (VHP) uses the general 

USGS social media outlets. For very specific information 

relating to volcano hazards, however, the VHP employs datafeed services and has communicated up-to-the-minute eruption 

information via event-specific Twitter feeds. For example, 

the Volcano Notification Service is an RSS (“really simple 

syndication”) feed, and a Twitter account was used during the 

2009 eruption of Redoubt, in Alaska, when eruption updates 

were automatically pushed to followers of the Redoubt-2009 

Twitter feed. In future volcanic crises, the VHP will again use 

event-specific accounts with social media Web sites and data 

feeds to communicate critical information to stakeholders. 

For everyday communication of news releases, observatory 

operations, and noncritical eruption updates, HVO and the 

VHP will continue to use the more broad USGS umbrella 

social media outlets (http://www.usgs.gov/socialmedia/).

Continuing Integration into National-Scale 

Volcano-Monitoring Efforts

The Hawaiian Volcano Observatory predates by 70 years 

the establishment of the second USGS volcano observatory—

the Cascades Volcano Observatory (CVO), which was 

formally established in 1982 after the 1980 eruption of Mount 

St. Helens. The USGS’s family of observatories then grew 

to five with the addition of the Alaska Volcano Observatory 

(AVO) in 1988, Long Valley Observatory (LVO) in 1999, 

and Yellowstone Volcano Observatory (YVO) in 2001. In 

a reorientation of the VHP in 2012, LVO ceased to exist 

formally but its functions were incorporated into the newest 

USGS observatory—the California Volcano Observatory 

(CalVO). For more information about all USGS volcano 

observatories, the interested reader is directed to the VHP Web 

site (http://volcanoes.usgs.gov/).

Expanding technical capabilities in the 21st century 

now allow rapid communication and sharing of data among 

the USGS volcano observatories, thereby promoting a much 

higher degree of interoperability between them. For example, 

the VALVE graphic-display software (Cervelli and others, 
44 Characteristics of Hawaiian Volcanoes

2002b, 2011), which was developed at HVO during the early 

2000s, is now being upgraded and installed at all U.S. volcano 

observatories. A more recent development, an instrument 

site database that stores everything from land-access permits 

and instrument serial numbers to a log of site visits and 

instructions to find monitoring sites, was also developed at 

HVO and is being deployed at all USGS volcano observatories 

and other sites involved in volcano hazards studies. The 

increase in the number of observatories, together with the VHP 

emphasis on national focus and framework, thus mutually 

benefits all volcano observatories.

Many aspects of volcano monitoring can now be 

conducted remotely, by both ground- and space-based 

systems. Because of its early development, the buildings of 

HVO sit on the rim of Kīlauea Caldera with commanding 

views in all directions. Continuous visual observation was 

a key component of Jaggar’s monitoring routine, and he 

needed visual line-of-sight to the volcanoes. This visual 

observation is now more consistently achieved with webcams 

and time-lapse photography, but there is still much insight 

to be gained by first-hand human observations of volcanic 

processes when opportune, practical, and safe. Former HVO 

Scientist-in-Charge Donald A. (Don) Swanson makes a cogent 

and eloquent case for the importance of on-site geological 

observations (Swanson, 1992). Increasing use of satellite 

imagery parallels and complements this trend to remote 

monitoring, extending the on-site observations of geologists 

while expanding the number of locations at which monitoring 

can take place.

Significantly increased interoperability and the drive 

toward much more remote monitoring of volcanoes is changing 

the way that U.S. volcanoes are monitored under the National 

Volcano Early Warning System (NVEWS) framework (Ewert 

and others, 2005), with expanding capabilities unforeseen even 

at the end of the 20th century. Indeed, a prime goal of NVEWS 

is that HVO will continue to serve as an important development 

and testing ground for volcano-monitoring and research efforts, 

not only in the United States but also in other countries with 

active or potentially active volcanoes. In so doing, despite its 

geographic insularity in the middle of the Pacific Ocean, HVO 

is now fully integrated into the national-scale monitoring effort 

of the USGS Volcano Hazards Program. If Thomas Jaggar were 

still living today, he doubtless would be utterly amazed, but also 

delighted to see how his creation a century ago has grown and 

thrived, using greatly increased scientific knowledge and a huge 

assortment of monitoring techniques and tools to observe and 

measure volcanoes on Earth and beyond. 

Acknowledgments

As former Scientists-in-Charge of the Hawaiian Volcano 

Observatory (HVO), we (R.I.T. and J.P.K.) readily appreciate 

the dedication and contributions of our HVO colleagues, 

past and present. In writing this introductory chapter, we 

drew heavily from their published scientific papers and 

from informal discussions (“talk story”) with them to learn 

valuable personal impressions regarding the challenges and 

accomplishments during their tenure at HVO. So, to them we 

give our heartfelt thanks for sharing with us their perspectives 

and insights on Hawaiian volcanism. We also wish to thank 

the many friends and colleagues (too numerous to mention 

by name) in the global volcanologic community who have 

shared their knowledge and expertise with us in the pursuit to 

better understand how volcanoes work. We would be remiss 

not to acknowledge the close cooperation and support we 

have always received from Hawai‘i Volcanoes National Park, 

Hawai‘i County and State Civil Defense, and the Hawai‘i 

Pacific Parks Association (formerly called the Hawai‘i Natural 

History Association). 

Earlier drafts of this paper were reviewed by Steven 

Brantley (HVO) and Christina Neal (USGS/Alaska Volcano 

Observatory); we also received constructive informal comments and suggestions from Michael Poland (HVO) and Wendy 

Stovall (USGS/Menlo Park). The reviews by Brantley and Neal 

were exceptionally thorough and incisive, and our responses 

in addressing their concerns and helpful recommendations 

required introduction of new content and (or) clarifications of 

the text discussion that sharpened the “take-home” messages 

we wanted to emphasize in this chapter. Consequently, we 

invited them to join us as co-authors, and they accepted. The 

assistance of editor Poland with the final formatting of the illustrations and text was a huge help, as was the careful checking 

of the references by Taeko (Jane) Takahashi. Last, but not least, 

neither of us could have successfully accomplished our jobs 

without the unwavering support, understanding, and patience of 

our wives, Susan Tilling and Jeri Gertz. 

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Geology, v. 24, no. 6, p. 495–498, doi:10.1130/0091-

7613(1996)024<049.
The viewing tower of the Hawaiian Volcano Observatory, illuminated by moonlight and by the glow from Kīlauea’s summit eruptive vent 

within Halema‘uma‘u Crater. USGS photograph by D. Dzurisin, April 30, 2008.
Jerry P. Eaton, an important figure in the modernization of the seismographs in operation in Hawai‘i. Here, he examines a seismogram on 

a drum recorder in 1960. Photograph courtesy of the Honolulu Star-Advertiser.
1

U.S. Geological Survey.

Abstract

In the century since the Hawaiian Volcano Observatory 

(HVO) put its first seismographs into operation at the edge of 

Kīlauea Volcano’s summit caldera, seismic monitoring at HVO 

(now administered by the U.S. Geological Survey [USGS]) has 

evolved considerably. The HVO seismic network extends across 

the entire Island of Hawai‘i and is complemented by stations 

installed and operated by monitoring partners in both the USGS 

and the National Oceanic and Atmospheric Administration. The 

seismic data stream that is available to HVO for its monitoring 

of volcanic and seismic activity in Hawai‘i, therefore, is built 

from hundreds of data channels from a diverse collection of 

instruments that can accurately record the ground motions 

of earthquakes ranging in magnitude from <1 to >8. In this 

chapter we describe the growth of HVO’s seismic monitoring 

systems throughout its first hundred years of operation. Although 

other references provide specific details of the changes in 

instrumentation and data handling over time, we recount here, in 

more general terms, the evolution of HVO’s seismic network. We 

focus not only on equipment but also on interpretative products 

and results that were enabled by the new instrumentation and 

by improvements in HVO’s seismic monitoring, analytical, and 

interpretative capabilities implemented during the past century. 

As HVO enters its next hundred years of seismological studies, it 

is well situated to further improve upon insights into seismic and 

volcanic processes by using contemporary seismological tools.

Introduction

The U.S. Geological Survey (USGS) Hawaiian Volcano 

Observatory (HVO) has been conducting routine volcano monitoring and continuous measurement programs from the rim of 

Kīlauea Volcano’s summit caldera since 1912. HVO put its first 

seismographs into operation on July 31, 1912, which began 

more than 100 years of seismic monitoring at Kīlauea and on 

the Island of Hawai‘i. Over the past century, thousands of scientific reports and articles have been published in connection 

with Hawaiian volcanism, and an extensive bibliography has 

accumulated, including numerous discussions of the history of 

HVO and its seismic monitoring operations, as well as research 

results. From among these references, we point to Klein and 

Koyanagi (1980), Apple (1987), Eaton (1996), and Klein and 

Wright (2000) for details of the early growth of HVO’s seismic 

network. In particular, the work of Klein and Wright stands 

out because their compilation uses newspaper accounts and 

other reports of the effects of historical earthquakes to extend 

Hawai‘i’s detailed seismic history to nearly a century before 

instrumental monitoring began at HVO. Doing so required that 

they account for seismic monitoring capabilities throughout 

HVO’s history in order to better evaluate historical accounts 

when instrumental records were unavailable or limited.

We present here an updated discussion of the seismic 

monitoring systems at HVO. Rather than casting our discussion 

to include extensive details of instrumentation, our aim is to 

speak to the evolution of HVO’s seismic network and monitoring practices in terms of the capabilities and data collection 

thus afforded by changes to the network. Readers who seek 

greater detail regarding instrumentation can consult the publications mentioned above, as well as the series of HVO weekly, 

monthly, and even annual reports and summaries (see compilations by Fiske and others, 1987, and Bevens and others, 1988).

Early Instrumental Monitoring, 1912–50

To begin systematic seismic monitoring of Kīlauea Volcano in 1912, HVO founder Thomas A. Jaggar, Jr., purchased 

two instruments from Fusakichi Omori of the University of 

Tokyo and installed them in HVO’s Whitney Vault (see Tilling 

and others, this volume, chap. 1). Omori had pioneered seismological research in Japan, including the study of seismicity 

of both volcanic and nonvolcanic origins, and Jaggar’s interest 

was to establish similar observing and monitoring capabilities at 

Kīlauea. One of the Omori instruments installed in 1912 was an 

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 2

The Evolution of Seismic Monitoring Systems at the 

Hawaiian Volcano Observatory

By Paul G. Okubo1, Jennifer S. Nakata1, and Robert Y. Koyanagi1
68 Characteristics of Hawaiian Volcanoes

“ordinary” seismograph, meant to record relatively strong local 

earthquakes; the other was referred to as a “heavy” seismograph 

that, because of its mass, was capable of registering relatively 

weak ground motions, such as those produced by teleseisms.

To set up the seismographs at HVO, Jaggar enlisted 

Harry Wood, who was working at the University of California, 

Berkeley, at the time. Wood arrived in Hawai‘i in summer 1912 

and, in addition to installing and maintaining the seismographs, 

documented the seismic activity recorded by them. Wood left 

HVO in 1917 and, after World War I, was commissioned to 

establish a seismic network in Southern California in cooperation with the California Institute of Technology (Caltech). 

Among Wood’s contributions at Caltech was his construction, 

with astronomer John Anderson, of the Wood-Anderson torsion 

seismometer (Anderson and Wood, 1925). Regional deployment 

of these seismometers led to Charles Richter’s publication of his 

magnitude scale for local earthquakes (Richter, 1935).

At Kīlauea, the heavy Omori seismograph routinely 

and clearly registered microseismic background noise that, 

in Hawai‘i, is closely linked to oceanic swells. In addition, 

the instrument recorded signals that generally correlated 

with visible lava activity at Kīlauea’s summit and that were 

subsequently called “volcanic vibrations” (Wood, 1913). The 

November 25, 1914, eruption of Mauna Loa was preceded 

by about 2 months of seismicity recorded on the two HVO 

seismographs. Although instrumentally determined epicenters 

were not available from the single station that was in operation 

at that time, estimated distances of the recorded earthquakes 

suggested their origin to be beneath Mauna Loa (Wood, 1915).

Like other contemporary instruments, the Omori seismographs were large mechanical devices whose sensitivity to (or 

magnification of) ground movement resulted from their overall 

size (Wood, 1913). The two Omori instruments first brought to 

Hawai‘i did not feature viscous damping. To more accurately 

Omori

seismograph

Bosch-Omori

seismographs

record ground motion, the J.A. Bosch firm of Strasburg modified the original Omori design to add damping, and Jaggar 

subsequently complemented his two Omori instruments with 

Bosch-Omori seismographs (fig. 1). HVO continued to operate 

these Bosch-Omori instruments until 1963 (Apple, 1987).

Seismology in the early 20th century was at an early 

evolutionary stage, with much effort dedicated to collecting 

more and better data by improving the designs of seismometers, 

timing mechanisms, and recording instruments. To promote its 

volcano and seismic monitoring, HVO also undertook its own 

seismograph design and fabrication efforts to achieve greater 

instrument sensitivity and ease of installation and operation. 

The resulting instruments were deployed in Hawai‘i and later 

installed at stations in Alaska and California.

HVO installed additional seismographs as they became 

available, at Hilo in 1919, Kona in 1922, and Hīlea in 1923. The 

distribution of seismic stations in 1923, reflecting the need to 

locate instruments at relatively accessible locations where hosts 

agreed to serve as observers and record changers, is mapped in 

figure 2A. Along with the instruments installed on Kīlauea at 

HVO, these additional stations expanded volcano and earthquake monitoring coverage on the island. HVO staff compiled 

lists of event times, including tremor, local earthquakes, and 

teleseisms, which were reported in HVO’s weekly and monthly 

bulletins (Bevens and others, 1988; Fiske and others, 1987). 

Sizes of local earthquakes, as determined from amplitudes 

measured on available instrumental records, were also provided 

(Klein and Wright, 2000).

A principal—possibly the most basic—goal of seismic 

monitoring was, and continues to be, the accurate cataloging 

and reporting of earthquake time, location, and size. Time reference at the early HVO seismographic stations was provided 

by means of Howard precision astronomical clocks that were 

wired to produce reference marks on the records by lifting the 

recording pens once per minute. Earthquake epicenters were 

determined graphically on the basis of estimates of the respective distances between recording stations and earthquakes, as 

derived from seismic-wave traveltimes. Wood (1914) initially 

selected traveltime-versus-distance curves, compiled by Conrad 

Zeissig for earthquakes and seismic stations in Europe, to provide these estimates for Hawai‘i.

Ruy Finch, who assumed seismological tasks at HVO in 

1919, resumed estimating the distances between earthquakes 

and seismographic stations, which had been abandoned since 

Wood’s departure from HVO in 1917. Finch also changed this 

procedure by using the time-distance tables, compiled and 

published by Omori, which were tabulated over shorter distance ranges and considered to be more suited for application in 

Hawai‘i, to determine earthquake locations (Finch, 1925). 

After working for several years to establish seismographic 

stations and volcano observatories on Lassen Peak in California and in the Aleutian Islands, Austin Jones came to HVO in 

1931 to study the relationships between local earthquakes and 

episodes of ground tilt at Kīlauea. As described by Jones (1935), 

sufficient data had been compiled to produce traveltime tables 

from the data compiled for earthquakes in Hawai‘i, and to more 

Figure 1. Thomas A. Jaggar, Jr., founder of the Hawaiian Volcano 

Observatory, in the Whitney Vault, with Omori and Bosch-Omori 

seismographs at left edge and back, respectively.
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 69

routinely determine and publish earthquake locations from the 

data recorded at HVO’s seismographic stations. The HVO seismic network configuration in 1934, with three stations located 

in the Kīlauea summit caldera region, one in Hilo, and another 

in Kona, is mapped in figure 2B. The Hīlea station was discontinued in 1927 (Jones, 1935), but another station had been added 

at Waiki‘i, on the west flank of Mauna Kea, and the seismic 

network began to approach an islandwide monitoring footprint. 

Jones also initiated a more quantitative means of determining 

earthquake size from the amplitudes of ground motion on seismograms (Klein and Wright, 2000), although he did not account 

156° 155° 

20° 

19° 

Hilo

Kona

HVO

(Whitney Vault)

Hīlea

1923

A

156° 155° 

Uēkahuna

Halema‘uma‘u

Waiki‘i

1934

20° 

19° 

Desert Outlet

Pāhoa

Nā‘ālehu

North Pit

Whitney

Kamuela

1958

MLO

1950

B

C D

0

0

20 40 KILOMETERS

10 20 MILES

for the distances between recording stations and earthquake 

locations, as is required for determining earthquake magnitudes 

on the basis of amplitude.

Throughout the 1940s, HVO’s volcano and seismic 

monitoring capabilities remained stable while being somewhat 

bolstered by attention to instrument design and fabrication. 

Significant to the monitoring was the addition in 1938 of seismic 

station MLO on Mauna Loa’s southeast flank, as shown in figure 

2C. Tabulation of earthquakes and seismic activity continued 

all the while, evolving to the point where graphically estimated 

earthquake locations, including focal depths, were routinely 

Figure 2. Island of Hawai‘i, showing locations of seismographic stations (triangles) operating in 1923 (A), 1934 (B), 1950 (C), and 1958 (D).
70 Characteristics of Hawaiian Volcanoes

being reported. The compilation by Klein and Wright (2000) 

speaks clearly to the quality and commitment of the effort that 

was made in data collection and record keeping. The Hawaiian 

earthquake catalog built by Klein and Wright (2000) included 

more than 16,000 earthquakes, from the time when HVO’s first 

seismometer was put into operation in 1912 through 1953.

More Sensitive Instruments, Data 

Telemetry, Earthquake Magnitudes, and 

Increased Recording Capacity, 1950–70

In the 1940s, HVO staff recognized that seismic monitoring required upgraded capabilities and increased network 

sensitivity, which would improve HVO’s ability to detect, 

reliably locate, and catalog seismicity of interest in Hawai‘i. 

The plan was to first deploy more seismographs, and in 1950 

the first Loucks-Omori instrument, built by HVO machinists 

Burton Loucks and John Forbes, was installed at the Hilo 

station (Pleimann, 1952; Eaton, 1996). As suggested by its 

name, the Loucks-Omori seismograph was a modification of 

the original Omori design. Though still a mechanical system, 

the nominal magnification of ground motion afforded by the 

Loucks-Omori instruments was times-200, somewhat greater 

than that of the Bosch-Omori and other HVO instruments in 

operation at the time.

Hired to take part in HVO’s seismic network expansion 

and modernization, Jerry Eaton arrived at HVO in fall 1953. 

His Ph.D. dissertation, “The theory of the electromagnetic 

seismograph” (subsequently published in brief form as Eaton, 

1957), established him as an authority in the field. Bringing 

this background to HVO, Eaton transformed seismic monitoring by creating the means to expand the HVO seismic network 

to include more stations and, by tuning the instrumentation, to 

be able to record the many small earthquakes related to active 

volcanism in Hawai‘i (Eaton, 1996).

Eaton experimented by recording earthquakes on instruments built to have different frequency-response characteristics. His observations allowed him to specify the design of 

an electromagnetic seismograph, eventually referred to as the 

HVO-1, that was better suited to record small earthquakes 

in Hawai‘i than the Bosch-Omori or the Loucks-Omori 

seismographs. The instrument that resulted from this effort 

was significantly more sensitive; its peak magnification was 

~25,000 at a period of 0.2 s. Seismographs of this design 

eventually replaced older instruments at the Uēkahuna, Hilo, 

Pāhoa, Kamuela, and Nā‘ālehu stations.

Eaton also advanced the real-time seismic monitoring of 

volcanic activity (Eaton, 1996). He simply wanted a means 

to record visibly, as opposed to photographically, small 

earthquakes related to eruptions. Eaton experimented with 

different combinations of available seismometers, preamplifiers, 

and pen recorders until he identified a combination that offered 

the desired sensitivity. During testing of the system, the HVO 

building (where the recorder was on display) was found to be 

too noisy for the sensitive seismometer. The seismometer and 

preamplifier were eventually moved to the Outlet Vault, some 

3 km away from HVO, and the signals were transmitted over 

that distance via cable to the recorder at HVO. This was the 

prototype of a telemetered seismograph, called the HVO-2.

Eaton’s HVO-2 seismometers were more sensitive than 

the HVO-1, with a peak magnification of 40,000 at 0.2 s, and 

they also afforded, by way of telemetry between HVO and 

the seismic stations, the ability to extend monitoring coverage 

over greater distances. Because the available data-transmission 

scheme consisted simply of telephone cables between the 

seismometers and recorders, deployment was restricted to areas 

where it was feasible to establish cabled connections. By 1958 

literally miles of cables extended from HVO to the stations at 

Mauna Loa, Desert, Outlet, and North Pit, all of which were 

equipped with HVO-2 seismometers. With additional installations of Loucks-Omori instruments, an expanded and improved 

HVO seismic network was beginning to take shape (fig. 2D).

Throughout this period of network growth, earthquake 

data processing continued much the same way it had 

previously. Analysts read and compiled seismic-wavearrival times from paper records. Earthquake locations were 

determined graphically, fitting the measured traveltimes 

to those postulated from earthquakes at provisional or 

hypothetical locations among the recording stations. Such 

techniques as implemented at HVO were described by both 

Jones (1935) and Macdonald and Eaton (1964).

In addition to the sensitive seismographs that he designed, 

Eaton also installed standard Wood-Anderson seismographs 

(Anderson and Wood, 1925), initially at his residence in Hawai‘i 

Volcanoes National Park and subsequently at Haleakalā (Maui) 

in late 1956 and at Hilo in 1958. With these instruments, Eaton 

was also able to quantify earthquake magnitudes systematically 

according to the Richter scale, as used with the Wood-Anderson 

seismographs deployed in Southern California (Richter, 1935). 

Although the practice and procedures of determining earthquake 

magnitudes have since evolved and become computerized and 

automated, Eaton’s work can be viewed as the beginning of 

HVO’s production of modern seismicity catalogs, containing 

earthquake hypocenters, origin times, and magnitudes.

The evolution of HVO’s high-gain seismographic network resulted in orders-of-magnitude increases in the numbers of earthquakes recorded, precisely timed, and located 

(Eaton, 1996). Consistent and careful processing of larger 

numbers of earthquakes was necessary to identify temporal 

behaviors and source regions, as well as to infer or associate 

these data with source processes within the magmatic systems 

and volcanic edifices.

Using seismic-wave traveltimes to infer simple, depthvarying seismic-wavespeed distributions and earthquake 

locations derived from the expanded and improved HVO 

seismic network, Eaton was able to articulate a general 

structural model for an idealized Hawaiian volcano. Combining 

these insights with those obtained from deformation (the 

measurement of which was also improved by Eaton; see 
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 71

Poland and others, this volume, chap. 5), geologic studies, 

and geochemical monitoring, Eaton and HVO colleague K.J. 

Murata constructed the first model for the structure and growth 

of a Hawaiian volcano (fig. 3) in their seminal research paper 

titled “How Volcanoes Grow” (Eaton and Murata, 1960). 

Eaton further used the improved seismic data to argue for 

a slight deepening of the crust/upper mantle boundary, or 

Mohorovičić Discontinuity, directly beneath the island. He also 

suggested that the locations of volcanic tremor and swarms 

of deep earthquakes, 60 km directly beneath Kīlauea Caldera, 

highlighted a mantle source feeding the volcano and that the 

ground-surface deformation and earthquake patterns noted 

during eruptions or accompanying inferred magma movement 

suggested a shallow magma reservoir several kilometers beneath 

Kīlauea’s summit caldera (Eaton, 1962).

The growth of HVO’s seismic network over time, 

beginning with its push toward modernization in 1950, 

is plotted in figure 4, along with numbers of station sites, 

recorded ground-motion components, and processed 

earthquakes, as well as indications of HVO’s principal seismic 

data-recording platforms. We begin the curve with data points 

representing earthquakes processed in 1959, marking the 

earliest date for which reading sheets with phase-arrival time 

entries for local earthquakes have been recovered at HVO 

and their arrival-time readings transferred to computer files. 

Although the footprint of HVO’s seismic network evolved 

over the years, earthquake hypocentral solutions dating back 

to 1959 were recomputed to ensure catalog consistency over 

time, across significant changes to HVO’s routine earthquake 

location and magnitude determination procedures. The curve 

ends in 2009, when HVO made significant improvements to 

both its field seismographic stations and its seismic computing 

environment. Data since 2009 are yet to be finalized (as of 

September 2014) with the upgraded capabilities afforded by 

these changes.

Abrupt changes appear in the curves in figure 4 as 

changes affording increased data flow or new stations were 

introduced. Overlap among the different recording/analysis 

platforms, indicated across the bottom of this figure, should 

be understood to reflect HVO’s commitment to preserving 

continuity and consistency of the record where appropriate. 

For the earlier times shown in figure 4, data were written 

exclusively to drums with either smoked or photosensitive 

papers. Throughout the entire interval shown in figure 4, 

HVO continued to record signals from key stations on 

drum papers. The drum recorders showed as much as a full 

day’s continuous and real-time seismic record, affording 

rapid visual recognition and assessment of seismic activity. 

Over time, though still providing critical value to HVO’s 

monitoring and interpretative capabilities, the paper records 

Figure 3. Idealized model of an active Hawaiian volcano, originally published by Eaton and Murata (1960), as inferred from local earthquake 

traveltimes and locations and corroborated by deformation, geologic, and geochemical data.

5

0

-5

-10

-15

-20

-30

-40

-50

-60

Perpendicular to rift zones

Parallel to rift zones

Vertical exaggeration 2x

3.0

4.0

5.3

6.8

8.25

Density

2.77 g/cm

Crust

Mantle Mohorovičić Discontinuity

VP

km/s

Density

g/cm

1.03

4.3 2.3

6.6 2.84

7.9 3.27

Elevation, in kilometers

Sea level

VP

km/s

0

0

20 40 KILOMETERS

10 20 MILES

3

3
72 Characteristics of Hawaiian Volcanoes

gave way to other recording platforms, including the 

Develocorders and the Eclipse and CUSP systems (see 

below), for HVO’s routine seismic data processing and 

earthquake cataloging.

Overall, the number of components grew steadily 

with the number of station sites. A notable exception to this 

trend occurred between 1958 and 1960, when the number 

of recorded components plotted in figure 4 appears to jump 

because new stations featured multiple components, including 

horizontal as well as vertical sensors. As described above, 

some of the stations were equipped with Wood-Anderson 

horizontal seismographs that provided HVO with the capability 

of determining earthquake magnitudes according to Richter’s 

formulation for Southern California (Richter, 1935).

In 1961, Eaton moved to Denver, Colo., to join the 

USGS Crustal Studies Branch and conduct crustal refraction 

and earthquake studies in the western United States. He 

left HVO with its modern seismic instruments telemetering 

data in a truly networked operation. Just as important, he 

provided a template for HVO’s seismic network expansion. 

The years after Eaton’s departure from HVO were a period 

of steady seismic network growth and improvement. 

Seismographic stations were added, both as “outstations” 

that recorded locally at the remote sites and as “networked” 

stations that were connected to recorders at HVO by way 

of upgraded cables laid within and near Kīlauea’s summit. 

By 1967, as plotted in figure 4, HVO was collecting several 

dozen smoked-paper and photographic seismograms daily 

from its seismic network. 

Microearthquake Monitoring Developments at HVO 

and in California

In 1965 Eaton moved from Denver to Menlo Park, Calif., 

to join and help build the USGS’s earthquake research program, 

the creation of which was spurred by the great Alaska earthquake 

of 1964; the group in Menlo Park was particularly focused 

on the San Andreas Fault System. Two related aspects of this 

program were determination of crustal structure and mapping 

of the details of active fault structures, using relatively dense 

networks of sensitive seismometers to record microearthquakes. 

Eaton’s experience with instrument development and network 

deployment at HVO demonstrated the importance of matching 

overall network system response to the earthquakes of interest and 

the region of study. This approach was embraced as an important 

design principle for the microearthquake monitoring efforts 

undertaken in California (Eaton, 1977, 1992).

Eaton and others (1970) demonstrated the value of 

microearthquake recording with closely spaced, high-gain, shortperiod instruments in their study of aftershocks of the June 27, 

1966, Parkfield earthquake that ruptured the San Andreas Fault in 

central California. With sensitive seismometers deployed in close 

proximity to the fault (and, therefore, the earthquakes of interest), 

Eaton and coworkers were able to locate hundreds of aftershocks 

in this sequence. Hypocentral precision was sufficient to map 

fault complexity at depth that could be associated with surface 

observations. Although a planar rupture surface is a reasonable 

approximation for a faulting event, the distribution of hypocenters 

in the aftershock sequence of the 1966 earthquake suggested a 

Figure 4. Evolution of the Hawaiian Volcano Observatory (HVO) seismic network from 1950 through 2009, in terms of numbers of station sites 

(diamonds), recorded ground-motion components (dots), and processed earthquakes (triangles). Plotted across bottom are HVO intervals of seismic 

acquisition and processing systems, with their operational times reflected by lengths of horizontal lines.

0

20

40

60

80

100

120

0

50

100

150

200

250

300

Number of components and sites

Cumulative number of processed earthquakes, in thousands

Develocorder B

Develocorder A

1" Magnetic tape

CUSP

Eclipse Earthworm

1950 1960 1970 1980 1990 2000

EXPLANATION

Components

Sites

Processed earthquakes
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 73

Figure 5. Develocorder seismic-recording equipment used at the 

Hawaiian Volcano Observatory (HVO) from 1967 to 1997. A, Recording 

unit, with photographic system containing film reels and film-developing 

stage at top and lit screen for viewing film shortly after being developed 

and fixed at bottom. B, Offline viewer for Develocorder films, with HVO 

analyst measuring seismic-wave-arrival times.

more complex fault structure in which associated patterns of 

measured fault displacements and tectonic-strain-energy release 

varied, both along fault strike and with depth. The data also 

suggested that the geometric complexity of the San Andreas 

Fault, as indicated by the clustering of aftershock hypocenters, 

affected the dynamic rupture associated with the main shock 

(Eaton and others, 1970). 

As USGS microearthquake monitoring efforts in California progressed, valuable products and insights became 

available to the seismic monitoring community. Eaton and 

colleagues in Menlo Park published numerous reports related 

to USGS’s California seismographic network, essentially 

defining and shaping microearthquake-monitoring practice 

in the United States (see summary discussions by Lee and 

Stewart, 1981). The Menlo Park USGS group, led by Eaton, 

shared seismic network resources with HVO, allowing HVO 

to keep pace with evolving seismic monitoring technologies.

One of the key pieces in the early Menlo Park microearthquake monitoring program was the Develocorder, a 

20-channel photographic recorder that continuously captured 

seismic-trace inputs onto 16-mm reels of microfilm. The 

Develocorder automatically fixed and stored the microfilm so 

that a record showing the 20 channels of data was available 

for viewing within 11 minutes of receiving the data inputs. In 

1967, a Develocorder was delivered to HVO (fig. 5). By the 

end of the year, it was in operation and became a mainstay of 

HVO’s seismic monitoring effort until 1997.

The Develocorder afforded HVO a way to quickly 

and visually determine where an earthquake had occurred. 

A Develocorder display of earthquakes recorded during a 

swarm in Kīlauea’s upper East Rift Zone in 1992 is shown in 

figure 6. The data traces from top to bottom of the film were 

assigned and grouped according to station location. By viewing the times of seismic-wave arrivals at different stations, 

relative to one another, the proximity of the earthquake to the 

stations on display could be visually determined. Depending 

on whether and how the patterns of arrivals changed, possible migration of hypocenters or other changes in seismic 

activity could also be visually assessed.

Though proceeding somewhat deliberately, HVO’s seismic 

network growth and the high rates of seismicity recorded by the 

network produced volumes of individual paper-drum records 

that were awkward, if not difficult, to read, interpret, and store. 

The introduction of the Develocorder and a companion offline 

viewer with a magnifying screen provided a means to efficiently 

scan the record for events and other seismic activity of interest. 

The viewer’s magnification also afforded greater timing precision for measuring seismic-wave arrivals. Arrival times read from 

seismograph drum-paper records were typically read to the nearest 

0.1 s, whereas the Develocorders were read to the nearest 0.05 s. 

With the practice of recording two time references, along with 18 

seismic data traces on both Develocorder films, the traces from 

different stations all shared a common reference time base.

Figure 6. Closeup of Develocorder viewer, showing earthquakes 

recorded during a dike intrusion into Kīlauea’s East Rift Zone in 1992. 

Data traces on Develocorders at the Hawaiian Volcano Observatory 

are arranged by region to allow qualitative visual assessment of 

earthquake locations by looking at arrival patterns among displayed 

data channels.

A

B
74 Characteristics of Hawaiian Volcanoes

The interval throughout the 1960s and mid-1970s shows 

steady increases in the number of station sites and in the 

number of components or channels of data recorded (fig. 4). 

Beginning in the mid-1960s, gradual replacement of cables 

along the ground in favor of radio telemetry made it somewhat 

easier to establish data links. Radio telemetry also facilitated 

expansion of the network into new and remote areas. A second 

Develocorder, operational in 1971, provided recording capacity 

for 18 more seismic data channels.

Computer Processing of Earthquake 

Locations, 1970–85

By 1970, the HVO seismic network had expanded into 

Kīlauea’s rift zones (fig. 7). The increase in the number of 

stations there improved upon the ability to discriminate 

between shallow rift-zone seismicity and deeper earthquakes 

beneath Kīlauea’s south flank. To more appropriately 

and efficiently work with the expanded and improved 

microearthquake datasets, HVO abandoned its graphical 

procedures for manually locating earthquakes in 1970 in 

favor of computer-based earthquake location procedures (for 

example, Endo and others, 1970).

Computer processing of HVO seismic data was a joint 

project between HVO and the USGS National Center for 

Earthquake Research in Menlo Park, Calif. Eaton had written the 

Figure 7. Island of Hawai‘i, showing location of seismographic 

stations operating in 1970.

computer program HYPOLAYR (Eaton, 1969) to locate local 

earthquakes in California. This program featured the calculation 

of P-wave traveltimes between trial hypocenters and recording 

stations from a specified seismic-velocity model composed of 

flat-lying, homogeneous crustal layers, instead of interpolating 

between entries of precompiled traveltime tables, as in other 

programs. HYPOLAYR ran in batch mode, locating earthquakes 

from P-wave arrival-time readings, and computed earthquake 

hypocenters and magnitudes, along with formal errors.

The work involved with the computer locations was 

divided between Menlo Park and HVO. Analysts at HVO 

scanned Develocorder films and paper records for earthquakes, 

identified first-arrival picks, and measured the arrival times on 

Develocorder films and paper records. Reading sheets were sent 

to California, and the arrival times were punched onto computer 

cards. The readings were input to HYPOLAYR, which ran on 

the USGS computer in Menlo Park. Both input card decks and 

output decks created by HYPOLAYR were saved for possible 

further analysis.

The basic product of these efforts was an analystreviewed catalog of Hawaiian seismicity. The start of routine 

computer-based earthquake data analysis is recognized 

in figure 4 as the increase in the number of earthquakes 

processed after 1970. The greater numbers of located or 

processed earthquakes bolstered and added dimensions to 

interpretative discussions of volcanic and seismic processes 

in Hawai‘i. HVO seismicity catalogs were initially published 

quarterly (see Nakata, 2007). After the implementation of 

computer-based hypocentral estimation at HVO, as described 

below, the catalogs were produced as annual summaries of 

seismicity (for example, Nakata and Okubo, 2009).

Koyanagi and Endo (1971) presented the first computerbased HVO compilations for seismic activity recorded in 

1969. Beginning in 1966, HVO had begun to distinguish 

earthquake types beneath Kīlauea’s summit region on the 

basis of the appearance and frequency content of their 

radiated waveforms, and the “long period,” or LP, designation 

for Kīlauea summit earthquakes was added to the HVO 

seismicity catalogs (fig. 8). They noted temporal associations 

of earthquake swarms and of different types of earthquakes 

with eruptive episodes in Kīlauea’s East Rift Zone at Mauna 

Ulu. Clustering of earthquakes, evident in their cross section 

and indicating different event types through different phases 

of eruption, added detail to the general model proposed by 

Eaton and Murata (1960). The earthquake locations and 

classifications subsequently factored into an early model of 

the relation between seismicity, dike intrusion, and seaward 

movement of Kīlauea’s south flank (Koyanagi and others, 

1972) and discussion of the processes related to the 1969 

Kīlauea East Rift Zone eruption (Swanson and others, 1976). 

Koyanagi and others (1976) added interpretative details 

beneath Kīlauea’s summit caldera to the sketch suggested 

by Koyanagi and Endo (1971). Koyanagi and others (1976) 

focused on swarms of earthquakes of different types near 

Kīlauea Caldera to describe the magma conduits and 

reservoirs in that area.

156° 155° 

20° 

19° 

1970

0

0

20 40 KILOMETERS

10 20 MILES
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 75

Figure 8. Seismic activity in 

southeastern part of the Island of 

Hawai‘i in 1969, as originally presented 

by Koyanagi and Endo (1971). A, 

Kīlauea and southeast flank of Mauna 

Loa, showing locations of earthquake 

epicenters, with symbols indicating 

earthquake magnitude. B, Cross section 

projected onto plane parallel to line A-A’ in 

part A, showing locations of earthquakes 

lying within 4 km of line A-A'.

155°30'

19°30'

155°20' 155°10'

19°20'

19°10'

A

A'

EXPLANATION

4.0<M≤5.0

X 3.0<M≤4.0

+ 2.0<M≤3.0

▪ M=2.0

Seismic

station

0

10

20

30

40

Depth, in kilometers

10 20 30

A A'

0

Distance, in kilometers

0

0

5 10 KILOMETERS

2.5 5 MILES

A

B
76 Characteristics of Hawaiian Volcanoes

Ryan and others (1981) constructed a physical model 

to display seismicity throughout Kīlauea as a step toward a 

model of Kīlauea’s magma-transport system. Using event 

classifications as well as hypocentral distributions, they 

identified both a primary conduit feeding magma into Kīlauea 

from the upper mantle and conduits feeding magma into 

Kīlauea’s upper East Rift Zone; they also articulated features 

within Kīlauea’s summit caldera complex. Their resulting model 

offered a multidisciplinary interpretative framework that built 

upon earlier work, such as that of Eaton and Murata (1960).

Though not immediately included in HVO routine data 

processing, a Bell & Howell 3700 FM analog tape recorder 

was brought to HVO in summer 1975 and used to record the 

full complement of telemetered HVO seismic data onto analog 

tapes. HVO did not immediately acquire an accompanying tapeplayback utility, but HVO analysts compiled lists of events of 

interest and sent those, along with the FM tapes, to Menlo Park 

for dubbing or copying selected events onto another FM tape. 

With 14 recording tracks, each capable of storing the equivalent 

of eight seismic channels, the full capacity of the Bell & Howell 

recorder was 112 data channels. This extra capacity was used 

to add more stations and to augment stations with horizontal 

seismometers in order to improve the recognition of seismic 

S-waves that provide added constraint on computed earthquake 

hypocenters. After 1980, relatively few new station sites were 

added to HVO’s seismic network, although the addition of horizontal seismometers at existing sites continued.

In 1979 computer-based earthquake data processing at 

HVO began when HVO acquired a Data General Eclipse computer and equipment to set up FM tape playback and digitizing 

capabilities. With its new platforms to acquire and manipulate 

seismic data, HVO modified its seismic processing routine 

accordingly. Data analysts continued to scan Develocorder 

films to assemble hourly and daily earthquake counts. Instead of 

measuring the arrival times of larger events on the Develocorder 

viewer, however, they compiled lists of events longer than 

40 s—the equivalent of a M1.5 earthquake. Analysts used the 

lists for transferring data from the FM tapes to the computer, 

where waveform data, nominally sampled at 100 Hz, from the 

entire telemetered network were reviewed. Events were picked 

using a graphics computer terminal that registered arrival-time 

or pick information onto the Eclipse computer. Parametric and 

waveform data were subsequently archived to nine-track computer tapes. The timing precision of arrival times measured on 

the graphics terminal improved to 0.01s for all of the telemetered 

data channels.

An important piece of HVO’s Eclipse computing 

configuration was the HYPOINVERSE earthquake location 

program, written by Fred Klein (Klein, 1978). Central to 

HYPOINVERSE and to earthquake location programs in 

general is the method used to solve the problem of 

forward-calculating seismic wave traveltimes through a 

specified model of seismic wavespeeds. The misfit between 

observed and calculated traveltimes based on the locations 

of postulated trial hypocenters and seismic recording stations 

is used to adjust the trial location iteratively until specified 

solution criteria are met. HYPOINVERSE allowed the use of 

crustal seismic-velocity models composed of flat-lying layers 

with horizontal interfaces, in which the seismic velocities are 

specified in terms of depth-varying linear gradients. For such 

gradient models, including the one used at HVO for routine 

earthquake locations (Klein, 1981), fewer parameters are 

required to specify realistic depth-varying seismic-wavespeed 

distributions in the Earth, while maintaining the efficiency 

desired or required for routine hypocentral estimation. 

HYPOINVERSE also accounted for the elevation differences 

between stations by allowing for specification of adjustments, 

called station delays, to be applied to respective arrival times.

HYPOINVERSE also calculated earthquake magnitudes. 

The program included specification of seismograph calibrations for several seismometer and recording system types. Data 

analysts were able to pick maximum phase amplitudes on the 

Eclipse graphics terminal or to enter peak amplitudes, as measured from drum-paper records, intothe HYPOINVERSE phasearrival input files to determine amplitude-based earthquake magnitudes. In addition, analysts measured event coda durations on 

the Develocorder viewing screen as part of their daily scanning 

practice. They entered these durations into the HYPOINVERSE 

calculations to compute coda-duration magnitudes.

HVO’s onsite computer with HYPOINVERSE afforded 

useful flexibility. For episodes of abnormal seismicity, the 40-s 

minimum cataloging threshold was typically relaxed, and the 

intention was to simply locate as many earthquakes as possible. 

On occasion, during volcanic microearthquake swarms, arrivaltime readings from the Develocorder viewers were hand-entered 

into HYPOINVERSE runs to obtain computer-generated 

hypocentral information while the swarm was in progress. Of 

particular interest on these occasions would be estimates of 

earthquake location and focal depth that might reflect shallowing of seismicity before a possible eruption. Together with 

HYPOINVERSE, Klein also developed collections of computer 

utilities for graphical display of earthquakes and manipulation 

of earthquake catalogs (for example, Klein, 1983, 1989a).

Among the numerous reports that have drawn from the 

HVO seismicity catalogs produced from Eclipse data and 

HYPOINVERSE, perhaps none deserves mention more than 

the authoritative compilation by Klein and others (1987) that 

described earthquake distributions at Kīlauea in both space 

and time from 1960 through 1983. From the details observed 

in these patterns, Klein and coauthors systematically described 

individual earthquake swarms related to magmatic intrusions 

and eruptions, and articulated the relationships between 

seismicity and magma movement through Kīlauea’s magma 

system. Koyanagi (1987) offered a similar compilation and 

discussion of Mauna Loa’s magma system.

Summary plots of seismicity, presented in map and cross 

section views of the two volcanoes from these compilations, 

are shown in figures 9 and 10. These plots evoke a comparison 

with those of Eaton and Murata (1960) or Eaton (1962), but 

with the increasing numbers of earthquakes available in these 

later compilations, hypocentral distributions were used to infer 

greater details in the magma systems of the volcanoes. For 
0

10

20

0 10 20 30

Coastline Kīlauea Caldera

Depth, in kilometers

10 20 30

WNW ESE

Kīlauea Caldera Nāpau Crater

Distance, in kilometers

Distance, in kilometers

0 10 20 30

Mauna Iki Kīlauea Caldera

S N SW NE

19°10'

19°20'

19°30'

19°40'

19°

155°40' 155°30' 155°20' 155°10' 155°

0

0

0

10 20 KILOMETERS

5 10 MILES

1

1

2

3

3

2

The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 77

Figure 9. Seismicity on Kīlauea, Island 

of Hawai‘i, between 1960 and 1983, as 

originally presented by Klein and others 

(1987). Epicenters of earthquakes from 

0 to 60 km deep, with boxes labeled for 

reference to cross sections 1, 2, and 3 

(which focus on the upper 25 km of each 

section).
78 Characteristics of Hawaiian Volcanoes

155° 40' 155° 30' 155° 20'

19° 20'

19° 30'

Y

X

X′

Y'

0.0+

1.0+

2.0+

3.0+

4.0+

5.0+

6.0+

5.0+

13.0+

20.0+

0.0+

EXPLANATION

Depth, in

kilometers

Magnitude

X (SW) (NE) X'

10

0

20

30

40

50

60

0 10 20

Depth, in kilometers

Y (NW) (SE) Y'

10

0

20

30

40

50

60

0 10 20

Depth, in kilometers

Distance along transect, in kilometers Distance along transect, in kilometers

A

B C

0

0

5 10 KILOMETERS

2.5 5 MILES

Figure 10. Seismicity beneath Mauna Loa, Island of Hawai‘i, from January 1974 through April 1984, as originally presented by Koyanagi (1987). A, 

Earthquakes beneath Mauna Loa’s summit and adjacent flanks, with epicentral symbols indicating focal depth, and symbol size indicating earthquake 

magnitude. B, Cross section of earthquake hypocenters, projected onto a vertical plane parallel to line X-X′ in part A. C, Cross section of hypocenters 

projected onto a vertical plane parallel to line Y-Y′ in part A. Hypocenters within 5 km of lines X-X′ and Y′-Y′ are included in cross sections.
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 79

example, Klein and others (1987) pointed to gaps in seismicity 

centered 5 km beneath Kīlauea as possible indicators of low 

rigidity that would be consistent with a magma chamber in the 

summit caldera region. They also identified seismicity beneath 

Mauna Loa that extends to depths of 50 km and may represent 

a magmatic root to that volcano. Koyanagi (1987) discussed 

this feature at greater lengths in his study of Mauna Loa’s 

magma system and established that Mauna Loa and Kīlauea 

share similarities in seismicity patterns involving summit and 

rift-zone eruptive features. Seismic network coverage and 

Kīlauea’s higher rates of intrusive and eruptive activity have 

allowed greater refinement of the concepts that he contended 

also apply to Mauna Loa.

Aki and Lee (1976) demonstrated the feasibility of 

using measured local-earthquake seismic-wave arrival times, 

along with cataloged hypocentral locations and origin times, 

to determine the distribution of seismic wavespeeds in 

three dimensions beneath a section of the USGS’s Northern 

California Seismic Network. As their technique was 

embraced and adapted, the procedure, which became known 

as local earthquake tomography, has since been applied in 

various settings. Thurber (1984) was the first to apply local 

earthquake tomography to Kīlauea, using data from the 

HVO seismicity catalog consisting of P-wave-arrival times 

and the locations of 85 earthquakes recorded in and around 

Kīlauea’s summit caldera region in 1980 and 1981. The 

results of P-wave tomographic modeling revealed a region of 

anomalously high VP values interpreted as Kīlauea’s summit 

magma complex, as well as high-velocity anomalies in the 

intrusive cores of Kīlauea’s East and Southwest Rift Zones.

In addition to studies using the parametric hypocentral 

information compiled in the HVO seismicity catalogs, the 

Eclipse recording system preserved digital seismic waveforms, 

played back from the analog FM tape-recording system and 

reviewed to pick arrival times for building HYPOINVERSE 

catalogs. Got and others (1994) used cross-correlation 

techniques on the digital waveforms recorded on the Eclipse 

system between 1979 and 1983 in an early study of the fault 

geometry beneath Kīlauea’s mobile south flank. Rubin and 

others (1998) conducted a similar study of microearthquakes 

associated with the magmatic dike intrusion leading to 

Kīlauea’s currently ongoing East Rift Zone eruption, which 

began in 1983. Identifying earthquakes whose waveforms 

closely correlated with those of other earthquakes established 

multiplets, or families of earthquakes, defined by their highly 

correlating seismic waveforms, occurring in specific regions 

beneath Kīlauea’s south flank (Got and others, 1994) and 

East Rift Zone (Rubin and others, 1998). By determining the 

arrival-time differences among earthquakes in a multiplet from 

waveform cross-correlations, they were able to locate earthquakes within the multiplets to levels of precision significantly 

beyond that of routinely produced HYPOINVERSE catalogs.

From their original selection of earthquakes in the 

HVO seismicity catalogs, Got and others (1994) defined 

earthquake multiplets as families of earthquakes in which 

each event has at least 90 percent coherency with at least 30 

other events in the same family. These criteria resulted in an 

earthquake multiplet containing 252 events, whose hypocenters are shown in cross section in figure 11B. Got and others 

(1994) suggested that a significant fraction of Kīlauea’s 

south-flank seismicity occurs within a relatively thin layer 

or fault at ~8 km below the Earth’s surface. Although they 

analyzed seismicity associated with a rather small part of 

Kīlauea’s south flank, they suggested that the details of fault 

geometries derived from precise relocations of earthquake 

multiplets influence Kīlauea’s geodetically measured deformation and surface-displacement patterns.

Automated, Near-Real-Time, and 

Enhanced Seismic Data Processing at 

HVO, 1985–2009

Although the Eclipse/HYPOINVERSE system 

performed well in Hawai‘i, attempts to use it in Menlo 

Park were less successful, most likely because of the larger 

number of seismographic stations operated by the USGS 

and its monitoring partners in Northern California. Through 

the late 1970s, Carl Johnson at Caltech designed and built 

a system for the seismic network in Southern California, 

for real-time digital data acquisition and automated, nearreal-time event detection and processing on an online 

minicomputer. Interactive data analysis and postprocessing 

were carried out on a second, offline minicomputer (Johnson, 

1979). Johnson joined the USGS and continued working in 

Pasadena to develop an improved seismic processing system, 

which evolved into the Caltech-USGS Seismic Processing 

(CUSP) system (Dollar, 1989; Lee and Stewart, 1989). He 

later served HVO from 1985–89.

As with Johnson’s system at Caltech, CUSP received 

telemetered seismic data signals and converted them into a 

digital stream, using an analog-to-digital (A/D) converter. It also 

passed the digital data through a series of automated processing 

steps beyond A/D, leading to automatic hypocentral estimation 

on event “trigger” files if sufficient numbers of arrival-time 

picks could be assigned automatically, on an online or real-time 

computer system. Then the event triggers were passed to an 

offline analysis and postprocessing minicomputer for interactive 

review on graphics terminals. Readings were added from paper 

records, especially for determining earthquake magnitudes. The 

automation implemented in CUSP obviated the need to play 

back analog tapes in order to produce digital files for analysis 

and archiving, as well as providing seismic analysts with preliminary phase picks and hypocentral coordinates as they began 

their interactive data review. A custom database was integral to 

CUSP’s event-processing and archiving environment.

The Northern California Seismic Network group in 

Menlo Park was optimistic with regard to how CUSP could 

help address its seismic processing needs and, in 1984, set up a 

test CUSP system. HVO was recommended to follow suit as a 
80 Characteristics of Hawaiian Volcanoes

Figure 11. Earthquake relocations beneath Kīlauea, Island of Hawai‘i, modified from Got and others (1994). A, Epicenters of initial selection of 

earthquakes, 1<M<2.5, with hypocenters of 5 to 10 km deep, with readings from more than 12 seismographic stations used in computer location, 

along with HVO stations and geologic and volcanic features. B, North-south cross section of hypocenters of largest multiplet of earthquakes 

consisting of more than 250 events within box plotted in part A. Stars, HVO catalog hypocenters; ellipses, relocated hypocenters, constrained such 

that centroids of HVO catalog and relocated hypocenters are identical. Sizes of ellipses are proportional to uncertainties in relative relocations.

means of modernizing its seismic data acquisition and analysis 

capabilities and procedures by using CUSP’s automated 

utilities. In late 1985, CUSP systems were set up at HVO 

with two Digital Equipment Corp. VAX minicomputers and a 

graphic analysis terminal. At that point, the three largest USGSsupported regional seismographic networks—in Northern 

California (based at Menlo Park), Southern California (at 

Pasadena), and Hawai‘i (at HVO)—had brought CUSP online. 

Adding CUSP’s offline computer at HVO also expanded 

HVO’s general, as well as seismic, computing capabilities.

Implementation and further improvement of CUSP 

proceeded somewhat independently among the various 

USGS CUSP sites. Specific local needs were addressed, and 

procedures were adapted to the local network setting. By 1992, 

CUSP development and support consolidated into a group based 

at the USGS Pasadena office. Among the principal advances 

achieved were the addition of continuous recording of seismic 

traces onto small-format digital tapes and the successful porting 

and implementation of CUSP utilities to take advantage of the 

availability of capable workstation computers configured in 

local-network clusters. Such a configuration greatly enhanced 

the passing of data among real-time acquisition and processing 

computers and analysis workstations. Also noteworthy was 

the creation in the mid-1990s of ISAIAH, or Information on 

Seismic Activity in a Hurry, a utility that would automatically 

issue earthquake notifications by email in near-real time (Wald 

and others, 1994).

With CUSP’s automated earthquake data acquisition 

and processing, HVO was able to catalog greater 

numbers of earthquakes (see fig. 4), although scanning 

of Develocorder films continued for compiling daily and 

hourly counts of earthquakes. To maintain consistency 

with the reporting threshold of M1.5 in earlier published 

seismicity catalogs, earthquakes whose durations read 

from the Develocorders exceeded 40 s were flagged, and 

their arrival times were extracted from the CUSP database. 

Hypocenters and magnitudes for these earthquakes were 

calculated with HYPOINVERSE, maintaining continuity 

with the HYPOINVERSE catalog that began when the 

Eclipse computer was installed in 1979. When the last 

HVO Develocorder failed irreparably in 1997, HVO data 

analysts built the lists of earthquakes for the published HVO 

HYPOINVERSE catalogs by using durations measured from 

drum-paper records.

Working in Menlo Park, Klein continued to support 

and adapt HYPOINVERSE through several updates (Klein, 

1989b, 2002). Across updates of HYPOINVERSE, HVO 

reran older seismicity catalogs to maintain backward 

consistency through all HYPOINVERSE-computed locations 

and magnitudes. HVO also processed phase-arrival data 

dating back to 1959. Ultimately, HVO built an internally 

consistent, HYPOINVERSE-computed catalog of Hawaiian 

seismicity of nearly 150,000 earthquake hypocenters and 

magnitudes spanning 50 years, from 1959 to 2009.

19°30 ́

19°20 ́

19°10 ́

155°30 ́ 155°20 ́ 155°10 ́ 19°21 ́ 19°20 ́30 ̋ 19°20 ́ 19°19 ́30 ̋

NORTH SOUTH

6.5

7.0

7.5

8.0

8.5

9.0

9.5

Depth, in kilometers
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 81

The principal products derived from HVO’s CUSP systems 

continued to be a seismicity catalog and an archive of seismic 

waveforms associated with the cataloged events. Fulfilling 

HVO’s seismic monitoring and cataloging requirements, while 

adding the capability to closely track seismicity in near-real 

time, CUSP also afforded the ability to record and process 

significantly greater numbers of earthquakes in Hawai‘i. 

CUSP’s modular design and the introduction of computer 

workstations allowed tasks, including interactive review and 

timing of earthquakes, to be distributed among a group of 

seismic analysts and computer platforms. In turn, along with 

improved computing capabilities, this distribution of tasks 

motivated further implementation of extended seismic analysis 

techniques that, with repeated or more routine application, 

could be thought of as enhanced or more interpretative seismic 

monitoring products for HVO’s volcanic setting.

Seismic tomography, as pioneered by Aki and Lee (1976), 

is one of the more immediate extensions of a seismicity catalog, 

especially if the seismic-wave-arrival times that are used in 

routine earthquake hypocentral estimation are carefully read and 

confirmed by seismic analysts. Okubo and others (1997) and 

Benz and others (2002) revisited Thurber’s initial tomographic 

P-wave imaging of Kīlauea’s summit caldera region, using early 

HVO CUSP data from 1986 through 1992. In their calculations, 

they included more than 110,000 P-wave-arrival times from 

more than 4,700 earthquakes and expanded the size of the 

tomographically imaged region to include the summit and 

southeastern flank of Mauna Loa (fig. 12; see Denlinger and 

Morgan, this volume, chap. 4).

Other tomographic imaging projects have followed, 

often complementing the arrival-time dataset assembled 

by HVO staff, with arrival times derived from temporary 

Depth = 1 km

Depth = 5 km

Depth = 9 km Depth = 11 km

Depth = 7 km

Depth = 3 km

Mauna Loa

Mauna Ulu

Kīlauea Caldera

Makaopuhi

Maunaiki

Ka‘ōiki

Pu‘ukoa‘e

Kalalua

Mauna Loa

Mauna Ulu

Kīlauea Caldera

Makaopuhi

Maunaiki

Ka‘ōiki

Pu‘ukoa‘e

Kalalua

6.0 6.5 7.0 7.5

P-wave velocity, in km per second

HAWAI‘I

Study area

Figure 12. Selected 

slices from tomographic 

models of Benz and others 

(2002), showing volcanic 

craters and rift zones and 

seismic P-wave-velocity 

(VP) distributions at 

different depths. Zones 

of high velocities are 

interpreted as olivine 

cumulates that resulted 

from repeated cycles of 

intrusion and eruption 

and that may be a driving 

force behind instability of 

Kīlauea’s south flank.
82 Characteristics of Hawaiian Volcanoes

Figure 13. Precise relocations of 1,322 long-period earthquakes that occurred during January 1997–December 1999 beneath Kīlauea, Island of 

Hawai‘i, modified from Battaglia and others (2003). A, Original Hawaiian Volcano Observatory catalog earthquake locations in map view (upper right), 

north-south cross section (left), and east-west cross section (bottom). B, Relocated hypocenters of earthquakes in map view (upper right), north-south 

cross section (left), and east-west cross section (bottom).

-10 -8 -6 -4 -2 0

-10

-8

-6

-4

-2

0

-10

-8

-6

-4

-2

0

-10 -8 -6 -4 -2 0

19°24'

19°26'

155°18' 155°16' 155°14'

B

19°24'

19°26'

155°18' 155°16' 155°14'

A

Depth, in kilometers Depth, in kilometers Depth, in kilometers

Depth, in kilometers

0

0

2 KILOMETERS

1 MILE

1

deployments or surveys targeting more specific imaging targets on and around the Island of Hawai‘i, such as Kīlauea’s 

summit caldera (Dawson and others, 1999), the offshore 

region adjacent to Kīlauea’s south flank (Park and others, 

2009), and Kīlauea’s subaerial south flank (for example, 

Hansen and others, 2004; Syracuse and others, 2010).

The precise relocation of earthquake hypocenters by Got 

and others (1994) is another extension of the basic seismic 

network cataloging products. Following their example of 

relocating hypocenters beneath Kīlauea’s south flank, precise 

relocation analyses that include the cross-correlation of 

seismic waveforms have been applied to other earthquake 

source zones and earthquake types, as identified in the 

HVO catalogs. Striking differences between cataloged and 

subsequently relocated hypocenters afforded very different 

views of seismicity distributions and earthquake processes. 

For example, Battaglia and others (2003) precisely relocated 

hypocenters of two families of LP earthquakes within 

Kīlauea’s summit caldera and showed that one family of LP 

events lies directly beneath the east rim of Halema‘uma‘u 

Crater 500 m below the surface, whereas another lies 

5 km beneath the east rim of Kīlauea caldera northeast of 

Keanakāko‘i Crater (fig. 13). Instead of a diffuse pattern of 

hypocenters suggestive of a volume of LP seismic sources, 

as indicated in the traditional HVO seismicity catalogs, the 

relocated seismicity suggests that the systems of cracks 

comprising the crater rims in these areas may extend as 

deep as 5 km and that LP sources do not fully trace magma 

pathways beneath Kīlauea’s summit caldera. Instead, 

combinations of stress, pressure, temperature, and crack 

geometry in these places are favorable to generating LP 

seismic-energy release (Battaglia and others, 2003).

A similar waveform correlation-based treatment of an 

unusual sequence of deep LP seismicity beneath Mauna 

Loa from 2002 and 2004 established that the distribution 

of cataloged hypocentral solutions, forming a nearly 

vertical source region >20 km in vertical extent, could be 

quite misleading (Okubo and Wolfe, 2008). The relocated 

hypocenters, again using waveform cross-correlations, 

reduced the spread (as routinely located) of these LP 

hypocenters to two clusters, both centered beneath Mauna 

Loa’s summit caldera at depths of 36 and 45 km (fig. 14). 

Rather than tracing a magma pathway from the mantle 

into Mauna Loa, the LP hypocenters reflect regions where 

conditions are favorable to LP seismic energy release.

Systematic relocation of the entire HVO CUSP catalog 

is currently underway, using the waveform cross-correlation 

and cluster analysis techniques described by Matoza and 

others (2013). Rather than targeting specific earthquake 

source regions for relocation, Matoza and coworkers have 

thus far systematically relocated 101,390 of the 130,902 

events spanning the 17-year interval between 1992 and 2009. 

The sharpening of details visible in the relocated earthquake 

distributions along faults, fault streaks, and magmatic features 

will provide important constraints on future studies of fault 

geometry and tectonic and volcanic processes. Research 

continues to include CUSP data from 1986 through 1991 and to 

improve upon the systematic relocation of volcanic LP events.
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 83

 155.7° 155.6° 155.5° 19.3°

19.4°

19.5°

19.6°

20

30

40

50

60

70

Depth, in kilometers

 155.7° 155.6° 155.5° 19.3° 19.4° 19.5° 19.6° 19.3°

19.4°

19.5°

19.6°

Latitude

20

30

40

50

60

70

Depth, in kilometers

19.3° 19.4° 19.5° 19.6°

Latitude

A

B

C

HVO locations

Relocated

 156° 155.75° 155.5° 155.25°

19.25°

19.5°

19.75°

Mauna Loa

Mauna Kea

Kīlauea

Hualālai

AHU

HSS

KII MLO

WIL

POHA

0

0

20 KILOMETERS

10 MILE

10

5

Figure 14. Island of Hawai‘i, showing 

precise relocations of long-period 

earthquakes beneath Mauna Loa, modified 

from Okubo and Wolfe (2008). A, Seismic 

network used for relocations. Open 

triangles, vertical-component stations; 

solid triangles, three-component stations 

(with station names); square, Global 

Seismographic Network station. Box 

shows area covered in parts B and C. B, 

Hawaiian Volcano Observatory catalog of 

earthquake locations below 20-km depth 

between August 2004 and December 2005 

in map view (left) and depth section (right). 

C, Relocated hypocenters of earthquakes 

shown in part B in map view (left) and 

depth section (right).
84 Characteristics of Hawaiian Volcanoes

Beyond Microearthquake Monitoring

Broadband Monitoring

The basic CUSP requirement was to compile regional 

microearthquake catalogs from seismic networks. When CUSP 

was conceived and built, USGS seismic network operations 

were constructed around the USGS high-gain, short-period, 

analog FM, voice-grade telemetry systems designed, built, 

and improved in Menlo Park. CUSP was therefore written to 

drive the A/D conversion of seismic data inputs and had to be 

written to work with specific A/D units. What was originally 

a design specification in terms of the seismic data source 

therefore became, in a sense, a limitation of CUSP as digital 

capabilities became increasingly more accessible to seismic 

data telemetry. CUSP was also designed and principally 

operated in California and Hawai‘i, where relatively large 

USGS seismic networks were recording high levels of seismic 

activity. These aspects also limited CUSP, because its design 

around high-bandwidth A/D conversion, using moderately 

capable minicomputers and workstations, made it difficult to 

adapt CUSP to seismic monitoring efforts that were smaller in 

operational scope, staffing, and (or) support.

Microearthquake cataloging remains an aspect of earthquake and volcano monitoring of the utmost importance. 

Simultaneously, the scope of regional seismic monitoring 

needs to be extended to lower frequencies in the seismic 

spectrum in order to provide information about seismic source 

physics by using broadband seismic instrumentation. Computing technologies, including the Internet and the Global Positioning System (GPS), have developed and matured, allowing 

digital seismic telemetry, recording, and data processing to 

expand and diversify. 

As personal computers (PCs) became increasingly more 

common, seismologists began to explore PC-based computer 

systems for acquiring and processing seismic data, notably 

Willie Lee of the USGS in Menlo Park, who developed what 

has become the IASPEI Seismological Software Library (Lee, 

1989). The capability and value of IASPEI PC systems were 

clearly demonstrated by their use during the Mount Pinatubo 

volcanic crisis in 1991 (Harlow and others, 1996). The USGS 

network group in Northern California was quite interested in 

adding broadband monitoring to its operational scope. Lee 

and colleagues developed a PC software module to match a 

digital seismic telemetry (DST) field unit designed and built 

by USGS engineer Gray Jensen.

HVO, Lee, and USGS Menlo Park colleague Bernard 

Chouet agreed to conduct a broadband seismic monitoring 

field test at Kīlauea in late 1994. Chouet provided the seismometers, borrowing several broadband instruments from 

colleagues to match 10 DST units built by Jensen. Lee brought 

IASPEI PC systems to HVO to record DST seismic data. The 

DST effort evolved from field test to a long-standing HVO 

operational research collaboration with Chouet’s project in 

Menlo Park. The locations of the DST seismographic stations 

in the Kīlauea summit caldera region are shown in figure 15.

Establishing the broadband seismometer subnet within 

Kīlauea Caldera resulted in important improvements to 

HVO’s ability to monitor seismicity in the summit region. 

Adding 10 more seismographic stations within this region 

necessarily meant smaller interstation distances, with 

commensurate improvement in resolving finer details in 

hypocentral distributions and seismic velocity heterogeneity 

beneath the caldera. The Kīlauea broadband seismometer 

subnet served as the anchor for a temporary deployment of 

116 seismometers at Kīlauea in 1996 (McNutt and others, 

1997). Dawson and others (1999) were able to update 

the tomographically derived seismic velocity models of 

Kīlauea’s summit caldera region by using nearly 8,000 Pand S-wave arrival-time readings from 206 earthquakes. 

With the additional stations available for their calculations, 

Dawson and others (1999) cited a tomographic model 

resolution of 500 m, in comparison with the 2- to 3-km 

resolution available to Thurber (1984) and Okubo and others 

(1997). The resulting seismic velocity models for both VP

and VP/VS (fig. 16) suggested distinct but connected magma 

reservoirs beneath the southern part of the summit caldera 

and extending into Kīlauea’s upper East Rift Zone.

Broadband recording at Kīlauea revealed dynamic 

excitation of the volcano that was not previously known because 

it is composed of seismic wave oscillation frequencies of 

0.05 Hz (20 s) or lower (Dawson and others, 1998). This 

frequency range, which is considered to be very-long-period 

(VLP), lies considerably below the traditional 0.5 to 5 Hz LP 

band of volcanic and seismic processes and is poorly recorded 

155°20'

19°22'

19°24'

19°26'

155°18' 155°16' 155°14'

0

0

2 KILOMETERS

1 MILE

1

Figure 15. Kīlauea Caldera, Island of Hawai‘i, showing locations 

of Hawaiian Volcano Observatory (HVO) seismographic stations in 

1996. White triangles, high-magnification, short-period stations with 

data recorded by HVO’s CUSP system; solid black triangles, stations 

equipped with U.S. Geological Survey Digital Seismic Telemetry units 

that were eventually upgraded to broadband seismic sensors.
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 85

Kīlauea Caldera

A H A'

0

5

10

15

0 5 10 20 25 30

Kīlauea Caldera

B H B'

0

5

0 5 10

Depth, in kilometers Depth, in kilometers

19°25'

19°30'

19°20'

155°20' 155°15' 155°10'

A

A'

B' B

3.81 4.81 5.24 5.61 6.04 6.81

P-wave velocity, in kilometers per second

1.50 1.86 2.10 2.31 2.55 3.16

VP/VS, in kilometers per second

0

0

2.5 5 KILOMETERS

1 2 3 MILES

Distance, in kilometers

Distance, in kilometers

Figure 16. Seismic velocity structure of Kīlauea Caldera 

region, Island of Hawai‘i, showing cross sections of VP

and VP/VS from Dawson and others (1999). Reference 

map at top left shows locations of earthquakes used in 

tomographic imaging calculations, as well as locations of 

cross sections A–A′ and B–B′ that are hinged at a point 

adjacent to Halema‘uma‘u Crater within the summit caldera. 

Anomalously low VP and high VP/VS features within the 

summit caldera and extending into Kīlauea’s Southwest 

and East Rift Zones are interpreted to indicate locations of 

magma bodies.

on short-period microearthquake instruments (which made 

up the bulk of the Kīlauea seismic network in the 1990s). 

Taking advantage of the full waveform content of broadband 

recordings, Chouet and colleagues have cataloged and modeled, 

using the moment tensor seismic source representation, VLP 

sources at Kīlauea (for example, Ohminato and others, 1998; 

Dawson and others, 2010), and they continue to develop seismic 

source models to explain the geometry and physics of VLP 

excitation. Most recently, Chouet and Dawson (2011) have 

articulated the geometry of the shallow conduit system feeding 

the current eruption at Halema‘uma‘u Crater. They modeled 

seismic waveforms of tremor bursts related to degassing at 

Halema‘uma‘u by using the seismic velocity model of Dawson 

and others (1999), and suggested that VLP seismic sources 

beneath the east rim of Halema‘uma‘u Crater are associated 

with a complex of intersecting dikes within Kīlauea’s shallow 

magmatic plumbing system (Chouet and Dawson, 2011).

Earthworm and Tsunami Hazard Mitigation

In late 1993, the USGS chose to support the creation 

of a new seismic processing system to address many shared 

concerns identified by seismic network operators both within 

and external to (but supported by) the USGS. USGS software 

developers Alex Bittenbinder and Barbara Bogaert consulted 

with, and ended up collaborating with, Carl Johnson (who 

had moved from the USGS to the University of Hawai‘i, 

Hilo) to design and eventually build and implement the new 

system. Key features of this new system included modularity and scalability to meet the needs of both large and small 

seismic networks, as well as the ability to support data 

exchange among a diversity of digital data sources, ranging from A/D converters to seismic instrumentation or data 

streams from cooperating network partners. Design goals 
86 Characteristics of Hawaiian Volcanoes

were also sensitive to allowing network operators to configure 

and adapt their systems as needed, with sustainability even 

for seismic networks where staffing support is limited. The 

project was named Earthworm (Johnson and others, 1995).

Earthworm was designed as an object-oriented messagebroadcasting system with network utilities built into program 

modules. Interacting modules can be set up on separate 

computers to communicate by using a local area network 

infrastructure, or on the same computer to communicate by 

way of shared memory buffers. Other important decisions were 

to strive toward platform neutrality by initially programming 

for both Sun and Intel computer architectures, to help achieve 

the goal of scalability according to an individual network’s 

needs and resources. The Earthworm system software, except 

for proprietary code specific to particular hardware (like 

software to control an A/D converter), was freely and openly 

available. Earthworm has been embraced by numerous, 

typically smaller, seismic networks across the United States 

and around the world, allowing for the implementation 

of a community-support model for offering and obtaining 

Earthworm technical assistance.

In 1999, upon a recommendation from the National 

Tsunami Hazard Mitigation Program (NTHMP), HVO and the 

National Oceanic and Atmospheric Administration (NOAA) 

Pacific Tsunami Warning Center (PTWC) were given the 

opportunity to improve notification and warning capabilities 

in connection with tsunamis generated by large earthquakes 

in Hawai‘i. Installing Earthworm systems at both centers 

allowed HVO and PTWC to replace an analog FM radio link 

for seismic data with Earthworm data imports and exports 

over the Internet, enabling greater sharing of data between the 

two centers. This implementation allowed HVO to collect all 

of its incoming analog seismic data by using an Earthworm 

A/D converter. The NTHMP upgrades also resulted in the 

installation of three combined broadband and low gain, or 

strong-motion, seismographic stations outside Kīlauea caldera. 

Available Internet bandwidth limited the initial Earthworm 

exchanges to ~40 data channels, which were nonetheless a 

considerable increase over the 8 channels sent by way of the 

earlier FM link. Because no interactive utility had been written 

at that time for Earthworm to efficiently review earthquakes, 

HVO continued to use its CUSP systems for seismic monitoring and cataloging on the Island of Hawai‘i.

Advanced National Seismic System, M6+ 

Earthquakes in 2006

The timing of NTHMP upgrades coincided with that of 

discussions concerned with the seismic monitoring infrastructure of the United States. In 1997 the U.S. Congress enacted 

Public Law 105–47, which required that the USGS conduct 

an assessment of seismic monitoring across the United States, 

with particular focus on improving and updating monitoring 

infrastructure and expanding monitoring capabilities to include 

recording strong ground motions, especially in urban areas. To 

comply with Public Law 105–47, in 1999 the USGS released 

Circular 1188, “Requirement for an Advanced National 

Seismic System: An Assessment of Seismic Monitoring in the 

United States” (U.S. Geological Survey, 1999). In 2000, Congress authorized full implementation of the Advanced National 

Seismic System (ANSS), as framed in USGS Circular 1188, 

over a 5-year period and required an implementation plan.

The first ANSS operational efforts focused on strongmotion monitoring in urban areas exposed to high seismic 

risk. In addition to recording data on strong ground shaking 

in urban areas caused by large local earthquakes, ANSS 

goals and performance were also articulated in terms of the 

generation of data and information products derived from 

recorded ground motion. One of the most important ANSSderived data products is ShakeMap, which predicts the 

ground shaking caused by an earthquake (Wald and others, 

1999a). Although the principal data for ShakeMap came from 

seismographic records, ShakeMap has since been enhanced 

to also use reports of earthquake effects and intensities, as 

compiled on USGS Community Internet Intensity Maps (Wald 

and others, 1999b), in automatically triggered calculation of 

ground-shaking distribution.

Instruments operated by the USGS National Strong 

Motion Project (NSMP, previously the National Strong 

Motion Program) were critical to the creation of ShakeMap. 

The NSMP operates strong-motion accelerometers in 

seismically hazardous regions, including Hawai‘i, to record 

large earthquakes without exceeding the instrumental limits, 

in order to document the effects of large earthquakes on 

structures and the built environment. In 1999, NSMP and 

HVO agreed to upgrade 12 existing NSMP sites on the 

Island of Hawai‘i with digitally recording accelerographs. In 

addition, these sites were to be set up to automatically connect 

and to upload their data by modem to NSMP servers in Menlo 

Park, after earthquakes large enough to satisfy their triggering 

algorithm. Neither HVO’s CUSP nor its Earthworm systems 

were able to acquire and process data from NSMP instruments.

On October 15, 2006, a pair of damaging, M6+ 

earthquakes struck the Island of Hawai‘i. The first event, 

M6.7, was centered ~40 km beneath Kīholo Bay on the 

northwest coast of the island. Just 7 minutes later, an M6.0 

earthquake occurred at a depth of ~20 km, centered north of 

the earlier earthquake, ~20 km offshore from Māhukona. In 

2001, the USGS had implemented its Community Internet 

Intensity Map, or “Did You Feel It?” Web utility (Wald and 

others, 1999b), for earthquakes in Hawai‘i. This utility allows 

reports of earthquake effects to be contributed by the general 

public to USGS servers. The reports are compiled and posted 

as map displays that show distributed earthquake effects. 

Along with data from NSMP digital instruments in Hawai‘i, 

the “Did You Feel It?” reports (fig. 17A) were used in creating 

ShakeMaps of these two earthquakes (fig. 17B) that were 

posted the following day. The ShakeMap utility has since been 

configured to automatically generate ShakeMap products after 

earthquakes of M>3.5 in Hawai‘i.
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 87

Figure 17. U.S. Geological Survey Community Internet Intensity Map (A), also referred to as “Did You Feel It?”, and ShakeMap (B) computed 

after the M6.7 Kīholo Bay earthquake of October 15, 2006.

The 2006 earthquakes were large enough to saturate the 

recordings of HVO’s entire microearthquake monitoring 

system. The NSMP strong-motion instruments, including 

those installed with telemetry as part of the NTHMP 

upgrades, as well as those storing their data only onsite, 

recorded these earthquakes without clipping and provided an 

unprecedented collection of records for a damaging Hawaiian 

earthquake. Yamada and others (2010) used the available 

strong-motion records, along with records of aftershocks that 

defined a northwest-trending inferred rupture plane, in a study 

of the Mw6.7 main shock. Yamada and others determined that 

the rupture propagated northwestward at >3 km/s, and 

inferred that aftershocks clustered about asperities, or regions 

on the rupture plane that sustained relatively large fault slip 

(fig. 18). This sequence was the first opportunity to study the 

rupture process of a damaging earthquake in Hawai‘i in such 

detail, as afforded by the number of onscale strong-motion 

records recovered after the main shock.

The 2006 earthquakes were the largest to strike Hawai‘i 

since the June 25, 1989, M6.1 earthquake near Kalapana, 

along the southeast coast of the Island of Hawai‘i, and thus 

Figure 18. Faulting model of the M6.7 Kīholo Bay earthquake of October 

15, 2006, as inferred by Yamada and others (2010), showing contours (at 

0.1-m intervals) of fault slip projected onto a south-dipping, east-trending 

fault plane defined by aftershock locations (red circles). Aftershocks are 

clustered on edges of asperities, or regions of large fault slip.

-16

-8

0

-32 -24 -16 -8 0 8

Distance along strike (N88E) from main-shock

hypocenter, in kilometers

Distance down dip from mainshock

hypocenter, in kilometers

West East

19 ̊

20 ̊

21 ̊

22 ̊

158 ̊ 157 ̊ 156 ̊ 155 ̊

B

Percieved

shaking

Potential

damage

Peak acceleration,

in percent g

Peak velocity, 

in cm per second

Intensity I II-III IV V VI VII VIII IX X+

>116

>124

Very heavy

Extreme

60–116

65–124

Heavy

Violent

31–60

34–65

Moderate/Heavy

Severe

16–31

18–34

Moderate

Very strong

8.1–16

9.2–18

Light

Strong

3.4–8.1

3.9–9.2

Very light

Moderate

1.1–3.4

1.4–3.9

none

Light

0.1–1.4

0.17–1.4

none

Weak

<0.1

<0.17

none

Not felt

0

0

50 100 KILOMETERS

25 50 MILES

158 ̊ 157 ̊ 156 ̊ 155 ̊

A
88 Characteristics of Hawaiian Volcanoes

provided an opportunity to assess the ability of seismic 

monitoring in Hawai‘i to respond to large earthquakes, using 

modern instrumentation and methods. With respect to their 

principal monitoring missions and requirements, no failures 

occurred among HVO, NSMP, and PTWC with regard to their 

performances in response to these earthquakes. Simultaneously, it was recognized that there was room for a great deal of 

improvement in terms of rapid delivery of reliable, consistent, 

and robust earthquake information (Chock, 2006; William 

Leith, written commun., 2006). The delays may have been 

exacerbated by the shutdown of Internet services resulting 

from power outages across the entire state, which compromised electronic exchange of data and information among 

HVO, PTWC, and the USGS National Earthquake Information 

Center (NEIC) in Golden, Colo., and resulted in difficulties in 

coordinating earthquake reporting.

One recommendation included in the USGS assessment 

in the aftermath of the 2006 earthquakes was to install, at 

HVO, seismic data-processing software implemented by 

the California Integrated Seismic Network (CISN) by the 

USGS offices in both Menlo Park and Pasadena, along with 

their major university operating partners at Caltech and the 

University of California, Berkeley. This installation would 

afford HVO significantly upgraded data acquisition and 

processing capabilities relative to its legacy CUSP systems. The 

CISN software was designed to complement Earthworm data 

acquisition systems collecting digital seismic data from diverse 

instruments. Data from contemporary digital field dataloggers, 

used in conjunction with broadband and low-gain or strongmotion seismic sensors, could be seamlessly merged into 

HVO data-processing streams. In other words, strong-motion, 

broadband, short period, and other seismic data could be 

analyzed concurrently within the CISN processing environment 

to determine earthquake locations and magnitudes. In addition, 

data could be easily shared among HVO, PTWC, NSMP, and 

NEIC to not only meet their respective requirements but also 

to afford earthquake-reporting capabilities consistent with 

expectations and requirements of the ANSS.

The installation of California Integrated Seismic Network 

processing software in Hawai‘i, which was the first attempt to 

migrate the software from its established seismic processing 

environments in California, represented a test of how the 

CISN software could best be ported to other seismic networks. 

Since its initial installation at HVO in 2008, the software has 

been renamed the “ANSS Quake Management Software” 

(AQMS) and installed at USGS-supported regional seismicnetwork processing centers in Washington, Utah, Tennessee, 

New York, and Alaska.

Into HVO’s Second Century of Seismic 

Monitoring: 2009 and Beyond

In April 2009, HVO decommissioned its long-running 

CUSP systems and transferred routine seismic-data-processing 

operations to the AQMS system. In addition to the DST 

seismic network of broadband sensors that continues in 

operation in Kīlauea’s summit caldera (fig. 15), the locations 

of seismographic stations operating on the Island of Hawai‘i 

and monitored by HVO as of 2009 are shown in figure 19. 

With support from the NTHMP, data from PTWC stations 

on the Island of Hawai‘i were also added into HVO’s AQMS 

streams, along with data from the Incorporated Research Institutions for Seismology (IRIS) Global Seismic Network (GSN) 

stations operated at Pōhakuloa Training Area (on the south 

flank of Mauna Kea) and at Kīpapa, O‘ahu (blue squares, fig. 

19). With facilitated acquisition of data from diverse digital 

sources using AQMS, HVO is now also collecting strongmotion seismic data from USGS NSMP stations and USGS 

NetQuakes recorders (installed in the homes of private citizens 

and connected to the Internet), which automatically transmit 

event triggers to HVO through data servers in Menlo Park.

Also in 2009, HVO and the USGS in general were given 

unprecedented opportunities for network modernization with 

the passage of the American Reinvestment and Recovery Act. 

At HVO specifically, significant funds went to an extensive 

overhaul of HVO’s seismic field network, including upgrades 

to seismic sensors and the telemetry needed to transmit data 

from remote station sites to the seismic-data-processing center 

at HVO (Kauahikaua and Poland, 2012). As of 2013, HVO’s 

seismic telemetry backbone is entirely digital, affording sufficient bandwidth and dynamic range to achieve continuous 

broadband seismic monitoring coverage of the entire Island 

of Hawai‘i, as well as the ability to accommodate a greater 

diversity of sensors and allow all earthquakes of interest, 

ranging in magnitude from <1 to >8, to be recorded onscale 

in real time. Establishing such capability extends the possible 

range of seismic source characterization of earthquakes in 

Hawai‘i to include moment tensor characterizations and other 

studies of earthquake rupture processes like that demonstrated 

by Yamada and others (2010). Although study (to the level of 

detail now accessible) of the rupture processes of Hawai‘i’s 

largest historical earthquakes—in 1868 (predating instrumental recording in Hawai‘i) and in 1975 (when local recording 

and analysis were focused on microearthquake locations and 

patterns)—is impossible, we can begin to explore and articulate interactions between future large earthquakes (such as 

those beneath volcanic flanks) and the magmatic storage and 

transport systems of active volcanoes.

A project of considerable interest is analysis of ambient 

seismic noise to detect changes in seismic wavespeed within 

Kīlauea (Ballmer and others, 2013), following the approach 

of Brenguier and others (2008) and Duputel and others 

(2009) at Piton de la Fournaise Volcano, La Réunion. At La 

Réunion, magmatic intrusions were accompanied by small 

but measurable decreases in seismic wavespeed within the 

volcanic edifice, eruptions were associated with increases in 

wavespeed, and a summit caldera collapse was marked by a 

large decrease in wavespeed. Ambient noise monitoring offers 

great promise for tracking changes in seismic wavespeed more 

reliably and consistently over time, because the measurements 
The Evolution of Seismic Monitoring Systems at the Hawaiian Volcano Observatory 89

 155°

20°

19°

 156°

USGS Strong Motion

(NSMP or NetQuakes)

Pacific Tsunami Warning Center

or IRIS GSN station site

Hawaiian Volcano

Observatory station site

EXPLANATION

0

0

20 40 KILOMETERS

10 20 MILES

Figure 19. Map of the Island of Hawai‘i, showing locations of seismographic stations as of 2009.

do not depend on earthquakes, which do not occur at steady 

rates, but rather on background microseismic noise—for 

example, noise caused by ocean waves on the island’s shores. 

Ballmer and others (2013) have implemented ambient-noise 

analyses for HVO data and, in the process of looking for 

changes in seismic wavespeed, have succeeded in identifying 

the contribution to the ambient seismic noise wavefield on 

the Island of Hawai‘i resulting from Kīlauea’s continuing 

eruptions at the summit and East Rift Zone.

HVO, throughout its first century of seismic monitoring, has compiled a comprehensive catalog of earthquakes 

in Hawai‘i and an extensive collection of records of the 

range of seismic events associated with active volcanism 

and tectonism. Implementation of improved monitoring and 

analysis technologies has allowed monitoring products to 

expand beyond simply counts and qualitative descriptions 

of seismic events, as seen on the earliest drum-paper or film 

seismograms, to maps of earthquake locations and estimated 

distributions of strong ground shaking, earthquake magnitudes 

and focal mechanisms, and earthquake notifications, many of 

which are now produced and distributed automatically in nearreal time.

Numerous investigators have drawn from HVO’s digital 

collections of earthquake hypocenters and seismic waveforms 

to apply and develop quantitative analysis and modeling techniques and to extend geophysical and seismologic insights. 

Imaging of seismic-source distributions and three-dimensional 

seismic wavespeed distributions is improving, indicating 
90 Characteristics of Hawaiian Volcanoes

plausible positions of magma bodies and pathways. Moreover, 

by using three-dimensional seismic models of the volcanoes 

with advanced seismic source modeling tools, features of the 

magma transport system that feeds intrusions and eruptions 

can now be seismologically mapped.

A great deal of interest in, and a need for, increasing the 

content of automatically generated, near-real-time earthquake 

reporting products, including more dynamic updating of 

seismicity catalogs, already exists. Thus, it is not unreasonable 

to expect that greater emphasis will be placed on implementing 

such products for volcano monitoring and reporting purposes. 

Earthquake cataloging will continue to be an important 

enterprise, because future research will also depend on the 

availability of reliably reviewed seismicity catalogs. As analysis 

tools and techniques—both mentioned here as well as awaiting 

implementation and demonstration—are more routinely 

applied to data from HVO’s modernized seismic systems, we 

can reasonably expect that routine reporting products from 

seismic networks in volcanically active regions will extend well 

beyond hypocentral location and magnitude to more completely 

describe volcanic and related seismicity.

Acknowledgments

From its very beginnings at HVO, like the observatory 

itself, seismic monitoring has been possible only through the 

dedicated and sustained efforts—both at HVO and at USGS 

offices in Menlo Park and Pasadena, California—of the many 

individuals committed to its continued success. We have 

named several of them as we documented HVO’s seismicnetwork evolution through its first 100 years. Many important 

contributions over that period were not such as would lead 

to authorship on the reports and publications chosen here to 

mark evolutionary network growth. We would now like to 

acknowledge those who contributed from critical staff positions 

at HVO during its transitions through computer-based seismic 

network operations. Their efforts represent the groundwork 

for HVO’s seismic network operation eventually becoming 

a conspicuous regional component of the USGS’ Advanced 

National Seismic System. In approximately chronological 

order of joining HVO’s staff, they are as follows: George 

Kojima, Kenneth Honma, Erwin MacPherson, Irene Tengan, 

Alex Bittenbinder, Thomas English, Robert Cessaro, Bruce 

Furukawa, Wilfred Tanigawa, Alvin Tomori, Gary Honzaki, 

Pauline Fukunaga, Renee Ellorda, Jon Tokuuke, Carol Bryan, 

Laura Kong, Alice Gripp, Steven Fuke, Stuart Koyanagi, 

Jeff Uribe, and David Wilson. In recognition of HVO’s 

long-sustained operation of the USGS CUSP systems, we 

acknowledge vital contributions and support from Robert 

Dollar, Douglas Given, and Allan Walter of the USGS 

Pasadena, Calif., office. Finally, we thank Fred Klein, Wes 

Thelen (since 2011, HVO’s seismic network manager, who can 

begin to write the next chapter of seismic network evolution), 

and Michael Poland for their reviews of the manuscript.

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Abstract

The 19 known shield volcanoes of the main Hawaiian 

Islands—15 now emergent, 3 submerged, and 1 newly born 

and still submarine—lie at the southeast end of a long-lived hot 

spot chain. As the Pacific Plate of the Earth’s lithosphere moves 

slowly northwestward over the Hawaiian hot spot, volcanoes are 

successively born above it, evolve as they drift away from it, and 

eventually die and subside beneath the ocean surface.

The massive outpouring of lava flows from Hawaiian 

volcanoes weighs upon the oceanic crust, depressing it by as 

much as 5 km along an axial Hawaiian Moat. The periphery 

of subsidence is marked by the surrounding Hawaiian Arch. 

Subsidence is ongoing throughout almost all of a volcano’s life.

During its active life, an idealized Hawaiian volcano 

passes through four eruptive stages: preshield, shield, 

postshield, and rejuvenated. Though imperfectly named, these 

stages match our understanding of the growth history and 

compositional variation of the Hawaiian volcanoes; the stages 

reflect variations in the amount and rate of heat supplied to the 

lithosphere as it overrides the hot spot. Principal growth occurs 

in the first 1–2 million years as each volcano rises from the 

sea floor or submarine flank of an adjacent volcano. Volcanic 

extinction ensues as a volcano moves away from the hot spot.

Eruptive-stage boundaries are drawn somewhat arbitrarily 

because of their transitional nature. Preshield-stage lava is alkalic 

as a consequence of a nascent magma-transport system and less 

extensive melting at the periphery of the mantle plume fed by the 

hot spot. The shield stage is the most productive volcanically, and 

each Hawaiian volcano erupts an estimated 80–95 percent of its 

ultimate volume in tholeiitic lavas during this stage. Shield-stage 

volcanism marks the time when a volcano is near or above the 

hot spot and its magma supply system is robust. This most active 

stage may also be the peak time when giant landslides modify the 

flanks of the volcanoes, although such processes begin earlier and 

extend later in the life of the volcanoes.

Late-shield strata extend the silica range as alkali basalt 

and even hawaiite lava flows are sparsely interlayered with 

tholeiite at some volcanoes. Rare are more highly fractionated 

shield-stage lava flows, which may reach 68 weight percent 

SiO2. Intervolcano compositional differences result mainly 

from variations in the part of the mantle plume sampled by 

magmatism and the distribution of magma sources within it.

Volcanism wanes gradually as Hawaiian volcanoes move 

away from the hot spot, passing from the shield stage into the 

postshield stage. Shallow magma reservoirs (1–7-km depth) 

of the shield-stage volcanoes cannot be sustained as magma 

supply lessens, but smaller reservoirs at 20–30-km depth persist. 

The rate of extrusion diminishes by a factor of 10 late in the 

shield stage, and the composition of erupted lava becomes 

more alkalic—albeit erratically—as the degree of melting 

diminishes. The variation makes this transition, from late shield 

to postshield, difficult to define rigorously. Of the volcanoes old 

enough to have seen this transition, eight have postshield strata 

sufficiently distinct and widespread to map separately. Only 

two, Ko‘olau and Lāna‘i, lack rocks of postshield composition.

Five Hawaiian volcanoes have seen rejuvenated-stage 

volcanism following quiescent periods that ranged from 2.0 

to less than 0.5 million years. The rejuvenated stage can be 

brief—only one or two eruptive episodes—or notably durable. 

That on Ni‘ihau lasted from 2.2 to 0.4 million years ago; on 

Kaua‘i, the stage has been ongoing since 3.5 million years ago. 

As transitions go, the rejuvenated stage may be thought of as the 

long tail of alkalic volcanism that begins in late-shield time and 

persists through the postshield (+rejuvenated-stage) era.

Because successive Hawaiian volcanoes erupt over long 

and overlapping spans of time, there is a wide range in the age 

of volcanism along the island chain, even though the age of 

Hawaiian shields is progressively younger to the southeast. 

For example, almost every island from Ni‘ihau to Hawai‘i had 

an eruption in the time between 0.3 and 0.4 million years ago, 

even though only the Island of Hawai‘i had active volcanoes 

in their shield stage during that time.

Once they have formed, Hawaiian volcanoes become 

subject to a spectrum of processes of degradation. Primary 

among these are subaerial erosion, landslides, and subsidence. 

The islands, especially those that grow high above sea level, 

1

Monterey Bay Aquarium Research Institute.

2

U.S. Geological Survey.

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 3

Growth and Degradation of Hawaiian Volcanoes

By David A. Clague1 and David R. Sherrod2
98 Characteristics of Hawaiian Volcanoes

experience mean annual precipitation that locally exceeds 

9 m, leading to rapid erosion that can carve deep canyons in 

less than 1million years.

Hawaiian volcanoes have also been modified by giant 

landslides. Seventeen discrete slides that formed in the past 

5m.y. have been identified around the main Hawaiian Islands, 

and fully 70 are known along the Hawaiian Ridge between 

Midway Islands and the Island of Hawai‘i. These giant 

landslides displace large amounts of seawater to generate 

catastrophic giant waves (megatsunami). The geologic 

evidence for megatsunami in the Hawaiian Islands includes 

chaotic coral and lava-clast breccia preserved as high as 

155m above sea level on Lāna‘i and Moloka‘i.

Large Hawaiian volcanoes can persist as islands through 

the rapid subsidence by building upward rapidly enough. But 

in the long run, subsidence, coupled with surface erosion, 

erases any volcanic remnant above sea level in about 15 m.y. 

One consequence of subsidence, in concert with eustatic 

changes in sea level, is the drowning of coral reefs that drape 

the submarine flanks of the actively subsiding volcanoes. 

At least six reefs northwest of the Island of Hawai‘i form a 

stairstep configuration, the oldest being deepest.

Introduction

Most volcanism on Earth is focused along the global 

network of tectonic plate boundaries, either at the midoceanridge spreading axes or at volcanic arcs located above 

subduction zones. Other volcanism does occur in midplate 

locations and has been attributed to “hot spots” (Wilson, 

1963) or mantle plumes (Morgan, 1972). A primary postulate 

Moloka‘i F.Z.

Clarion F.Z.

Murray F.Z.

Mendocino F.Z.

60°

50°

40°

30°

20°

10°

180 154 148 140 132 127 120 84

Age of ocean floor, in millions of years (Ma)

68 56 48 40 33 20 11 0

A l e u t i a n I s l a n d s

Eight main Hawaiian Islands

Midway Islands

Kure Atoll

Kamchatka

114 mm/yr

85 mm/yr

Pacific Plate

motion vector

HAWAIIAN RIDGE

EMPEROR

SEAMOUNTS

N O RT H

A M

E

R I C A

75.8

 Ma

49.2 Ma

55.5 Ma

200° 180° 160° 140° 120°

Figure 1. Map showing the Hawaiian Islands (red) and Hawaiian–Emperor volcanic chain, most of which consists of submarine 

seamounts, all depicted by their outlines at 2-km water depth (generalized from Clague and Dalrymple, 1987). Vectors indicate 

Pacific Plate motion relative to presumed fixed mantle hot spot in millimeters per year (mm/yr; from Simkin and others, 2006). 

Fracture zones (F.Z.) from Atwater and Severinghaus (1989). Isochrons along Emperor Seamounts chain show age of volcanism 

in millions of years (Ma; Duncan and Keller, 2004). Ocean floor age (Müller and others, 1997) from imagery available on 

EarthByte Web site (http://www.earthbyte.org/Resources/Agegrid/1997/digit_isochrons.html#anchorFTPa0). Mercator projection.
Growth and Degradation of Hawaiian Volcanoes 99

is that these commonly linear chains of midplate volcanoes 

form as the lithospheric plate migrates over a fixed or slowly 

moving magma source in the mantle. Of such chains, that 

which includes Hawai‘i is by far the longest lived and most 

voluminous example. This paper is an attempt to summarize 

the geologic history of these remarkable volcanoes, focused 

on the eight main Hawaiian Islands.

Geologic Setting

The Hawaiian Islands are formed by the youngest 

volcanoes in the chain comprising the Hawaiian Islands, the 

Hawaiian Ridge, and the Emperor Seamounts—an alignment 

of more than 129 volcanoes that stretches across more than 

6,100km of the North Pacific Ocean (fig. 1; Clague, 1996). 

This chain is the type example of an age-progressive, hotspot-generated intraplate volcanic province (see, for example, 

Clague and Dalrymple, 1987; Duncan and Keller, 2004; Sharp 

and Clague, 2006), with the oldest volcanoes, about 81 Ma in 

age (Keller and others, 1995), located east of the Kamchatka 

Peninsula of northeastern Asia. A prominent bend in the chain, 

now dated as a gradual transition occurring from 55 to 45 Ma, is 

located more than 3,550 km west-northwest of Hawai‘i (Sharp 

and Clague, 2006). The Hawaiian-Emperor chain cuts obliquely 

across magnetic lineaments and fracture zones, for the most 

part without regard to preexisting structure of the oceanic crust 

(fig.1; Clague and Dalrymple, 1987), although the Hawaiian 

Ridge is broader and higher near the chain’s intersections with 

the Moloka‘i and Murray Fracture Zones (Wessel, 1993).

The sizes and spacing of the volcanoes are nonuniform. 

Along the Hawaiian leg of the chain, the highest magma flux, 

as total crustal magmatism (probably including some small 

preexisting Cretaceous seamounts), peaked at 18 and 2 Ma 

(near 8 m3/s), whereas the lowest flux (<4 m3/s) was from 

about 48 to 25 Ma (Van Ark and Lin, 2004). The magma 

flux during formation of the entire Emperor chain was low 

(<~4m3/s), with the greatest flux at about 50 Ma. Only 24 

of the entire chain’s volcanoes failed to breach sea level and 

become islands (as well as Lō‘ihi Seamount, which has not 

yet grown to sea level), but few islands were ever large or 

survived as high islands for more than 1–2 m.y. Only eight 

volcanoes northwest of the main islands grew to 1,500 m or 

more above sea level (Clague, 1996).

The volcanic chain represents an enormous outpouring of 

basaltic lava since 81 Ma. The volume of the main Hawaiian 

Islands, accounting for flexural depression of the crust, is 

nearly twice that calculated from bathymetry alone (Robinson 

and Eakins, 2006). Simply applying this correction to the 

entire chain yields a total volume of about 2×106 km3, with 

about 25 percent erupted since 6 Ma to form the present 

Hawaiian Islands. This estimate roughly doubles an early 

estimate (Bargar and Jackson, 1974) that did not account 

for crustal flexure. Moreover, an average extrusion rate of 

3.1 m3/s (0.07 km3/yr) for the 81-m.y. history of the entire 

Hawaiian-Emperor chain can be derived by graphically 

integrating the flux curve of Van Ark and Lin (2004). That rate 

suggests the magmatic volume of the chain is not twice, but 

closer to six times, that calculated from bathymetry alone.

Much of what we know about Hawaiian volcanoes is 

derived from studies conducted by scientists of the Hawaiian 

Volcano Observatory (HVO) on the active Kīlauea and Mauna 

Loa volcanoes. These studies, combined with early work by 

Harold Stearns and Gordon Macdonald at all the main Hawaiian 

Islands, provide the framework for understanding the growth 

and degradation of the older islands in the chain. The present 

review emphasizes results obtained during the past 25 years, 

the period since summaries were published on the occasion of 

HVO’s 75th anniversary (Decker and others, 1987).

Volcano Growth—A Brief History of 

Ideas on Eruptive Stages

An idealized model of Hawaiian volcano evolution 

involves four eruptive stages: preshield, shield, postshield, 

and rejuvenated stages (fig. 2; Stearns, 1946; Clague, 1987a; 

Clague and Dalrymple, 1987; Peterson and Moore, 1987). 

These stages likely reflect variation in the rate at which heat 

is supplied to the lithosphere as the Pacific Plate overrides 

the Hawaiian hot spot (see, for example, Moore and others, 

1982; Wolfe and Morris, 1996a). Although five volcanoes 

of the main islands likely have all four stages present, none 

has products of all stages exposed, in part because preshieldstage lava is commonly buried but also because several of the 

volcanoes have not completed their eruptive development. 

Volcanic extinction follows as a volcano moves away from 

the hot spot. Dissection by large landslides may occur at any 

time in the growth or quiescence of a volcano, and subaerial 

erosion is ongoing whenever the volcano is emergent.

This synoptic view of volcano growth originated with 

Harold Stearns 70 years ago, long before the advent of 

plate tectonic theory (Stearns, 1946). Details have been 

added, subtracted, or rearranged as the timing of events has 

been better established, especially through studies of the 

submarine Lō‘ihi Seamount, the submarine flanks of the 

islands, and substantial mapping, radiometric dating, and 

geochemical analysis of the subaerial and submarine rocks 

that form the volcanoes.

For example, in Stearns’s original assessment, shield 

growth culminated in a caldera-forming stage, whereas 

now it is known that calderas can form and fill repeatedly 

during much of the history of the volcano. The shield stage 

itself was frequently subdivided to emphasize whether 

summit eruptions are occurring mainly in the submarine 

environment, at sea level, or almost entirely subaerially, 

owing to the tendency to discharge effusive or explosive 

eruptive products in the different settings (for example, 

Peterson and Moore, 1987). The capacity for groundwaterdriven explosivity in the volcanic record is widespread until 

the summit grows above the main rain belt on the volcano 
100 Characteristics of Hawaiian Volcanoes

Sea level

Subsidence Subsidence declines

Submarine Subaerial

Ocean floor

Depressed strandl ni e traces Ocean floor

Preshield

Magma chamberMajor structuralfeaturesMajor erosional ormarine-depositionalevents

Clague and Dalrymple (1987) Shield

Submarine Peterson and Moore (1987)

Early submarine Stearns (1946) Subaerial Caldera formation Trachyte and cinder-cone stage Secondary eruptions

Sea-level Subaerial Capping stage Renewed volcanism Atoll

Alkalic postshield Alkalic rejuvenated Erosional stages

Preshield Stage

Substage

Previous usage

Shield Postshield Rejuvenated Erosional stages

Scoria cones ofrejuvenated stagePreshield stage buried

Fringing reef Barrier reef

Submergence Atoll

Landslides

Rift zones

Shallow magma chambers Magma chambers deepen

Calderas

Figure 2. Diagram of the evolution of Hawaiian shield volcanoes as a sequence of generalized growth stages.
Growth and Degradation of Hawaiian Volcanoes 101

35 40 45 50 55 60 65 0

2

4

6

8

10

12

14

16

Na O + K 2 O, in weight percent 2

SiO2, in weight percent

Mugearite

Benmoreite

Tephriphonolite

Phonolite

Trachyte

Phonotephrite

Basanite

Foidite

Picrobasalt

Basalt

Basaltic

andesite

Andesite Tholeiitic

Alkalic

Hawaiite

Classic hawaiite-trachyte postshield-stage sequence

EXPLANATION

Fields for late shield-stage sequence

Mauna Loa (darker) and Kīlauea (lighter) 

fields, showing data symbols

Kohala, Pololū Volcanics

Mauna Kea, Hāmākua Volcanics

Haleakalā, Honomanū Basalt

East Moloka‘i, lower member

(Clague and Dixon, 2000). Such explosivity may persist 

regardless of a volcano’s height, judging from explosion 

debris that mantles the northwest and southeast rims of the 

caldera atop 4,169-m-high Mauna Loa (Macdonald, 1971; 

Trusdell and Swannell, 2003). Groundwater perched in the 

dike swarms of the summit and upper rift zones may be the 

cause of these explosions.

The transitions from alkalic preshield to tholeiitic shield 

to alkalic postshield stage are commonly gradational if defined 

on the basis of chemical composition. For example, many of 

the Hawaiian volcanoes have interbedded alkalic and tholeiitic 

lava flows near the end of the shield stage, as they make the 

transition to the postshield stage (fig. 3). A similar chemical 

transition marks the earlier change from the preshield to the 

shield stage, at least at Lō‘ihi volcano3. These stratigraphic 

complexities raise the question about how best to define the 

stage boundaries. Should the preshield stage end when the first 

tholeiitic shield lavas erupt or when the last of the preshield 

alkalic lavas erupt? Should the end of the shield stage be the 

youngest tholeiitic basalt or the oldest of the postshield alkalic 

lava? In this paper, we emphasize the gradational character of 

these stage boundaries at many volcanoes.

3

Noncapitalized “volcano” is applied informally, whereas capitalization of 

“Volcano” indicates adoption of the word as part of the formal geographic 

name, as listed in the Geographic Names Information System, a database 

maintained by the U.S. Board on Geographic Names. Lō‘ihi Seamount is the 

formal geographic name.

Figure 3. Alkali-silica diagram (Na2O+K2O versus SiO2) composited from several Hawaiian volcanoes. Rock classification 

grid from Le Maitre (2002); shown dashed is boundary separating tholeiitic from alkalic basalt (Macdonald and Katsura, 

1964). Data for Kīlauea and Mauna Loa from Wolfe and Morris (1996b); corresponding fields shown generalized to enclose 

all but a few outlying points. Bold black lines, variously solid, dashed, or dotted, indicate fields for increasingly alkalic late 

shield or transitional postshield basalt and minor hawaiite of Haleakalā, Mauna Kea, East Moloka‘i, and Kohala. Postshield 

lava is commonly even more alkalic, plotting in the field that ranges from hawaiite to trachyte, and encompasses all data 

from Wai‘anae (Pālehua Volcanics), Kohala (Hāwī Volcanics), and Mauna Kea (Laupāhoehoe Volcanics). The several 

exceptions to this fundamental pattern of increasingly alkalic composition across the late shield to postshield stages are 

discussed in the text. Listed in the appendix are the specific data sources for these many fields, on the basis of a published 

Hawaii-statewide whole-rock geochemistry GIS database (Sherrod and others, 2007).
102 Characteristics of Hawaiian Volcanoes

The evolutionary stages, though rooted in geologic 

mapping, are an interpretation of stratigraphic sequences 

imposed after a geologic map is completed. In the submarine 

realm, the assignments necessarily rely on geomorphology and 

petrologic analysis of samples collected by remotely operated 

vehicles or manned submarines. The stage boundaries 

are somewhat arbitrary, because a volcano’s evolution 

is commonly gradational. Two frequently controversial 

transitional periods are (1) the transition from shield to 

postshield stage and (2) the transition from postshield to 

rejuvenated stage.

Of some recent interest is the boundary between 

postshield- and rejuvenated-stage volcanism. An eruptive 

hiatus has long been inferred between the two stages, which 

suggests this boundary could be uniquely defined; indeed, 

the term “rejuvenated” arose to classify the subsequent 

reawakening of a volcano. Stearns (1946) referred to the 

rejuvenated stage as secondary volcanism, and Macdonald 

(1968) called it posterosional, a term that implied that 

substantial time was required to erode the large valleys 

later filled by the youngest volcanic rocks. But radiometric 

dating has shown that little, if any, time elapsed between 

emplacement of volcanic sequences once separated into 

postshield and rejuvenated stages at Wai‘anae (O‘ahu; Presley 

and others, 1997) or Haleakalā (Maui; Sherrod and others, 

2003). All these alkalic lavas are now thought to be of the 

postshield stage. On Kaua‘i, age and chemical data (discussed 

in detail in the section titled “For Kaua‘i, Rejuvenated Stage is 

Gradational from Postshield Stage”) suggest that the boundary 

between postshield and rejuvenated stages is gradational and 

that the eruptive hiatus once used to distinguish the two stages 

is lacking there.

We wrestled with introducing new names for the volcanic 

stages but decided it would only add confusion. Instead, the 

four stages are retained, but with emphasis on the transitional 

or gradational boundaries between sequential stages. As will 

be shown, these transitions vary from volcano to volcano. 

The evolutionary stage model of Hawaiian volcanoes remains 

a robust predictive tool for scientific exploration, but each 

volcano has peculiarities that temper the model’s application.

Some objections persist in the choice of names. For 

example, the shield shape that inspired the name of the 

“shield” stage forms only during the subaerial phase of shieldstage growth, in contrast to steeper slopes built during the 

submarine phase of the shield stage. As used here, shield 

stage includes the entire period when voluminous tholeiitic 

lavas are erupted. The term “preshield” might be interpreted 

as encompassing volcanic growth from inception until the 

development of the subaerial shield shape, but we use the 

term to indicate only the early alkalic part of volcano growth. 

A similar objection might apply to the term “postshield,” 

because postshield alkalic lavas simply veneer the shield 

and so maintain the volcano’s shield shape. Regardless, we 

use postshield to describe alkalic lava that begins erupting at 

the end of the shield stage. The terms “rejuvenated,” “posterosional,” or “secondary” all imply eruption following a time 

gap during which erosion took place. We now know that, at 

least on Kaua‘i, postshield-stage alkalic lava and strongly 

alkalic rejuvenated-stage lava erupted over a lengthy period 

that lacks major time gaps—the earliest rejuvenated-stage lava 

is similar in age to alkalic basalt, hawaiite, and mugearite of 

the postshield stage.

To be clear, the term “shield stage” (as a growth stage) 

should not be equated directly with the term “shield volcano,” 

the term for any broad, typically large volcano. Hawaiian 

volcanoes have long been the archetypal shield volcano.

Brief geographic data for the 17 volcanoes encompassed 

by the main Hawaiian Islands are compiled in table 1. Fifteen 

of those volcanoes are emergent (above sea level), and two 

(Māhukona, Lō‘ihi) are now fully submarine. Also included 

on the list are another two (or three) whose origin as discrete 

volcanoes remains uncertain. Latitude and longitude are 

taken from summit points for most of the volcanoes, although 

approximate caldera centers are included for Kīlauea, Mauna 

Loa, and Māhukona (from topographic maps).

Measuring the Growth of Hawaiian Volcanoes

The rate of growth of Hawaiian volcanoes is typically 

quantified in two ways. Volumetric rates (km3/yr) have been 

described for volcanoes, such as Kīlauea, for which fairly 

precise eruptive volumes have been measured for periods of 

decades or centuries. Also, volumetric rates averaged over 

long time periods have been assigned to a few volcanoes 

where the bulk volume and a fairly good estimate of eruptive 

duration are known. Stratigraphic accumulation rates 

(m/k.y.) are determined by dating sequences of lava flows, 

either in natural exposures or from drill core, which can 

provide substantially thicker sections for sampling. Unlike 

volumetric rates, stratigraphic accumulation rates vary 

widely, simply because of differing geographic distances 

from a volcano’s summit or rift zones, where volcanic 

accumulations are thickest.

Preshield Stage

The earliest growth stage of Hawaiian volcanoes was the 

latest to be discovered, because its products are deeply buried 

in older volcanoes, and substantial technological advances 

were required for sampling in deep water. The preshield stage, 

now known from Lō‘ihi Seamount and Kīlauea, Kohala, and 

perhaps Hualālai volcanoes, appears to be entirely submarine 

and consists of tholeiitic, transitional, alkalic, and strongly 

alkalic lavas. Compositionally transitional volcanic rocks 

were once thought to represent a preshield stage at Māhukona 

volcano because of their high He isotopic ratios (Garcia and 

others, 1990), similar to that of Lō‘ihi lavas, but these rocks 

are now known to be of postshield stage on the basis of 

radiometric ages of about 0.3 Ma (Clague and Calvert, 2009). 

Any preshield-stage strata at Māhukona are probably deeply 
Growth and Degradation of Hawaiian Volcanoes 103

Table 1. Location and summit altitude of volcanoes from Ni‘ihau to Lō‘ihi.

[Geographic coordinates referable to World Geodetic System 1984. Altitude is in meters and feet above mean sea level, as read from topographic maps; negative 

altitudes indicate bathymetric depth. Footnotes explain the variation between ours and other reported onland summit altitudes]

Volcano Longitude Latitude Summit Feature name; topographic map or other reference

 Known volcanoes

1 Ni‘ihau –160.0834 21.9386 392 m; 1,286 ft1 Keanauhi Valley (1989, 1:24,000)1

2 Kaua‘i –159.4974 22.0585 1,598 m; 5,243 ft Kawaikini; Wai‘ale‘ale (1983)

3 Wai‘anae (O‘ahu) –158.1416 21.5072 1,227 m; 4,025 ft Ka‘ala; Hale‘iwa (1983)

4 Ko‘olau (O‘ahu) –157.7881 21.3581 960 m; 3,150 ft North of Kōnāhuanui; Honolulu (1983)

5 West Moloka‘i –157.1570 21.1422 421 m; 1,381 ft Puu Nana; Moloka‘i Airport (1952, 1:24,000)

6 East Moloka‘i –156.8684 21.1065 1,515 m; 4,970 ft2 Kamakou; Kamalo (1968)

7 Lāna‘i –156.8731 20.8121 1,030 (+) m3; 3,379 ft Lāna‘i South (1984, scale 1:25,000)3

8 Kaho‘olawe –156.5715 20.5617 452 m; 1,483 ft Spot elevation west of Pu‘u ‘O Moa‘ula Nui; 

Kaho‘olawe East (1991)

9 West Maui –156.5863 20.8904 1,764 m; 5,788 ft Pu‘ukukui; Lahaina (1992)4

10 Haleakalā –156.2533 20.7097 3,055 m; 10,023 ft Red Hill summit; Kilohana (1983)

11 Māhukona (submarine) –156.1399 20.1315 –1,100 m; 3,610 ft Clague and Moore (1991)

12 Kohala –155.7171 20.0860 1,678 m; 5,505 ft5 Kaunu o Kaleiho‘ohie; Waipio (1916)5

13 Mauna Kea –155.4681 19.8206 4,205 m; 13,796 ft Summit benchmark; Mauna Kea (1982)

14 Hualālai –155.8644 19.6888 2,521 m; 8,271 ft Summit benchmark HAINOA; Hualālai (1982)

15 Mauna Loa summit –155.6054 19.4755 4,169 m; 13,679 ft Benchmark TU0145; Mauna Loa (1981)

Mauna Loa Caldera –155.5920 19.4722

16 Kīlauea summit –155.2868 19.4209 1,269 m; 4,163 ft Uēkahuna Bluff (benchmark TU2382)6

Kīlauea Caldera –155.2839 19.4064

17 Lō‘ihi –155.2601 18.9201 –975 m; –3,199 ft Fornari and others (1988); earlier reports cite slightly 

shallower summits, 969 m depth (Malahoff, 1987) 

and 950-m depth (Carson and Clague, 1995)

 Suspected volcanoes

18 Southwest of Ka‘ena 

Ridge (submarine)

–158.6490 21.7371 Eakins and others, 2003

60 km WNW of O‘ahu; 

may be same location as 

site 18, above

–158.8526 21.6685 About –3,000 m Eruption(?) 1956 C.E.; Macdonald (1959)

19 Penguin Bank(?) 

(submarine)

–157.6488 20.9722 –200 m Carson and Clague (1995); Price and Elliott-Fisk 

(2004); Xu and others (2007a)

1

Ni‘ihau: Highest point is at Pānī‘au benchmark (TU1870), 392 m orthometric altitude (local mean sea level). Previous version of the NGS data sheet reported 

352 m, a typographical error corrected during the preparation of this table. A slightly lower altitude, 381 m, corresponds to the altitude of a spot elevation 

northwest of Kamahakahaka, elsewhere along the coastal bluff (State of Hawaii databook, http://hawaii.gov/dbedt/info/economic/databook/).

2

East Moloka‘i: Summit of Kamakou on Kamalo quadrangle (1968) is marked at the junction of ahupua‘a and appears to be highest point. A slightly lower spot 

elevation is located a distance of 100 m southeast on the Moloka‘i East quadrangle (1983).

3

Lāna‘i: Summit point Lāna‘i hale has no surveyed benchmark but lies within (higher than) 1,030 m contour (Lāna‘i South, 1:25,000, 1984). Commonly cited is 

altitude 1,026 m (3,366 ft), a spot elevation at nearby Ha‘alelepa‘akai (State of Hawaii databook).

4

West Maui: 1992 topographic map has muddy printing of altitude annotations; the 5,788 ft-altitude of Pu‘ukukui resembles 6,788.

5

Kohala, Hawai‘i: Modern topographic maps do not show the summit altitude. A spot elevation 5,505 ft appears on Waipio topographic map (scale 1:62,500, 

surveyed 1911–1913 by R.B. Marshall, published 1916 and reprinted 1951). This altitude also appears in booklet “USGS index to topographic and other map 

coverage.” No spot elevation is shown on the 1:250,000-scale island map or the 30×60-minute topographic map. The summit lies within (higher than) the 

5,480-ft contour (40-ft contour interval; Kamuela topographic quadrangle, 1982), which is a commonly reported altitude slightly short of the summit.

6

Kīlauea: Uēkahuna Bluff (benchmark TU2382) (Miklius and others, 1994).
104 Characteristics of Hawaiian Volcanoes

buried. The transition from the preshield stage to the shield 

stage can be gradual with interbedded alkalic and tholeiitic 

lavas, as described below for Lō‘ihi Seamount, or abrupt, as 

suggested by existing data at Kīlauea Volcano4 (Calvert and 

Lanphere, 2006; Lipman and others, 2002, 2006).

Lō‘ihi Preshield Stage

Lō‘ihi Seamount, newest of the Hawaiian volcanoes, lies 

about 54 km south of Kīlauea Caldera and 975 m below sea 

level (fig. 4). Its summit is about 2.5 km above the adjacent sea 

floor. Lō‘ihi has well-developed rift zones (Moore and others, 

1982; Fornari and others, 1988) and a summit caldera complex 

(Malahoff, 1987; Clague, 2009) with three inset pit craters 

0.6–1.2 km in diameter, similar in scale to Halema‘uma‘u in 

Kīlauea’s summit caldera. The southernmost pit crater formed 

in 1996 during a strong seismic swarm (Lō‘ihi Science Team, 

1997; Davis and Clague, 1998; Garcia and others, 1998; 

overview in Garcia and others, 2006). Also, the summit and 

flanks of Lō‘ihi are scalloped by landslides (Fornari and 

Lō‘ihi

Hilina bench

(Kīlauea preshield)

Hilo Ridge

(Kohala)

Kaua‘i

Ni‘ihau O‘ahu

Lāna‘i

Kaho‘olawe Kohala

Mauna Loa

Mauna

Kea

Moloka‘i

Maui

Māhukona

Mauna Loa

submarine alkalic rocks

others, 1988) that expose lava sequences on the upper east 

flank (Garcia and others, 1995). Thus, Lō‘ihi demonstrates 

that the major structural features of Hawaiian volcanoes, such 

as calderas, rift zones, and flank failures, are well established 

during the preshield stage or in the transition to the shield stage.

Lō‘ihi whole-rock analyses are chiefly tholeiitic 

and alkalic basalt with sparse picrobasalt, basanite, and 

hawaiite (fig. 5). Dredged submarine rocks at Lō‘ihi 

show a compositional trend from alkalic to tholeiitic with 

diminishing age. It was recognized early on that the recovered 

alkalic lavas had thicker palagonite rinds than the tholeiitic 

lavas and were, therefore, likely older, on average (Moore and 

others, 1982). Recent work (Pauley and others, 2011) shows 

that alkalic glass alters more slowly than tholeiitic glass, so 

the alkalic glasses are even older, relative to the tholeiitic 

ones, than previously thought. Stratigraphic sections of lavas 

from two pit craters at Lō‘ihiʼs summit (Garcia and others, 

1993) and along a fault scarp on the east side of the summit 

(Garcia and others, 1995) show that, indeed, on average, the 

alkalic lavas are older than the tholeiitic lavas, albeit with 

considerable overlap. Interbedding of tholeiitic and alkalic 

glass fragments was also seen in an 11-m-thick section of 

volcaniclastic deposits emplaced during the last 5,900 years 

on the southeast part of the summit (Clague and others, 2003; 

Clague, 2009). Alkalic and tholeiitic lava flows exposed in 

Figure 4. Diagrammatic illustration showing oblique aerial view of main Hawaiian Islands. Base illustration courtesy of J.E. Robinson.

4

Of the large volcanoes among the Hawaiian Islands, only Kīlauea has the 

term “Volcano” as part of its formal geographic name and, hence, capital V 

(Geographic Names Information System, http://geonames.usgs.gov/).
Growth and Degradation of Hawaiian Volcanoes 105

Lō‘ihi whole rock

EXPLANATION

Submarine Kīlauea, Hilina bench upslope

Hualālai hawaiite (glass analyses)

Submarine Kīlauea, Hilina bench debris-flow clasts

Kīlauea surface lava flows

Mauna Loa surface lava flows

Kohala rock Hilo Ridge Kohala glass

35 40 45 50 55 60 65 0

2

4

6

8

10

12

14

16

Na O + K 2 O, in weight percent 2

SiO2, in weight percent

Mugearite

Benmoreite

Tephriphonolite

Phonolite

Trachyte

Phonotephrite

Basanite

Foidite

Picrobasalt

Basalt

Basaltic

andesite

Andesite Tholeiitic

Alkalic

Hawaiite

the walls of the summit pit craters are interbedded through 

a much thicker stratigraphic sequence, 300–370 m, and 

presumably represent a longer interval of time (Garcia and 

others, 1993). All these studies indicate that the compositional 

trend (older to younger) is alkalic to tholeiitic as the volcano 

evolves from the preshield to the shield stage.

Eruptions at Lō‘ihi’s present summit (at 975-m water depth), 

and deeper along the volcano’s upper rift zones have been both 

effusive (Moore and others, 1982; Malahoff, 1987; Umino and 

others, 2002) and explosive (Clague and others, 2003; Clague, 

2009; Schipper and others, 2010a,b), suggesting that the early 

stage may include an early deep (chiefly?) effusive substage and 

a later explosive and effusive substage. The explosive alkalic 

eruptions, due to the high magmatic volatile content of the alkalic 

magmas, appear to be mainly from fountains (called poseidic 

eruptions by Schipper and others, 2010a, to distinguish the 

quenching of volcanic fragments in water from that in air) or from 

Strombolian activity (Clague and others, 2003).

Growth rates for Lō‘ihi can be estimated from a 

500-m-thick section of lava flows on the volcano’s east flank. 

Unspiked K-Ar dating yielded ages ranging from about 

100 ka at the base to 5 ka at the top (Guillou and others, 

1997a), suggesting an average lava accumulation rate for the 

preshield stage of about 5 m/k.y. (By way of contrast, shield 

stage rates are 6–16 m/k.y., discussed later.) The K-Ar data 

are vexing, however, because ages from two of the five dated 

samples are inconsistent with their stratigraphic position, 

so we view this accumulation rate cautiously. Radiocarbon 

dating from a foraminifera-bearing volcaniclastic section 

11m thick produced growth rates as high as 3.7 m/k.y. during 

a few millennia (Clague, 2009). Thus, available evidence 

suggests volcanic growth during the preshield stage and the 

earliest shield stage lags somewhat behind the robust upward 

growth during most of the shield stage. Lō‘ihi’s volume now is 

roughly 1,700 km3 (Robinson and Eakins, 2006), and most of 

it apparently formed during the preshield stage.

When will Lō‘ihi breach sea level? This question of 

simple curiosity can be answered only speculatively. The 

volcano may form an island in as little as 50,000 years from 

now (DePaolo and Stolper, 1996). However, with its summit 

975 m below sea level and growth rate of 5 m/k.y., the island’s 

birth year may lie as much as 200,000 years in the future.

Figure 5. Alkali-silica diagram (Na2O+K2O versus SiO2) for preshield Lō‘ihi and Kīlauea whole-rock analyses and 

selected glass analyses from Kohala and Hualālai. Rock classification grid from Le Maitre (2002); shown dashed is 

boundary separating tholeiitic from alkalic basalt (Macdonald and Katsura, 1964). Data from Frey and Clague (1983), 

Hawkins and Melchior (1983), Garcia and others (1995), Sisson and others (2002), Hammer and others (2006), and 

Lipman and Calvert (2011, their electronic data repository appendix DR1).
106 Characteristics of Hawaiian Volcanoes

Kīlauea Preshield Stage

Kīlauea was born less than 300,000 years ago (Calvert 

and Lanphere, 2006). Its preshield-stage lavas range from 

strongly alkalic (including nephelinite) through alkalic basalt 

to transitional and tholeiitic (fig. 5; Lipman and others, 2002; 

Sisson and others, 2002; Coombs and others, 2006; Kimura and 

others, 2006; Lipman and others, 2006). Ages determined by 

40Ar/39Ar dating (Calvert and Lanphere, 2006) on a few samples 

suggest that the transition from preshield to shield stage occurred 

about 150 ka, but the sampling is insufficient to determine the 

temporal span of the transition (as indicated by the stratigraphic 

expanse of interbedded alkalic and tholeiitic lava). Volatiles 

trapped in glasses show that the earliest strongly alkalic lavas 

at Kīlauea erupted subaerially or in shallow water, implying 

that Kīlauea forms only a thin skin on the flank of Mauna Loa 

(Coombs and others, 2006), an idea that originated with Stearns 

and Macdonald (1946, p. 131–136 and plate 1 cross sections) 

and was developed more thoroughly by Lipman and others 

(2006). In contrast to Lō‘ihi Seamount lavas, which have a wide 

range of isotopic ratios and trace element characteristics (Frey 

and Clague, 1983; Staudigel and others, 1984; Garcia and others, 

1993, 1995, 1998, 2006), Kīlauea preshield-stage lavas have 

more uniform isotopic ratios and are interpreted as having been 

derived from a more homogenous source by variable degrees of 

partial melting (Kimura and others, 2006).

Lipman and others (2006) estimate growth rates near the 

end of Kīlauea’s preshield stage of 0.025 km3/yr and a total 

volume of Kīlauea Volcano of 10,000 km3, only 25–66 percent 

of previous estimates. Kīlauea’s alkalic and strongly alkalic 

lavas have an estimated volume of 1,250 km3, and the younger 

transitional basalts an additional 2,100 km3, for a total preshieldstage volume of 3,350 km3 preceding the transition to the shield 

stage (Lipman and others, 2006).

Kohala Preshield Stage

The Hilo Ridge (fig. 4) was long thought to be a rift zone 

of Mauna Kea volcano, but more recently Holcomb and others 

(2000) and Kauahikaua and others (2000) have shown that it 

is a rift zone of Kohala volcano. Lava samples collected from 

the distal Hilo Ridge have 40Ar/39Ar ages of about 1.1 Ma and 

are therefore older than any dated subaerial flows from Kohala 

(Lipman and Calvert, 2011). Compositionally they are tholeiitic 

basalt (fig. 5), although some of the chemical analyses plot close 

to the tholeiitic-alkalic boundary. These latter samples may 

indicate that lava of the preshield stage or earliest shield stage is 

exposed deep on the rift zone of Kohala volcano (Lipman and 

Calvert, 2011).

Hualālai Preshield Stage (?)

Hualālai volcano has a small sliver of alkalic lava exposed 

along a submarine ridge on its lowermost western flank that 

includes volcaniclastic rocks of hawaiite composition, which 

Hammer and others (2006) inferred to represent the alkalic 

preshield stage because of their stratigraphic position. Hammer 

and others (2006) also concluded that the preshield stage at 

Hualālai included a long time period with interbedded alkalic and 

tholeiitic lavas, similar to what was observed at Lō‘ihi Seamount 

(Moore and others, 1982; Garcia and others, 1993, 1995; Clague 

and others, 2003). An alternate interpretation is that these alkalic 

lavas may instead have been erupted during the shield stage, as 

discussed in a later section (“Shield-Stage Alkalic Volcanism”) 

or were perhaps emplaced by slumping of postshield-stage lavas 

(Lipman and Coombs, 2006). Submarine Kohala and Hualālai 

have not been sampled as extensively as Kīlauea or Lō‘ihi, 

so alkalic preshield-stage lava exposed on them may be more 

abundant than our current collections indicate.

Shield Stage

All Hawaiian volcanoes have a shield stage during which 

voluminous eruptions of tholeiitic basalt dominate. The shield 

stage is the most productive volcanically, marking the time when 

a volcano is near the underlying hot spot and its magma system 

is robust. An estimated 80–95 percent of the volcano’s ultimate 

volume is emplaced during this stage. The volumes of most 

volcanoes active since 6 Ma, from Ni‘ihau southeast to Lō‘ihi, 

are compared in figure 6. The pāhoehoe and ‘a‘ā lava flows of 

the ongoing eruption that began in 1983 along Kīlauea’s East 

Rift Zone, 20 km from the volcano’s summit, are characteristic 

of shield-stage volcanism in both style and composition.

The illustrative volume comparison uses the downwardrevised estimate for Kīlauea’s volume (Lipman and others, 2006), 

with the previously estimated volume shown dashed (Robinson 

and Eakins, 2006). Also shown dashed in figure6 is the 

proportion that Kīlauea’s loss would contribute to Mauna Loa’s 

volume. That reassignment, however, may not be warranted, 

because the volumes of overlapping volcanoes are customarily 

calculated by assuming vertical boundaries separating each 

volcano (Robinson and Eakins, 2006). If accuracy is the goal, 

then a bolstered Mauna Loa volume (courtesy of Kīlauea’s onlap) 

should be diminished accordingly by Mauna Loa’s position 

upon the flanks of Hualālai and Mauna Kea, which precede it in 

the volcanic chain. Likewise, the recognition that Hilo Ridge is 

a Kohala rift zone would transfer some Mauna Kea volume to 

Kohala volcano (not shown in figure 6 except for the symbols 

for overestimate and underestimate that accompanied the tabular 

data of Robinson and Eakins, 2006). Revisions like these are a 

reminder that error estimates for large volcano volumes are rarely 

better than 30 percent and commonly worse.

Submerged Volcanoes

At least two, and possibly as many as four, volcanoes 

lie submerged off the coasts of the major Hawaiian islands 

(figs. 4,7). From southeast to northwest, they are Lō‘ihi, 

Māhukona, Penguin Bank, and an edifice on the southwest flank 
Growth and Degradation of Hawaiian Volcanoes 107

Mauna Loa

Haleakalā

Kaua‘i

Wai‘anae

Mauna Kea (+)

Kohala (-)

Ko‘olau

West Moloka‘i

Kīlauea (+)

Kaho‘olawe

East Moloka‘i

Ni‘ihau

Lāna‘i

Hualālai

Māhukona (-)

West Maui

Lō‘ihi (+)

10

9.0

1.7

53.5

57.6

69.8

95.6 74.0

36.4

41.9

30.3

31.6 10.0

34.1

26.3

23.9

21.7

21.1

14.2

13.5

100 90 80 70 60 50 40 30 20 0

100 80 60 40 20 0

Volume, in thousands of cubic kilometers

of Ka‘ena Ridge. Lō‘ihi is best known because of its sporadic 

seismic and eruptive activity; it is discussed more fully 

above, as a volcano in the preshield stage. The other three are 

introduced here. Of the four, only Lō‘ihi and Māhukona are 

widely known to be discrete Quaternary volcanoes, as opposed 

to rift zones of already known volcanoes.

Māhukona

Māhukona lies adjacent to the Island of Hawai‘i 

(fig. 7A). Its summit is at about 1,100-m water depth. A 

small, circular depression may mark a caldera (Clague and 

Moore, 1991). The submarine slope continues upward from 

there, owing to onlap by Hualālai and Kohala volcanoes. 

The age of Māhukona’s inception is unknown but was 

likely about 1.5–1 Ma, on the basis of its present distance 

140 km or so from the Hawaiian hot spot. Dredged samples 

suggest that Māhukona survived at least briefly as a subaerial 

volcano (Clague and Moore, 1991); sunken coral reefs form 

a stairstepping series of terraces, one of the most notable 

geomorphic features of the volcano today (fig. 7A).

Penguin Bank

Penguin Bank is the bathymetric shelf extending 

southwest from West Moloka‘i (fig. 7B). It may be a West 

Moloka‘i rift zone or it may be a separate volcano. In some 

reconstructions, Penguin Bank is the first of the several 

volcanoes that coalesced to form Maui Nui (Big Maui), a 

land mass once larger than the present-day Island of Hawai‘i 

(Price and Elliott-Fisk, 2004). Dredged samples from 

Penguin Bank are subtly distinct, geochemically, from West 

Moloka‘i lava (Xu and others, 2007a), which may further 

substantiate Penguin Bank as a separate volcano. Its summit 

location is chosen to coincide with a closed bathymetric 

contour more or less centered in the western part of the bank; 

a distinct bathymetric saddle separates the shallowest part of 

Penguin Bank from West Moloka‘i.

Southwest Flank of Ka‘ena Ridge

Ka‘ena Ridge is a submarine ridge extending northwest 

from the Island of O‘ahu (fig. 7C). Recently it was proposed 

to be a volcanic feature separate from Wai‘anae volcano 

(Tardona and others, 2011). Two topographically distinct 

shields (labeled 1, 2) form likely eruptive vents on the 

southwest flank of the ridge (Smith, 2002; Eakins and others, 

2003). Whether either of these shields marks the summit of 

a separate volcano or whether they are, instead, related to a 

rift zone of Wai‘anae volcano remains to be resolved. The 

eastern shield, explored during a dive by remotely operated 

underwater vehicle in 2001, consists of ‘a‘ā lava flows and 

rounded beach cobbles of altered vesicular, tholeiitic basalt 

that was erupted and emplaced subaerially (Coombs and 

others, 2004). We indicate this eastern shield as the summit 

of a possible Ka‘ena volcano (fig. 7C), in part because the 

top of this shield is shallower than the one to the west, and 

because its location is nearer the shallow end of the Ka‘ena 

Ridge, which may be a rift zone extending from this summit.

Figure 6. Bar graph of volcano volumes. Shown in order of 

decreasing volume, for all the volcanoes of the eight major Hawaiian 

Islands. Data from Robinson and Eakins (2006) except Kīlauea, 

whose volume is revised downward from their 31,600 km3 (shown 

dashed) to 10,000 km3 (Lipman and others, 2006). The volumetric 

difference, 21,600 km3, is added to Mauna Loa’s volume, increasing it 

to 95,600 km3 (shown dashed). Symbols for likely overestimate (+) and 

underestimate (-) from Robinson and Eakins (2006). For the Mauna 

Kea-Kohala pair, the problem arises from the way in which the volcano 

boundaries partition Hilo Ridge, an offshore Kohala rift zone. For 

Kīlauea, the overestimate applies to the 31,600-km3 volume.
108 Characteristics of Hawaiian Volcanoes

Shield Extrusion Rates

Volumetric Rate

The most precise volumetric rate analyses are for 

Kīlauea Volcano, where the scientific eruption record now 

encompasses 100 years, the written historical record is 

slightly more than twice that, and the oral history of eruptions 

spans a millennium. Basic to the method is the presumption 

that during sustained eruptions, in the absence of deformation, 

the effusion rate is the magma supply rate (Swanson, 1972; 

Dvorak and Dzurisin, 1993). If deformation accompanies an 

eruption, then the deformation data can be recalculated as 

volume change, corresponding to magma storage or discharge 

within the volcanic edifice; these changes are then added to or 

subtracted from erupted volume to calculate the throughput. 

At Kīlauea, the long-term eruption rate for the past 100–

200years, 0.09–0.11 km3 of dense-rock-equivalent magma 

per year, is essentially the same as the long-term rate (0.12 

km3/yr) for the Pu‘u ‘Ō‘o eruption along Kīlauea’s East Rift 

Zone (Heliker and Mattox, 2003), which began in 1983 and 

continues at this writing.

Stratigraphic Accumulation Rate

Shield accumulation rates are based on stratigraphic 

positions and ages of multiple samples. Knowledge of 

stratigraphic separation between samples is commonly 

precise, but ages of samples often have poor precision, 

owing to low potassium content, incipient alteration, and 

sporadic extraneous argon in the tholeiitic lava of the shield 

stage. The derived rates range from 1 m/k.y. to as much as 

16m/k.y., depending on the part of the stratigraphic sequence 

sampled and distance from eruptive vents. Lower rates are 

characteristic of late shield-stage growth far from vents.

S

0

1,000

2,000

3,000

4,000

5,000

6,000

Water depth,

in meters

0

0 10 20

20 40

MILES

KILOMETERS

0

0 10 20

20 40

MILES

KILOMETERS

0

0 10 20

20 40

MILES

KILOMETERS

158° 157°45' 157°30' 157°15′ 157°

156°45′ 156°30' 156°15' 156° 155°45′

159° 158°45' 158°30' 158°15′

20°15′

20°

19°45′

21°15′

21°

20°45′

22°

21°45′

21°30′

A

B

C

Figure 7. Maps showing bathymetry of submerged 

volcanoes and suspected volcanoes along the reach of 

the major Hawaiian Islands. Base is from Eakins and 

others (2003); geographic names from Coombs and 

others (2004). Summit locations are from table 1. A, 

Māhukona, a known shield volcano. B, Penguin Bank, a 

suspected volcano that may only be a rift zone of West 

Moloka‘i volcano. C, Ka‘ena Ridge and topographic 

prominences (1, 2) on its southwest flank; the latter two 

features may form a shield volcano discrete from Ka‘ena 

Ridge (rift zone of Wai‘anae volcano). See Macdonald 

(1959) for discussion of the suspected 1956 C.E. 

submarine eruption.
Growth and Degradation of Hawaiian Volcanoes 109

Rates for Mauna Kea are calculated from dated core in 

the 2.7-km-deep Hawaii Deep Scientific Drilling (HDSD) 

holes (Sharp and others, 1996; Sharp and Renne, 2005). 

Representing Mauna Kea’s mid-to-late shield-stage growth, 

the rates are as great as 8.6 m/k.y. low in the section, 

diminishing upsection to 0.9 m/k.y. in the upper 120 m 

(fig. 8; Sharp and Renne, 2005). The HDSD drill site is 

located 40–45km from Mauna Kea’s center, and the lower 

section may have erupted from Kohala along the Hilo Ridge. 

Decreased rates late in the shield stage, to about 1 m/k.y., 

are also seen in some dated sections from Waiʻanae (O‘ahu; 

Guillou and others, 2000), West Maui (Sherrod and others, 

2007a), and Ko‘olau volcanoes (O‘ahu; fig. 8; Yamasaki and 

others, 2011).

Kīlauea data come from Scientific Observation Holes 

SOH-1 and SOH-4 (Trusdell and others, 1992, 1999), which 

penetrated about 1.7 km of lava flows along the axis of the 

East Rift Zone, at a location about as far from Kīlauea’s 

summit as the HDSD site is from Mauna Kea’s summit. 

These age-depth results have proven difficult to interpret. 

For example, a rate of 3–4 m/k.y. results from fitting a 

curve to the radiometric ages (Guillou and others, 1997b; 

Quane and others, 2000; Teanby and others, 2002), but this 

implies ages between 425 and 565 ka for the bottom of the 

drill hole, significantly older than dated submarine preshield 

alkalic rocks of Kīlauea (Calvert and Lanphere, 2006). A 

substantially higher rate, 16 m/k.y. during the past 45 k.y., 

was obtained by applying a model depth-age curve drawn 

from paleomagnetic inclination and intensity data for the 

upper 800m of strata in SOH-1 (Teanby and others, 2002). 

The higher rate may be reasonable for some episodes of 

shield-stage growth along a rift zone axis but is too high 

to characterize durations of 100 k.y. or longer. Calvert and 

Lanphere (2006) similarly urged caution when interpreting 

the complicated argon geochronologic results from the SOH 

samples and the calculated accumulation rates.

Using surface exposures for Kīlauea, a rate of 6 m/k.y. 

is estimated from outcrops in Hilina Pali, where 275–300 m 

of strata are exposed. The age of those strata is known from 

radiocarbon ages of 28.3 and ~43 ka (D.A. Clague, quoted in 

Riley and others, 1999) and the likely occurrence of the Mono 

Lake (35 ka) and Laschamp (40 ka) geomagnetic excursions 

in lava flows of the Hilina section (data of Riley and others, 

1999, interpreted by Teanby and others, 2002).

Shield-Stage Alkalic Volcanism

Rare submarine alkalic lavas from the base of the 

southeast flank of Kīlauea’s Puna Ridge (Clague and others, 

1995; Hanyu and others, 2005; Coombs and others, 2006) and 

from several cones on the west flank of Mauna Loa volcano 

(Wanless and others, 2006a) are apparently neither preshield 

nor postshield lavas, nor do they appear to have erupted 

during transitions from stage to stage. Emplaced during the 

shield stage, the Kīlauea lava, the youngest submarine lava 

recovered from Kīlauea (based on almost complete lack of 

glass alteration to palagonite), was termed “peripheral” alkalic 

lavas by Clague and Dixon (2000). The alkalic Kīlauea flow 

and Mauna Loa cones are located similar distances from 

the summits of Kīlauea and Mauna Loa, respectively, as 

preshield lavas of Lō‘ihi or postshield stage lavas of Hualālai 

and Mauna Kea are from their respective summits, but in 

directions perpendicular to the orientation of the chain rather 

than in line with the chain. The alkalic lava on the submarine 

west flank of Hualālai, interpreted to be preshield alkalic 

lava (Hammer and others, 2006), may instead be shield-stage 

peripheral alkalic lava. Additional sampling and radiometric 

dating should resolve its origin.

Shield Structure: Rift Zones, Radial Vents, and 

Calderas

Essential features of Hawaiian shield growth have been 

described for nearly a century, although not necessarily 

well understood. The past 50 years have seen increasingly 

sophisticated modeling supported by extensive datasets and 

spaceborne technology.

The intrusive structure of the shield develops as the 

volcano grows from the ocean floor. Magmatic pathways 

within the volcano, including the routes for rift zone intrusions, 

are established in earliest shield-stage time, as indicated by 

the existence of a summit caldera and rift zones on Lō‘ihi 

volcano (Malahoff, 1987; Clague and others, 2003). At the 

active, well-monitored Kīlauea Volcano, seismicity effectively 

tracks the passage of magma from mantle depths of 40–60 km 

(Eaton and Murata, 1960) as it rises into the crust beneath the 

summit area. As the rate of magma supply increases, storage 

reservoirs develop at intermediate depths (near the base of the 

oceanic crust) and shallow depths (1–7km below caldera floor) 

(Clague, 1987a; Ryan, 1987). Magma is shunted from the 

shallow reservoir into the volcano’s rift zones, as documented 

by (1) geodetic and seismic data (Cervelli and Miklius, 2003; 

Klein and others, 1987); (2) CO2 discharge, which is much 

higher from summit vents than from vents on the rift zones 

(Gerlach and Taylor, 1990); and (3) similar trace-element 

concentrations that suggest a shared magma source for lava 

erupted more or less concurrently from both Halema‘uma‘u 

(summit) and Pu‘u ‘Ō‘ō (19 km down the East Rift Zone) 

(Thornber and others, 2010) and even farther down the Puna 

Ridge (Clague and others, 1995).

Rift Zones

A debate still stirs about whether rift-zone intrusions are 

passive or forceful events (Poland and others, this volume, 

chap. 5). As the volcano gains in height, its unbuttressed flank 

or flanks tend to extend by gravitational spreading, which 

favors the development of rift zones by extensional fractures 

and lateral injection of dikes (Fiske and Jackson, 1972; 

Denlinger and Morgan, this volume, chap. 4). But geodetic 
110 Characteristics of Hawaiian Volcanoes

Wai‘anae 

Kepuhi flat-lying

Ka‘ena

Pu‘uma‘ili‘ili

Nānākuli

Kepuhi dipping lava

Waipili

Launiupoko

Waikapū

Honokōhau

Honolua

Pōhākea

50

0

100

200

300

400 0 0.5 1.0 1.5 2.0 2.5 3.0 3.5 4.0

Stratigraphic height downsection, in meters (datum is highest shield-stage or lowest postshield lava)

Age, in millions of years

Mauna Kea West Maui

Mauna

Loa

Mauna

Kea

Transitional

basalt

Tholeiitic basalt

Reversed magnetic

polarity of stratum

controls this point

8.9

0.9

0.3

1.5 0.8

Dated strata

substantially

deeper define

this slope

2.25 2.75

Ko‘olau

Figure 8. Graphs of age versus stratigraphic height for samples from exposed stratigraphic sequences and drill core at selected Hawaiian volcanoes. 

Datum is positioned to approximate the top of the shield stage in each volcano. For Mauna Kea, black symbols, shield-stage strata; gray symbols, 

transitional basalt. For West Maui, gray symbols are postshield lava. Numbers adjacent to some line segments show slope, expressed as meters per 1,000 

years. Sources: Mauna Kea data, Sharp and others (1996); Wai‘anae data, Guillou and others (2000); West Maui data, Sherrod and others (2007a); Ko‘olau 

data, Yamasaki and others (2011). Graphed lines that extend downward beyond data points are based on stratigraphically lower, dated samples (Mauna 

Kea) or remanent magnetic polarity (West Maui); in the latter case, the slope defines the highest permissible accumulation rate for that sequence.

surveys at Kīlauea have documented compressive uplift of the 

outer rift-zone flank over time spans of years; thus, forceful 

intrusion is a necessary component of rift-zone growth 

(Swanson and others, 1976). Doubtless several conditions are 

required for the lateral displacement or dilation of the flanks, 

such as high intrusion rate, low magma viscosity, and low 

fault strength (Dieterich, 1988).

The concept that Kīlauea’s mobile south flank is 

closely related to its rift-zone structure was suggested 

in the 1960s (Moore and Krivoy, 1964), but the analysis 

by Swanson and others (1976) was probably the first to 

depict dilation reaching to the root of the rift zone, as deep 

as 8 km and therefore near the volcano-seafloor interface 

(fig. 9), so that extension is accommodated largely within 
Growth and Degradation of Hawaiian Volcanoes 111

the edifice. Deep, low-angle (4°–5°) faults dipping back 

toward the volcano at such depth (now generally thought to 

be 9–10 km) were recognized first from seismic evidence 

(Ando, 1979; Furumoto and Kovach, 1979). These faults 

may be localized in abyssal sediment (Nakamura, 1982), 

which provides favorable properties of low strength and 

normal or excessive pore-fluid pressure. Alternatively, 

displacement may occur across a zone where “much of the 

displacement is taken up by many local adjustments within 

the pillow [lava] complex” (Swanson and others, 1976, 

p. 25). Lipman and others (1985) proposed, on the basis of 

deformation during the 1975 Kalapana earthquake, that the 

Hilina faults connected to this basal low-angle detachment 

and accommodated deep dilation.

Deep faults are required to model the patterns of 

subsidence and compression from geodetic measurements 

and deep seismic hypocenters along Kīlauea’s rift zones 

(Owen and others, 1995, 2000; Cayol and others, 2000). 

The precise results of the geodetic studies vary, depending 

on the time period studied, but they converge on a model 

in which rift-zone intrusions are blade-like dikes that 

extend upward from the ocean floor-volcano interface 

(now depressed to depths of 9–10km by crustal loading) 

to as shallow as 2–3-km depth, which is the shallow 

zone of frequent seismicity within the rift zone. Deep rift 

expansion, whether by passive or forceful intrusion, is 

interpreted as the primary force driving south-flank motion 

and seismicity (Owen and others, 2000). Deep horizontal 

stress may be increased by the formation of olivine 

cumulates, owing to the additional mass they provide and 

their ability to flow (Clague and Denlinger, 1994).

The quantitative results emphasize a rift zone’s 

magma-storage capacity and the magnitude of ground 

deformation. Historical values for dilation across Kīlauea’s 

East Rift Zone require magma emplacement at rates ranging 

from 0.025 to 0.06km3/yr (Delaney and others, 1993; 

Owen and others, 1995). Add to this the long-term erupted 

volume of 0.12 km3/yr of the ongoing East Rift Zone 

eruption (averaged over the period 1983–2002; Heliker and 

Mattox, 2003), and the resulting total magma-supply rates 

are in the range 0.15–0.18 km3/yr for much or all of the 

past 20 years (Cayol and others, 2000; Heliker and Mattox, 

2003). These values, within the range cited by Dvorak and 

Dzurisin (1993) for periods of years and even centuries, 

are larger by 60 percent than previous estimates for 

magma supply during sustained activity (0.09–0.1 km3/ yr; 

Swanson, 1972; Dzurisin and others, 1984; Dvorak and 

Dzurisin, 1993), owing to the magma stored by interpreted 

deep dike dilation along the East Rift Zone during the past 

two decades.

Mauna Kea appears to lack rift zones, and its eruptive 

vents form a shotgun scatter pattern across the summit 

region. The volcano also lacks submarine features that 

might be early rift zones. Kauaʻi also lacks obvious rift 

zones above sea level but below sea level, at least four 

ridges appear to be rift zones (Eakins and others, 2003).

Radial Vents

Radial vents are well documented on Mauna Loa, where 

linear eruptive fissures on the west and north flanks of the 

volcano trend away from the summit caldera (Lockwood and 

Lipman, 1987). Several examples are found offshore on the 

west flank, including the submarine 1887 C.E. vents (Fornari 

and others, 1980; Moore and others, 1985) and additional vents 

only recently discovered (Wanless and others, 2006b). Other 

Mauna Loa radial vents crop out as far from the summit as Hilo 

(Hālaʻi-Puu Honu vent alignment; Buchanan-Banks, 1993) and 

in the saddle between Mauna Loa and Mauna Kea. Radial vents 

are probably a consequence of spreading across the arcuate 

Northeast and Southwest Rift Zones of Mauna Loa, which 

causes extension on the west and north sides of the volcano. 

Dikes propagate into this wide extensional zone. Radial vents 

have not been identified on Kīlauea.

On older volcanoes, dikes are commonly the only 

exposed clues to the existence of radial vents. At Waiʻanae 

volcano (O‘ahu), dikes along a southeast-trending rift zone 

follow a radial pattern as they swing around the south side of 

the caldera, perhaps in response to the stress field created by 

caldera growth in late shield time (Zbinden and Sinton, 1988). 

Scattered dikes of almost all orientations have been mapped 

on Kaua‘i (Macdonald and others, 1960). Most are near the 

volcanic center, where varied local stresses may predominate.

Calderas

Calderas are common on Hawaiian shield volcanoes. 

Their topographic rims crown the summits of the active 

Kīlauea and Mauna Loa. Erosion has exposed their shallow 

or mid-depth reaches at older volcanoes, such as Kaho‘olawe, 

West Maui, East Moloka‘i, Ko‘olau, Waiʻanae, and perhaps 

Kaua‘i. Once thought to form only late in the shield stage, 

calderas are now known to appear as early as the preshield 

stage (for example, Lō‘ihi) and then likely develop and fill 

repeatedly throughout the shield stage. Thus, calderas go 

hand-in-hand with the high rate of magmatism that builds the 

Hawaiian Islands.

The repeated injection of magma into a volcano’s summit 

region results in a complex magma reservoir: a nexus of dikes, 

sills, and plugs. Over time this complex develops a greater 

bulk-rock density than the adjacent volcano, owing to the 

nonvesicular, massive character of the intrusions and crystal 

accumulation in the deeper parts of the reservoir. In general, 

Hawaiian calderas are coincident with a positive gravity 

anomaly centered at the point from which the rift zones radiate, 

owing in large part to their position above the dense magma 

column that extends downward through the crust and into 

the mantle (fig. 10) (Kauahikaua and others, 2000). In older 

calderas, such as Ko‘olau (O‘ahu), where the reservoir is fully 

crystallized, the resulting seismic signature includes P-wave 

velocities (VP) of 7.7 km/s at depths less than 2 km (Adams and 

Furumoto, 1965; Furumoto and others, 1965), contrasting with 

the low VP of 4.6 km/s in the surrounding material of the shield.
112 Characteristics of Hawaiian Volcanoes

Figure 10. Maps showing calderas, rift zones, and Bouguer gravity anomalies at selected 

Hawaiian volcanoes, all at same scale. A, Kīlauea Volcano, Island of Hawai‘i. Dark brown 

fill, modern topographic caldera; lighter fill, expanse of caldera defined by outer ring faults. 

Structural elements shown by faults (black lines). Erosion is too shallow to provide much 

structural sense from the few dike exposures. Gravity contours from dataset of J.P. Kauahikaua 

(written commun., 2011); gravity maximum, about 330 mGal, is coincident with caldera. B, 

Ko‘olau volcano and caldera, Island of O‘ahu. Note that illustration is rotated 96° clockwise to 

facilitate comparison with the depiction of Kīlauea Volcano in A. Shown by dark brown fill is 

caldera defined by Stearns (1939); an outer boundary (lighter fill) is inferred from map data of 

Walker (1987). Rift-zone dike orientation is generalization used by Walker (1987). Large gaps 

in downrift dike progression correspond to interfluves between canyons, where lava flows at top 

of volcano bury dike exposures. Bouguer gravity anomalies from Strange and others (1965); 

gravity maximum about 310 mGal. C, Kaua‘i volcano (all of the Island of Kaua‘i). Caldera 

boundary, faults, and dikes from Macdonald and others (1960). Bouguer gravity contours from 

Krivoy and others (1965). These contours match the more recent work by Flinders and others 

(2010), whose graphical depiction does not lend itself readily to presentation as isogal lines.

Oceanic crust

Inactive Active

0

-2

2

N S

N S

-4

-6

-8

-10

-12

-14

East Rift Zone

Mantle 

Sea level Altitude, in kilometers

Mauna Loa Kīlauea

PACIFIC

OCEAN

A

0

-2

2

Sea level

-4

-6

-8

-10

-12

Altitude, in kilometers

Coulomb

stress, in MPa

3

2

4

1

0

-1

-2

-3

-4

M6.1, 1989 M7.2, 1975

Locked

B

0 5 10 15 KILOMETERS

0 5 10 MILES

310

300

290

280

270

Complete Bouguer gravity

anomaly, in mGals

M6.1, 1989 M7.2, 1975 Creep

Figure 9. Cross sections showing 

structure and geodetic model of a Kīlauea 

rift zone. A, Profile from Swanson and 

others (1976, their A-A′) shows middle 

East Rift Zone of Kīlauea Volcano. No 

vertical exaggeration. Concept of lowangle fault at volcano-ocean floor interface 

had not yet been proposed, so forceful 

expansion of the rift zone was envisioned 

occurring along dispersed low-angle faults 

within the pillow lava complex. Division 

of rift zone dikes into inactive and active 

part originates from asymmetrical gravity 

profile and other evidence that rift-zone 

axis propagates southward through time 

as a consequence of dilation (Swanson 

and others, 1976). Gravity profile above 

section is from dataset of J.P. Kauahikaua 

(written commun., 2011); mGals, milligals. 

B, Geodetic model and Coulomb stress 

calculations depicted by Cayol and others 

(2000), at roughly the same geographic 

position and scale as cross section in 

panel A. Also shown from the original 

illustration are focal depths of two large 

south-flank earthquakes and the division 

of the deep fault into locked and creeping 

segments. Contrasting colors for arrows 

indicating dike expansion are solely for 

visibility. MPa, megapascals.
Growth and Degradation of Hawaiian Volcanoes 113

0 5 10 KILOMETERS

0 5 MILES

220

230

250

270

240

300

290

280

310

320

280

290

300

310

320

330

340

Honolulu

Kaua‘i

Ni‘ihau

O‘ahu

Maui

Moloka‘i

Kaho‘olawe

Hawaiian Islands

Lāna‘i

Hawai‘i PACIFIC OCEAN

280

270

270 250

260

260

250

B. O‘AHU

KO‘OLAU VOLCANO

WAI‘ANAE VOLCANO

220

230

250

270

240

A. HAWAI‘I

Bouguer gravity isogal, contour interval 10 mGal

Crack

Fault

Kīlauea Caldera

Outer bounding faults (Neal and Lockwood, 2003)

Bouguer gravity isogal, contour interval 10 mGal

Dike

Caldera (Stearns, 1939)

Caldera outer limit (Walker, 1987)

Bouguer gravity isogal, contour interval 10 mGal

Fault

Dike

Caldera (Macdonald and others, 1960)

MAUNA LOA VOLCANO

KĪLAUEA VOLCANO

300

290

280

310

320

Hawai‘i

O‘ahu

C. KAUA‘I 280

290

300

310

320

330

340

260

270

270

290

250
114 Characteristics of Hawaiian Volcanoes

Where they are youthful, Hawaiian calderas form 

topographic features 2–5 km across. They mark rudely circular 

zones of subsidence, lava infilling, and intrusion. Subtle 

arcuate faults may mark the outer limit of caldera-related 

deformation, which extends much farther outboard than the 

prominent cliff of the most mobile part of the caldera. Such is 

the case at Mauna Loa and Kīlauea volcanoes today (fig. 10A) 

(Neal and Lockwood, 2003). Some caldera floors subside at 

least 800–900 m cumulatively through time, to account for 

the depth of exposure seen in caldera-filling rocks at deeply 

eroded volcanoes like Ko‘olau (O‘ahu) (Walker, 1987). 

Whether the subaerially emplaced lava of a caldera sinks 

substantially farther may never be known, because Hawaiian 

volcanoes subside into the submarine realm before erosion can 

expose their deep structural levels.

Older calderas, which may lack any present topographic 

expression, are mapped on the basis of several features that 

define the caldera boundary (Macdonald, 1965; Walker, 

1987). Caldera-filling lava-lake flows commonly are thick 

and vesicle poor; others, from fissures across a caldera’s 

floor, may be thin and vesicular. Structurally the lava flows 

are subhorizontal or dip inward if deformed by caldera 

subsidence. Caldera-filling lava flows might even dip outward 

slightly if they were built up around a central-vent location, 

as around Halema‘uma‘u (Kīlauea), but the dips of those 

flows are less than the 2°–10° dips of lava flows on a shield’s 

flanks. In some instances, talus breccia exposed within the 

rock sequence marks the trace of cliffs that bounded the 

caldera, as on Kaua‘i (Macdonald and others, 1960) and at 

Wai‘anae volcano (O‘ahu; Stearns and Vaksvik, 1935; Sinton, 

1987). At Ko‘olau volcano (O‘ahu), hydrothermal alteration 

is extensive in the caldera-filling lava sequence (Stearns and 

Vaksvik, 1935), perhaps a deeper-seated equivalent to the 

modern fumaroles on the floor of Kīlauea Caldera (Casadevall 

and Hazlett, 1983). The Ko‘olau caldera is also defined, in 

part, by a paucity of dikes relative to the adjacent rift zone 

(Walker, 1987), because the caldera’s episodic collapse buried 

or destroyed that part of the stratigraphic sequence, only to be 

filled anew with intact lava flows.

An enlarging caldera may engulf a previous caldera or 

shift its central region of subsidence, resulting ultimately in 

a set of bounding faults that enclose an area larger than that 

affected in any single caldera-forming event. That process is 

shown by the overlapping concentric or elliptical rims that 

form Mauna Loa’s caldera today. Thus, what is mapped as a 

large caldera in older volcanoes may simply be the convenient 

boundary drawn to encompass exposures of several successive 

caldera formations. Calderas may expand by the capture 

of adjacent pit craters (one of the mechanisms suggested 

by Macdonald, 1965), but the structural relation would be 

difficult to show, and pit craters may simply be casualties of 

proximity, not features fundamental to caldera growth.

The caldera on Kaua‘i is unusual for its substantial 

breadth (see the map of Macdonald and others, 1960); at 

16 km by 20 km, it is two to three times larger than other 

Hawaiian calderas (fig. 10C). The Kaua‘i caldera is also 

unusual, though not unique, in that it does not coincide 

closely with peak gravity values (Krivoy and others, 1965; 

Flinders and others, 2010). The mapped caldera may be a 

subsidence-and-fill feature unrelated or only marginally 

related to what would have been the shield’s original summit 

caldera (Holcomb and others, 1997; Flinders and others, 2010). 

The key structural features that could confirm this idea are 

unidentified because little mapping has taken place on Kaua‘i, 

especially of the older parts of the volcano, since the 1940s 

and 1950s (Macdonald and others, 1960), a reminder that basic 

geologic map data remain incomplete for parts of the Hawaiian 

Islands, even as we forge ahead in the 21st century.

The prevailing view of caldera formation, at Hawai‘i and 

elsewhere, is by subsidence consequent upon withdrawal of 

magma. The withdrawal removes support from a volcano’s 

summit region, causing it to subside. In an alternative 

mechanism, subsidence is due to the load of intrusions 

and cumulate rocks—the trail of present and past magma 

chambers stacked one above the other as the volcano grows 

(see, for example, Walker, 1987, 1988). This mass becomes 

unstable, relative to less dense surroundings, and settles into 

the crust.

Where does the magma go? The magma-withdrawal 

hypothesis is budgetary, in that the volume of caldera 

subsidence should roughly equal the volume of magma 

withdrawn. Objections have arisen about its application to 

Hawaiian calderas, because contemporaneous lava flows 

rarely match the volume of subsidence. With this comparison 

in mind, Macdonald (1965) compiled data for 13 events 

presumed to be related to caldera collapse at Kīlauea Volcano 

in the period 1823–1955. The volumes of subsidence ranged 

from 0.1 to 0.6 km3, but subaerially extruded lava volumes 

(corrected to dense magmatic equivalent) were only a fraction 

of that, from as little as 2 to 80 percent5.

The missing volume for some of these caldera-forming 

events might be accounted for by lava flows erupted in the 

submarine realm. For example, lava flows with thin sediment 

cover and therefore relative youthfulness (identified by high 

sonar backscatter) surround the distal end of Puna Ridge, 

the submarine part of Kīlauea’s East Rift Zone (Holcomb 

and others, 1988). The youngest of these flows is probably 

<200 yr old, covers about 600 km2, and has a volume >6 km3

if assigned a thickness of 10 m, which is the low value for 

their flow-margin thickness (Holcomb and others, 1988; 

Holcomb and Robinson, 2004, their unit Qx). This flow, 

however, is the lone known example of a shield-stage alkalic 

lava flow on Kīlauea (Clague and others, 1995; Johnson and 

others, 2002; Coombs and others, 2006); thus, it probably 

did not pass through the subcaldera reservoir and rift zone 

of Kīlauea and is unrelated to caldera subsidence prior to 

1790. Adjacent extensive sea-floor lava flows, tholeiitic in 

5

It was noted during technical review that these calculations may have little 

bearing on the issue of caldera subsidence if the events were simply draining 

of fluid lava (broad “lakes”) into the rift zones, and not the downdropping of 

solidified lava flows (D.A. Swanson, written commun., 2012).
Growth and Degradation of Hawaiian Volcanoes 115

composition (Clague and others, 1995), are significantly 

older than the alkalic flow, as are all sampled or observed 

flows along the crest of the Puna Ridge (Clague and others 

1995, Johnson and others, 2002). The large lava flows at the 

submarine end of the rift could be related to caldera collapse 

events, but their ages and volumes are poorly constrained, 

so correlation with known subaerial volcanic events is 

speculative. Despite these uncertainties, their eruption may 

contribute substantially to balancing the budget of summit 

caldera collapses and rift-zone extrusion.

Another objection against the withdrawal hypothesis 

is that collapse need not accompany voluminous 

subaerial eruptions. No caldera collapse is known to have 

accompanied the ‘Ailā‘au eruptions, for example, which shed 

about 5 km3 onto the east flank of Kīlauea during a 50-year 

period about 1445 C.E. Instead, the summit area is thought 

to have remained intact so that lava could continue to spill 

eastward (Clague and others, 1999). ‘Ailā‘au may have had 

little obvious relation to caldera formation, mainly because 

it was essentially a summit eruption (at rates near historical 

rates) and therefore did not depressurize the summit magma 

chamber in the way that rift-zone eruptions do. Regardless, 

initiation of major collapse of the modern caldera occurred 

soon after the end of the ‘Ailā‘au eruption (Swanson 

and others, 2012), which leaves open the suggestion that 

sustained eruptions may prepare the ground for subsequent 

subsidence-related deformation, even if not producing a 

caldera immediately, as in the classic withdrawal model.

Perhaps the answer that overcomes all of these 

objections to the withdrawal model of caldera formation is 

the combination of large lava flows, submarine extrusion, 

the capacity of rift zones to store intruded magma, and the 

element of time. Caldera collapses may follow years-long 

periods when roof-rock strength and honeycombing of voids 

allows the summit region to remain intact while magma 

is withdrawn. For example, geodetic data suggest that 

Kīlauea’s East Rift Zone expanded along an 8.5-km-high 

dike to accommodate about 0.18 km3/yr of magma during 

a 6-year, essentially noneruptive time period between large 

south-flank earthquakes of 1975 and 1982 (Cayol and 

others, 2000). As modeled, the corresponding rift-zone 

fault was 47 km long, dilation was 40 cm/yr, and the area 

of slip on the décollement ultimately covered 132km2

(fig. 3 of Cayol and others, 2000). No caldera collapse 

followed this event, probably because input to the volcano 

was being shunted away from the summit (as opposed to 

draining the summit). Regardless, the volume of magma 

involved reaches into the realm needed for some collapse 

events. The episodicity of rift-zone intrusion and its possible 

correlation with specific caldera subsidence events are still 

to be determined. The geodetic constraint represented by 

an active rift zone episodically expanding, over a period 

of a few years, on the order of meters across kilometerslong dimensions of height and length, will likely prove 

fundamental for understanding shield magmatism, rift-zone 

growth, and caldera subsidence.

Interactions with Water During the Shield 

Stage

Water plays a key role in phreatic and phreatomagmatic 

eruptions and in determining the slope of Hawaiian volcanoes 

above and below sea level. The widely observed phreatic 

eruption on Kīlauea in 1924, the earlier 300-year-long period 

of phreatomagmatic eruptions at Kīlauea that produced the 

Keanakāko‘i Tephra Member of the Puna Basalt (culminating in 

1790 C.E.; Swanson and others, 2012), and the presence of extensive 

older ash deposits on parts of Kīlauea and Mauna Loa—examples 

are the Kalanaokuaiki Tephra (Fiske and others, 2009) and the 

Pāhala Ash (Easton, 1987)—make it clear that water influences 

the style of eruptions, and the volcanic hazards, at Kīlauea and, 

presumably, at older volcanoes in the chain.

Clague and Dixon (2000) used a model for formation and 

solidification of magma chambers (Clague, 1987a), involving 

hydrothermal cooling of those magma chambers, magmatic 

degassing, and timing of explosive eruptions, to propose that 

Hawaiian volcanoes undergo a number of important changes as 

they grow from the seafloor into tall subaerial mountains. When 

the volcano is submarine, such as Lō‘ihi Seamount, a shallow 

magma reservoir will persist as soon as magma flux is large 

enough that heat input exceeds heat extraction by hydrothermal 

circulation. The hydrothermal fluids are derived from 

seawater and interact with magma in the reservoir, where they 

contaminate resident magma with Cl, Na, and other components 

that are abundant in seawater (see, for example, Kent and 

others, 1999) but of low concentration in magma.

As the summit of the volcano grows into shallow water (as 

deep as 1 km) and then through sea level, significant degassing 

of magmatic volatiles can take place. Loss of these gases, 

especially of water, increases the density of the shallowest 

magma in the reservoir, leading to overturn and magma mixing 

in the reservoir (Dixon and others, 1991; Wallace and Anderson, 

1998). This is also when phreatic and phreatomagmatic 

eruptions start, although lava fountains can occur even deeper, 

driven by magmatic volatile losses (as on Lō‘ihi; Clague and 

others, 2003; Schipper and others, 2010b).

The next transition occurs when the top of the magma 

chamber reaches sea level as the hydrothermal fluids freshen 

and magma contamination by seawater diminishes. The final 

transition takes place when the top of the magma chamber 

rises above the orographic rain belt on the islands, where 

hydrothermal fluids are no longer replenished, and rapid 

cooling of the magma chamber by convection of hydrothermal 

fluids ceases, as also does fumarolic discharge. Mauna Loa 

has grown to this stage, but Kīlauea remains with the top of 

its magma chamber just at, or a little above, sea level, so that 

explosive eruptions remain common (Easton, 1987; Mastin, 

1997; Swanson and others, 2012). All these processes impart 

important characteristics on the erupted lavas and tephra, 

including contamination, degassing, mixing, and fractionation, 

but all also appear to be restricted in time to the late preshield, 

shield, and perhaps the beginning of the postshield stages.
116 Characteristics of Hawaiian Volcanoes

One notable characteristic of Hawaiian volcanoes is 

steeper submarine than subaerial slope. This topographic 

distinction results from the subaqueous chilling of subaerial 

lava flows by seawater, which tends to increase a lava flow’s 

effective viscosity and diminish its effective density because 

rock-water has a lower density contrast than rock-air (Mark 

and Moore, 1987). Fragmentation of lava at the shoreline and 

greater angle of repose in water also contribute to the increase 

in slope offshore. On older volcanoes, this slope change has 

been used to identify shorelines now submerged far below 

sea level around Hawai‘i (Clague and Moore, 1991; Moore 

and Clague, 1992), Maui Nui (Price and Elliot-Fisk, 2004; 

Faichney and others, 2009, 2010), the main Hawaiian Islands 

(Carson and Clague, 1995), and along the entire HawaiianEmperor Seamount chain (Clague, 1996). These shorelines 

define the areas and sizes of former islands and are a key to 

understanding subsidence of the volcanoes (the topic of a 

later section).

Source Components of Shield-Stage 

Lavas

Chemical differences between shield lavas from adjacent 

Hawaiian volcanoes have been recognized for many years 

(see, for example, Tatsumoto, 1978; Frey and Rhodes, 1993). 

These intervolcano chemical differences, especially in isotopic 

systematics, correlate strongly with the geographic locations of 

the volcanoes along two curved subparallel trends, called the Loa 

and the Kea trends (fig. 11), that are defined through at least the 

eastern part of the main islands and are named after the largest 

volcanoes included in each trend—Mauna Kea and Mauna Loa.

Chemical differences in major, trace, and volatile 

elements and in radiogenic and rare-gas isotopic compositions 

of Hawaiian lavas have been used widely to define different 

mantle components involved in partial melting to produce 

the lavas (for example, Hauri, 1996; Dixon and Clague, 

2001; DePaolo and others, 2001; Gaffney and others, 2004; 

Ren and others, 2004, 2006; Abouchami and others, 2005; 

Xu and others, 2005, 2007b; Dixon and others, 2008; Huang 

and others, 2009). A major objective has been to understand 

the chemical structure of the Hawaiian plume. These 

various mantle components are commonly named after the 

volcano whose lavas display the most extreme end-member 

compositions—hence the KEA and KOO components 

are named for Mauna Kea and Ko‘olau volcanoes. Other 

components include a depleted HA component for Hawaiian 

asthenosphere, and FOZO, an entrained lower mantle 

component (fig. 12; Dixon and Clague, 2001). All except the 

depleted upper mantle component are thought to represent 

different parts of ancient lithosphere subducted a billion or 

so years ago and stored in the mantle. Other authors have 

suggested different end-member components, but all models 

include at least four components, one of which is depleted 

upper mantle (Helz and others, this volume, chap. 6). Much 

of the current discussion among scientists is centered on 

defining the full spectrum of chemical characteristics of the 

Ni‘ihau O‘ahu

Island of Hawai‘i

Kaua‘i

Moloka‘i

Ko‘olau

Haleakalā

West Maui

Kīlauea

Lō‘ihi

Mauna Kea

Mauna Loa

Lāna‘i

Wai‘anae

Ka‘ena

Ridge

Penguin

Bank 

Kaho‘olawe

Hualālai

Māhukona

160°

22°

21°

20°

19°

159° 158° 157° 156° 155°

KEA TREND

LOA TREND

0

1,000

2,000

3,000

4,000

5,000

6,000

Water depth,

in meters

0 100 200 KILOMETERS

0 50 100 MILES

Figure 11. Map showing 

the Kea and Loa trends of 

volcanoes in the Hawaiian 

Islands. Diamonds are 

volcano summits, locations 

from table 1. Base from 

Eakins and others (2003). 
Growth and Degradation of Hawaiian Volcanoes 117

80

40

0

160

120

600 500 400 300 200 100 0 -100 -200

Depth, in kilometersDepth, in kilometers

High

Low

Shield-building

Postshield

Rejuvenated

Melt

production Temperature

Magmas

Distance from plume center, in kilometers

Vertical exaggeration ×2.5

A Kīlauea

O‘ahu Haleakalā Kīlauea Mauna Loa Lō‘ihi 0

80

40

160

120

200

Lithosphere

Mantle plume

Pacific Plate motion

HA HA

B

Focus zone (FOZO)—thin zone of 

heated and entrained lower mantle

Metasomatized mantle—region of interaction 

between upwelling plume and passively 

upwelling upper-mantle asthenosphere

Hawaiian asthenospheric component (HA)—

passively upwelling upper-mantle asthenosphere

Melt segregation-and-reaction paths

Figure 12. Cross sections showing the inferred structure of the mantle plume below the Hawaiian hot spot. Plate motion is 

from right to left. A, Isotherms and melt-production isopleths (from Ribe and Christensen, 1999). B, Compositional components in 

the plume and surrounding mantle, as illustrated by Rhodes and Hart (1995; based loosely on thermal modeling by Watson and 

McKenzie, 1991), with additional zonation from Dixon and Clague (2001). Lithospheric thickness differs between panels A and B

because of modeling decisions but is between 70 and 90 km. Both panels show vertical exaggeration ×2.5.

end-member components and how these different components 

are arranged in the mantle source region beneath Hawai‘i. Two 

recent papers specifically address the origins of the chemical 

differences between Loa-trend and Kea-trend volcanoes 

(Huang and others 2011; Weis and others 2011) and the role of 

bilaterally zoned plumes. Matzen and others (2011) recently 

suggested that primary Hawaiian tholeiitic melts contain 

19–21 percent MgO, significantly higher than most previous 

estimates. This estimate, if confirmed, implies that Hawaiian 

tholeiitic magmas are generated deeper and at higher potential 

temperatures than are primary melts with lower MgO.

Transition from Shield to Postshield 

Stage

Hawaiian volcanoes lose vigor as they drift away from 

the hot spot. Shallow magma reservoirs that lie only 1–7km 

deep during main stage activity cannot be sustained as magma 

supply decreases (Clague, 1987a), but reservoirs persist 

between 19-km (Bohrson and Clague, 1988) and 28-km depth 

(Frey and others, 1990), near the base of the downflexed ocean 

crust. This solidification of at least shallow magma reservoirs 

takes place near the end of the shield stage or within the 

postshield stage as magma supply dwindles.

As might be expected, the shift from shield to postshield 

stage is transitional. It is expressed geologically by the increasing 

number of alkalic basalt lava flows found interbedded in the upper 

part of the main stage sequence (table2). For example, at Kohala 

volcano, strata that were interpreted as shield stage show great 

compositional variation, spanning well into the alkalic range and 

even crossing from basalt into hawaiite (fig. 3, Pololū Volcanics; 

Wolfe and Morris, 1996a,b). It is doubtful that any of the Hawaiian 

volcanoes shift abruptly from shield- to postshield-stage volcanism 

without first sputtering through a period of transition.

Other lithologic and petrographic changes are associated 

with lava flows during the transition. At West Maui, for 

example, the upper 50 m of shield stage strata shows an 

increase in the proportion of ‘a‘ā and an increase in the size 

and abundance of olivine and clinopyroxene phenocrysts 

(Diller, 1982; Sinton, 2005). At both West Maui and Waiʻanae 
118 Characteristics of Hawaiian Volcanoes

Table 2. Hawaiian volcanic formations and their approximate place in the scheme of shield-, late shield-, postshield-, and rejuvenatedstage volcanism.

[Mauna Loa, Kīlauea, and Lō‘ihi are not listed because they are in the shield stage only]

Volcano Tholeiitic shield

exposed

Late shield/postshield

alkali basalt1 Postshield hawaiite-trachyte Rejuvenated stage

Ni‘ihau Pānī‘au Basalt Postshield plug (Ka‘eo), 

dike, and offshore cones

None Ki‘eki‘e Basalt

Kaua‘i Waimea Canyon 

Basalt, 

Nāpali and Olokele 

Members

Upper part of Makaweli 

and Olokele Members

Scattered flows in upper part 

of Makaweli and Olokele 

Members

Kōloa Volcanics

Southwest of Ka‘ena 

Ridge(?) (submarine)2

Tholeiite recovered by 

dredge

None None Sparse rejuvenatedstage(?) alkalic 

lava

Wai‘anae (O‘ahu) Wai‘anae Volcanics, 

Lualualei Member

Wai‘anae Volcanics, 

Kamaileunu Member

Wai‘anae Volcanics, Pālehua 

and Kolekole Members

None

Ko‘olau (O‘ahu) Ko‘olau Basalt None None Honolulu Volcanics

Penguin Bank(?) 

(submarine)3

Only tholeiitic rocks 

have been sampled

None None None

West Moloka‘i West Moloka‘i 

Volcanics

Spotty Wai‘eli and other late lava 

flows

None

East Moloka‘i Not exposed(?) East Moloka‘i 

Volcanics, 

lower member

East Moloka‘i Volcanics, 

upper member. As much as 

25 percent basanite on basis 

of published analyses

Kalaupapa 

Volcanics

Lāna‘i Lāna‘i Basalt None None None

Kaho‘olawe Kanapou Volcanics, 

main shield and 

caldera-filling 

strata

Kanapou Volcanics, 

late shield

Some Kanapou Volcanics are 

hawaiite and mugearite

None

West Maui Wailuku Basalt Spotty Honolua Volcanics Lahaina Volcanics

Haleakalā Not exposed(?) Honomanū Basalt Kula and Hāna Volcanics None

Māhukona

(submarine)

None known

Kohala Pololū Volcanics, 

lower part(?)

Pololū Volcanics Hāwī Volcanics None

Mauna Kea Drill core Hāmākua Volcanics Laupāhoehoe Volcanics None

Hualālai Submarine dredged 

samples; drill core

Hualālai Volcanics 

(Waawaa Trachyte 

Member at base)

Hualālai Volcanics, Waawaa 

Trachyte Member

None

1

It is not our goal to impose a new stage name (late shield), but instead to identify for the reader those stratigraphic units that pose the greatest 

problem of assignment within a rigid scheme of stages. Each Hawaiian volcano brings its own fingerprint to the story, which makes generalizations 

difficult.

2

A topographic prominence southwest of Ka‘ena Ridge was demarcated as a separate volcano by Eakins and others (2003). More commonly, it and 

Ka‘ena Ridge are considered a westerly rift zone of Wai‘anae volcano (O‘ahu). For example, Ka‘ena Ridge was considered part of Wai‘anae for 

volume calculations of individual volcanoes in the Hawaiian chain (Robinson and Eakins, 2006).

3

Penguin Bank was proposed as a separate volcano by Carson and Clague (1995). On that basis it was used in calculations that describe the 

topographic history of Maui Nui, the multivolcano complex that encompasses the islands of Moloka‘i, Lāna‘i, Kaho‘olawe, and Maui (Price 

and Elliot-Fisk, 2004). More common is the assignment of Penguin Bank as a rift zone of West Moloka‘i volcano (for example, Robinson and 

Eakins, 2006).
Growth and Degradation of Hawaiian Volcanoes 119

the trend is toward increasingly silicic lava upward in the 

Laupāhoehoe Volcanics (Wolfe and others, 1997). At Waiʻanae 

volcano (fig. 13B), the trend is opposite, from the hawaiitemugearite of the Pālehua Member of the Wai‘anae Volcanics to 

alkali basalt of the overlying Kolekole Member (as defined by 

Presley and others, 1997). The trend at Haleakalā (fig. 13C) is 

less sharp, but the lava is increasingly alkaline upsection from 

the lower to the upper part of the Kula Volcanics. The overlying 

Hāna Volcanics is geochemically similar to the upper part of the 

Kula (Sherrod and others, 2003).

Distribution

Lava of the postshield stage forms thick sequences on 

Wai‘anae (O‘ahu), East Moloka‘i, West Maui, Haleakalā, 

Kohala, Mauna Kea, and Hualālai volcanoes. The postshield 

stage is represented by only a few flows and small volumes on 

Ni‘ihau, Kaua‘i, West Moloka‘i, and Kaho‘olawe volcanoes. 

Postshield stage lava is absent from Ko‘olau and Lāna‘i 

volcanoes.

Onset

Postshield-stage volcanism follows immediately after 

the shield stage. At some volcanoes, like Kaua‘i, a discrete 

postshield sequence is ill-defined, but alkalic basalt, hawaiite, 

and mugearite—rocks characteristic of postshield strata at 

other volcanoes—are present in the upper part of the shieldstage Waimea Canyon Basalt.

No temporal gap between late shield and postshield 

stages is evident at volcanoes that have a distinct, readily 

mapped postshield stratigraphic sequence. If a hiatus exists, 

it is too brief to date by K-Ar or 40Ar/39Ar geochronology. 

This brevity is indicated in the dating archive by an overlap 

of radiometric ages (and their analytical errors), even where 

the field evidence indicates no stratigraphic interfingering. 

In the field the stratigraphic break is sharp but concordant, 

suggesting little erosion during the transition.

On West Maui, for example, ages of the shield-stage 

Wailuku Basalt range from about 2 to 1.35 Ma, and those of 

the postshield Honolua Volcanics range from 1.35 to 1.2Ma 

(McDougall, 1964; Naughton and others, 1980; Sherrod and 

others, 2007a). The statistical overlap of ages at about 1.3 Ma 

suggests that very little time elapsed during the switchover 

from stage to stage. No field-based evidence of interfingering 

is known, so on West Maui, the terminal period of shieldstage volcanism was relatively brief, not protracted across 

many hundreds of thousands of years.

A similarly brief or nonexistent hiatus marks the 

transition from shield to postshield stage on West Moloka‘i. 

There, two dated postshield cones have ages of 1.80 and 

1.73Ma (Clague, 1987b), whereas fractionated tholeiite 

near the top of the shield sequence has an age of 1.84 Ma 

(McDougall, 1964). The analytical error for these three ages 

permits them to be roughly synchronous.

(O‘ahu) volcanoes, ash beds and soil horizons are increasingly 

common upsection, the latter indicating a general decrease in 

eruption frequency (Macdonald, 1968; Sinton, 1987). The ash 

beds indicate a landscape increasingly speckled by scoria cones 

and small domes that produced ash and lapilli, in contrast to 

the spatter-rich vents characteristic of shield-stage eruptions. 

Some interpret this change as a general increase in explosivity 

of eruptions during this transitional time. Instead of being more 

explosive in the transitional and postshield phases, however, 

a Hawaiian volcano may produce less lava relative to vent 

deposits, so we see more of the Strombolian products. During 

this transition, vents also become scattered more widely, so the 

likelihood of widely distributed tephra might increase.

Postshield Stage

Lava flows and associated tephra in the stratigraphic 

formations attributed to postshield-stage volcanism are more 

alkalic and commonly more differentiated than the tholeiitic 

lava of the shield stage. Most common are rocks with 

compositions in the range hawaiite to benmoreite; trachyte is 

rare (fig. 13). Alkalic basalt may also be present.

The petrogenetic cause for these changes was stated 

succinctly by Sinton (2005):

A consistent characteristic of lava compositions from 

most postshield formations is evidence for post-melting 

evolution at moderately high pressures (3–7kb). Thus, 

the mapped shield to postshield transitions primarily 

reflect the disappearance of shallow magma chambers 

(and associated calderas) in Hawaiian volcanoes. . . . 

Petrological signatures of high-pressure evolution are 

high-temperature crystallization of clinopyroxene and 

delayed crystallization of plagioclase, commonly to 

<3percent MgO.

This description builds on earlier ideas about postshield 

volcanism (Clague, 1987a), such as loss of shallow magma 

chambers, the end of caldera formation, and the predominance 

of magma-storage zones at deep crustal to uppermost mantle 

depths in excess of 15 km. Deep differentiation of magmas by 

crystal fractionation leads to the more fractionated lava flows 

of the postshield stage.

Only two volcanoes among the Hawaiian Islands lack the 

alkalic rocks that characterize postshield volcanism. Ko‘olau 

(O‘ahu) entered the rejuvenated stage following a hiatus of 

1m.y. after ending the shield stage. Lāna‘i lacks any volcanism 

younger than the shield stage, which ended there about 1.3 Ma. 

Excluded from this count are the three youngest volcanoes 

(Mauna Loa, Kīlauea, and Lō‘ihi), which lack postshield 

volcanism simply because they have yet to progress fully 

through the shield stage. Figure 14 shows the stages and timing 

for volcanoes at each of the main Hawaiian Islands.

At some volcanoes the postshield-stage stratigraphic 

sequence can be subdivided into members that have fairly 

discrete geochemical groupings. At Mauna Kea (fig. 13A), 
120 Characteristics of Hawaiian Volcanoes

40 45 50 55 60 65 0

2

4

6

8

10

SiO2, in weight percent

Mugearite

Benmoreite

Tephriphonolite

Trachyte

Phonotephrite

Basanite

Foidite

Alkalic

Tholeiitic

Hawaiite

0

2

4

6

8

10

Na O + K 2 O, in weight percent 2

Mugearite

Benmoreite

Tephriphonolite

Trachyte

Phonotephrite

Basanite

Foidite

Picrobasalt

basalt

Basaltic

andesite andesite

Basalt

Basaltic

andesite Andesite

Basalt

Basaltic

andesite Andesite

Alkalic

Tholeiitic

0

2

4

6

8

10

12

Benmoreite

Tephriphonolite

Trachyte

Phonotephrite

Basanite

Foidite

Picrobasalt

Alkalic

Tholeiitic

A

Postshield 1

Postshield 2

B

C

Postshield 1

Postshield 2

Postshield 2

Postshield 1

Laupāhoehoe Volcanics

Younger volcanic rocks member

Older volcanic rocks member

Hāmākua Volcanics

Wai‘anae postshield (Kolekole Member)

Wai‘anae postshield (Pālehua Member)

Wai‘anae late shield (Kamaile‘unu Member)

Wai‘anae Volcanics, O‘ahu

Hāna Volcanics

Kula Volcanics (upper)

Kula Volcanics (lower)

Honomanū Basalt

Figure 13. Postshield 

volcanic sequences 

from A, Mauna Kea; 

B, Wai‘anae; and C, 

Haleakalā; compared 

using the alkali-silica 

diagram (Na2O+K2O 

versus SiO2). Rock 

classification grid 

from Le Maitre (2002); 

shown dashed is 

boundary separating 

tholeiitic from alkalic 

basalt (Macdonald and 

Katsura, 1964). Data 

from multiple sources 

compiled in a Hawaii 

statewide wholerock geochemistry 

GIS database 

(Sherrod and others, 

2007b). Stratigraphic 

nomenclature for 

Wai‘anae volcano is 

that of Sinton (1987) 

and Presley and others 

(1997).
Growth and Degradation of Hawaiian Volcanoes 121

0.35

2.32

4.33

6.0

4.0

5.8 

3.65

0.15

Postshield

or late shield

3.3

1.8

0.8

0.1

3.55

3.93

3.06

2.98

3.06

2.90 Kolekole

Pālehua

1.8

1.7

2.1

1.49

1.35

0.35

1.75

1.32

1.35

1.2

0.3

0.6

2.0

1.1

0.96

0.12

Kula

Hāna

1.1

0.26

0.12

V o l c a n o b i r t h ?

0.3

0.06

0.115

0

1

2

Age, in millions of years (Ma) 3

4

5

6

Distance northwest of Kīlauea, in kilometers

600 500 400 300 200 100 0 -50 Wai‘anae Ni‘ihau Kaua‘i Ko‘olau West Moloka‘i East Moloka‘i

Lāna‘i

West MauiKaho‘olawe

Haleakalā

M

Kohala(Hilo Ridge)

āhukona

Hual

Mauna Kea ālai

Mauna Loa

KīlaueaLō‘ihi

Rejuvenated stage

Postshield stage

Shield stage

Late part of shield stage

0.3

?

1.2

0.9

0.3

175 km, Haleakalā

Figure 14. Plot of age versus distance from Kīlauea, showing 

the duration of the different developmental stages for each 

volcano in the main Hawaiian Islands. From Clague and 

Dalrymple (1987) but updated according to a Hawaii statewide 

geochronologic GIS database (Sherrod and others, 2007b) and 

additional ages from Kīlauea (Lipman and others, 2006), Kaua‘i 

(Garcia and others, 2010), Kohala (Lipman and Calvert, 2011), 

and Ko‘olau (Yamasaki and others, 2011). Red line, suggesting 

age of volcano birth, has slope of 10 cm/yr (average plate 

motion). Some stratigraphic formations or members are listed 

for Haleakalā and Wai‘anae volcanoes, where the postshield 

stage comprises multiple units.

Duration of Postshield-Stage Volcanism

Once underway, postshield-stage volcanism typically 

persists for 100,000–500,000 years but may continue for as 

long as 1 m.y., as at Haleakalā (table 3; fig. 14). Short-duration 

examples include East Moloka‘i, where lava flows and domes as 

thick as 520 m accrued in the period between 1.49 and 1.35 Ma; 

and West Maui, where postshield strata 120 m thick accumulated 

in the period 1.35–1.2 Ma (Sherrod and others, 2007a). At 

Mauna Kea, postshield-stage volcanism has been ongoing for 

265,000 years (Wolfe and Morris, 1996a). Hualālai, too, is in 

the postshield stage, persisting already for about 115,000years 

(Cousens and others, 2003). On Wai‘anae volcano, the 

postshield Pālehua Member has a narrow age range from about 

3.06 to 2.98 Ma (fig. 15; Presley and others, 1997), suggesting 

eruption in as little as 100,000 years, and the overlying 

Kolekole Member extends the postshield duration by only 

another 140,000 years. Where transitions mark the beginning 

and end of postshield activity (as is the case at Kaua‘i), the 

duration of the postshield stage has large uncertainty.
122 Characteristics of Hawaiian Volcanoes

60 70

1

2

3

4

Kamaile‘unu Member Pālehua Member Kolekole Member

Lualualei

Member

Shield stage Postshield stage

A. Wai‘anae volcano

Shield stage Postshield stage

B. West Maui volcano

Wailuku Basalt Honolua Volcanics

Sequential order, by decreasing age within stratigraphic units

Age, in millions of years

0

0

1

2

3

10 20 30 40

2σ

50

Figure 15. Plots of sample ages showing the 

lack of a time gap between shield- and postshieldstage strata at (A) Wai‘anae (O‘ahu) and (B) West 

Maui volcanoes (Maui), which are the two Hawaiian 

volcanoes with deepest exposures into the shieldstage strata and most extensive dating of shield 

and postshield lava. Ages arranged from oldest to 

youngest within each stratigraphic unit, but precise 

stratigraphic position of a sample relative to others 

in the same formation is rarely known. Gray band 

shows likely range of stratigraphically valid ages 

across each formation, as a guide to recognizing 

ages too old or too young. Specific data sources are 

drawn from a Hawaii statewide geochronologic GIS 

database (Sherrod and others, 2007b). Error bars are 

two standard deviations (2σ). For Wai‘anae volcano, 

stratigraphic members are part of the Wai‘anae 

Volcanics as defined by Sinton (1987) and Presley 

and others (1997).

Ni‘ihau

At Ni‘ihau, the postshield-stage duration is ill-defined, 

owing to large analytical uncertainty in the radiometric 

ages. An alkalic dike and the eroded remnants of two cones 

represent the postshield stage onshore, and a few additional 

postshield cones are located offshore. Four onshore postshield 

samples yielded K-Ar ages ranging from 5.15±0.225 to 

4.67±0.16 Ma (G.B. Dalrymple, data in Sherrod and others, 

2007b), and 40Ar/39Ar isochron ages from two offshore 

postshield cones are 4.93±0.44 and 4.74±0.54 Ma (table 4), 

within the range of the on-land samples. These ages allow that 

postshield activity on Ni‘ihau could be as brief as 100,000–

200,000 years or as lengthy as about 500,000 years.

Kaua‘i

On Kaua‘i, postshield lava flows of hawaiite and 

mugearite with ages in the range 3.84–3.81 Ma occur at the 

top of the Olokele and Makaweli Members of the Waimea 

Canyon Basalt (Clague and Dalrymple, 1988). Clasts of 

other probable postshield-stage alkalic basalts occur in 

conglomerate recovered in water wells from the southeastern 

part of the island (Reiners and others, 1999). The oldest 

dated lava with postshield chemical affinity, an irregular dike 

of alkalic basalt from Kālepa Ridge, has an 40Ar/39Ar plateau 

age of 4.39±0.19 Ma and an isochron age of 4.39±0.07 Ma 

(table 4); the spread of ages from oldest to youngest suggests 

the postshield stage lasted roughly 600,000 years (more than 
Growth and Degradation of Hawaiian Volcanoes 123

Table 3. Thickness and duration of postshield-stage volcanism at Hawaiian volcanoes with well-defined hawaiite-trachyte suites above shield-stage 

stratigraphic sequences.

[Coverage indicates percent of volcano’s subaerial surface mantled by postshield deposits. Volume not corrected to dense-rock equivalence. Duration of postshield activity in millions of years (m.y.); the age of the activity is in millions of years ago (Ma)]

Volcano Postshield occurrence and coverage Thickness Volume,

in km3 Duration

Ni‘ihau One small intrusion n.a. --

Kaua‘i Scattered lava flows, not mapped separately -- -- --

Wai‘anae 20% covered ≤180 m 0.22 m.y. (3.06–2.84 Ma)

Ko‘olau None 0 0 0

West Moloka‘i 3 small lava flows, 16% covered n.d. <0.1 m.y. (?)

East Moloka‘i 50% covered ≤520 m n.d. 0.14 m.y. (1.49–1.35 Ma)

Lāna‘i None 0 0 0

Kaho‘olawe Thin extensive cover, 77%

West Maui 18% covered 120 m 0.15 m.y. (1.35–1.2 Ma)

Haleakalā Continuous thick cover, 95% (Kula and Hāna Volcanics) 1 km 3001 0.95 m.y., ongoing

Kohala 44% covered (Hāwī Volcanics) 0.14 m.y.

Mauna Kea 950 m 200–5002 0.25 m.y., ongoing

Hualālai 3003 0.1 m.y., ongoing

1

Haleakalā postshield volume from Sherrod and others (2003).

2

Mauna Kea postshield volume is crude estimate corresponding to area of Laupāhoehoe Volcanics and upper part of Hāmākua Volcanics multiplied by thickness 

range 50–200 m.

3

Hualālai postshield volume is crude estimate corresponding to volume of volcano above sea level.

Table 4. New 40Ar/39Ar incremental heating ages for samples from Ni‘ihau and Kaua‘i.

[Ni‘ihau ages analyzed by W.C. McIntosh, New Mexico Geochronology Research Laboratory, New Mexico Institute of Mining and Technology, Socorro, New 

Mexico. Kaua‘i ages analyzed by John Huard, Noble Gas Mass Spectrometry Lab, College of Earth, Ocean, and Atmospheric Sciences, Oregon State University, 

Corvallis, Oregon. Irrad. No., irradiation run sequence. The value n shows number of step increments used to determine the plateau age. MSWD, mean square of 

weighted deviates. WGS84, datum of World Geodetic System 1984. Sample altitude relative to mean sea level; negative altitudes are depths below sea level]

Sample Volcano

stage

Lab

no.

Irrad.

no.

Age 

method n

%

39Ar MSWD K/Ca±2σ Age±2σ, 

in Ma

Longitude, 

WGS84

Latitude, 

WGS84

Altitude, 

in meters

Island of Ni‘ihau

T322-R17 Shield 55320-01 NM-185F Plateau 7 90.8 2.6 0.2±0.1 5.42±0.11 –160.2233 22.09420 –1092

T321-R7 Postshield 55336-01 NM-185J Isochron 9 22.2 -- 0.4±0.6 4.93±0.44 –160.3529 21.9463 –821

T321-R6 Postshield 55332-01 NM-185H Isochron 8 22.3 -- 0.3±0.3 4.74±0.54 –160.3521 21.9458 –852

T317-R6 Rejuvenated 55316-02 NM-185F Plateau 2 27.4 4.6 0.1±0 1.37±0.32 -- -- --

T317-R9 Rejuvenated 55318-01 NM-185F Plateau 9 94.1 2.5 0.1±0.2 0.50±0.04 –160.2488 22.1310 –1692

T317-R8 Rejuvenated 55338-01 NM-185J Plateau 8 91.7 3.4 0.1±0.1 0.50±0.10 –160.2489 22.1305 –1700

T322-R6 Rejuvenated 55334-01 NM-185J Plateau 8 89.7 1.6 0.1±0.1 0.39±0.09 –160.2317 22.0913 –1407

Island of Kaua‘i

86KA3 Postshield 05c3516 OSUSF05 Plateau 3 65.1 0.2 0.58±0.19 4.39±0.19 –159.3576 22.0022 183

86KA2 Rejuvenated 05c3492 OSUSF05 Plateau 3 61.2 4.1 0.03±0.08 2.18±0.271 –159.6008 22.0125 451

75K1 Rejuvenated 05c3524 OSUSF05 Plateau 6 97.9 0.8 0.02±0.03 0.68±0.042 –159.5244 21.9153 244

76K1 Rejuvenated 05c3486 OSUSF05 Plateau 5 98.2 1.0 0.02±0.03 0.52±0.033 –159.3583 22.0376 37

For Kaua‘i samples, superscripted ages indicate samples also dated by Clague and Dalrymple (1988), with the following resulting K-Ar ages and ±1σ error:

1

1.914±0.023 Ma

2

0.648±0.034 Ma

3

0.554±0.023 Ma
124 Characteristics of Hawaiian Volcanoes

shown on fig. 14). Also assigned to the uppermost part of 

the Waimea Canyon Basalt (corresponding to postshieldstage volcanism) are three basanite samples with K-Ar

ages 

in the range 3.92±0.03 to 3.85±0.06 Ma (Garcia and 

others, 2010). These samples share chemical and isotopic 

characteristics with other younger Kōloa lavas. Their ages 

and isotope chemistry, in conjunction with those of the 

dated hawaiite and mugearite, suggest that the transition 

from postshield to rejuvenated stages may be gradational 

(interbedded stratigraphically) and that an eruptive hiatus 

need not necessarily occur in order to define the onset of the 

rejuvenated stage.

Wai‘anae

The postshield sequence on Wai‘anae volcano (O‘ahu) 

comprises the Pālehua and Kolekole Members of the Wai‘anae 

Volcanics (mapping of J.M. Sinton, plate 3 in Sherrod and 

others, 2007b). The Pālehua, 3.06–2.98 Ma in age, contains 

hawaiite and mugearite, whereas the overlying Kolekole, 

2.98–2.90 Ma, is tholeiitic (Presley and others, 1997). The 

Kolekole Member, which contains the youngest of the 

lava flows on Waiʻanae volcano, was once interpreted as a 

rejuvenated-stage formation, but it has been reclassified as late 

postshield stage, partly because it follows so closely on the 

deposition of the Pālehua Member (Presley and others, 1997). 

Chemically, Kolekole lava flows are distinct from rejuvenatedstage lavas by virtue of their higher SiO2 (compare figs. 13B

and 17). They and the underlying Pālehua Member form an 

unusual postshield sequence, however, because the younger 

part (Kolekole) is generally lower in total alkali content than 

the older part (Pālehua; fig. 13B).

Haleakalā

Much longer postshield stages characterize a few 

volcanoes. At Haleakalā volcano, the Kula Volcanics erupted 

from 0.96 Ma until 0.12 Ma, and activity was continuous 

with the overlying Hāna Volcanics (Sherrod and others, 

2003). These two formations form a cap 1 km thick across the 

summit of Haleakalā.

In a matter of clarification, the Hāna Volcanics 

stratigraphic unit was long considered a rejuvenated-stage 

formation on the basis of a presumed depositional hiatus that 

separated it from the Kula Volcanics (Stearns and Macdonald, 

1942). Erosion of large valleys on East Maui, which preceded 

emplacement of the Hāna, was thought to require substantial 

time. Detailed dating at Haleakalā, however, shows that as 

little as 0.03 m.y. may have been required to deeply scallop 

the landscape. The Hāna Volcanics unit is geochemically 

similar to the upper part of Haleakalā’s postshield formation, 

the Kula Volcanics. The lack of any intervening hiatus favors 

an interpretation that the Hāna is merely the waning phase of 

postshield volcanism (Sherrod and others, 2003).

Kaho‘olawe

The Kaho‘olawe volcanic complex exposes the youngest 

part of shield-building strata, including what were described as 

postcaldera strata (Stearns, 1940). More recently, geochemical 

analyses suggest that these postcaldera lava flows, which range 

from tholeiitic basalt to hawaiite, correspond to a transition 

into the postshield volcanic stage (Fodor and others, 1992; 

Leeman and others, 1994). The extent of likely postshieldstage strata, as shown on the geologic map of Hawai‘i (plate 6 

of Sherrod and others, 2007b), is derived from an unpublished 

compilation by Harold Stearns (courtesy of M.O. Garcia).

Kaho‘olawe’s youngest volcanic products are found 

on the east side of the island. These lava flows and tephra, 

tholeiitic in composition, overlie older strata with pronounced 

discordance, probably owing to a preceding episode of slope 

collapse. In the absence of dating, this discordance was 

taken as evidence that substantial time intervened before 

eruption of the mantling volcanic rocks (Stearns, 1940), 

which were thought to be part of rejuvenated-stage volcanism 

(Langenheim and Clague, 1987). However, new ages of about 

0.98 Ma (Sano and others, 2006) show that the mantling lava 

flows are coeval with Kaho‘olawe’s postshield strata, and their 

chemistry makes them similar to those strata. The structural 

discordance between the two is a rare instance where the age 

of a slope-failure event has been dated directly.

Distribution of Postshield-Stage Cones and 

Eruptive Fissures

Postshield cinder and spatter cones are concentrated 

along preexisting rift zones on West Moloka‘i, Haleakalā, 

and Hualālai, whereas they are more dispersed on Kohala, 

East Moloka‘i, and Wai‘anae. On Mauna Kea, the cones are 

scattered over a large region of the summit and upper slopes.

Xenoliths in Postshield Lava and Tephra

Xenoliths are abundant in postshield lavas and cinder 

on Mauna Kea (Fodor and Vandermeyden, 1988; Fodor and 

Galar, 1997; Fodor, 2001) and Hualālai (Jackson, 1968; 

Jackson and others, 1981; Bohrson and Clague, 1988; Clague 

and Bohrson, 1991; Chen and others, 1992; Shamberger 

and Hammer, 2006). Xenoliths in postshield lavas are also 

present at Kohala, East Moloka‘i, West Maui, and Wai‘anae 

but are much less common and smaller (Jackson and others, 

1982). All xenoliths in postshield eruptive products are midto-deep crustal cumulates, including rare ocean crust gabbro 

(Clague, 1987a). The ultramafic xenoliths are dominated by 

dunite, clinopyroxenite, and wehrlite rather than the mantle 

lithologies lherzolite and pyroxenite (with or without garnet) 

or harzburgite that predominate in rejuvenated-stage tephra 

and lavas and in rare preshield-stage lavas from Lō‘ihi 

Seamount (Clague, 1988).
Growth and Degradation of Hawaiian Volcanoes 125

Isotopes Indicate Changing Source Composition 

Through Time

Fractionation and changes in source composition 

accompany the development of postshield-stage volcanic 

sequences, although such distinctions are difficult to recognize 

at those Hawaiian volcanoes where the accumulated strata are 

thin or unevenly distributed. Where the postshield sequence 

is thick, however, or where the episode was sufficiently longlived, chemical variation through time can commonly be 

documented.

At Haleakalā, the Hawaiian volcano with the greatest 

exposed thickness of postshield strata, strontium isotopic 

ratios diminish upsection, from 0.70355 (in the lower part 

of Kula Volcanics) to as low as 0.70308 (upper part of Kula 

and overlying Hāna Volcanics; fig.16A; data of West and 

Leeman, 1987, 1994; West, 1988; Chen and others, 1991). An 

even greater span is seen if the stratigraphically transitional 

lava of the Honomanū Basalt is considered (fig.16A). The 

strontium isotopes indicate that the contribution of various 

source components in the underlying mantle changed during 

the lengthy history of postshield-stage volcanism at Haleakalā 

(Chen and others, 1991). For East Moloka‘i, a similar pattern 

may apply, but there the postshield stratigraphic sequence is too 

thin to define an extensive change (fig. 16B).

Rejuvenated-Stage Volcanism

The rejuvenated stage encompasses those volcanic deposits 

that form the latest stage in a Hawaiian volcano’s history. As 

classically defined, the term “rejuvenated stage” was conceived 

to account for volcanism that resumed following a period of 

quiescence. Indeed, rejuvenated-stage volcanism was first 

recognized because field evidence indicated an intervening 

period of inactivity—for example, weathering contrasts among 

older and younger lava flows, the preservation of primary 

geomorphic features in young lava flows or vents and their 

absence in older ones, or evidence that substantial erosion 

preceded the emplacement of lava flows in canyons or across 

alluvial fans. By these criteria, 5of the 15 emergent volcanoes 

on the 8 major Hawaiian Islands have rejuvenated-stage 

volcanism: Ni‘ihau, Kaua‘i, Ko‘olau (O‘ahu), East Moloka‘i, 

and West Maui volcanoes. Rejuvenated-stage eruptions include 

the youngest eruptions on the Islands of Ni‘ihau, Kaua‘i, O‘ahu, 

and Moloka‘i. Rejuvenated-stage lavas of West Maui volcano 

are older than postshield-stage eruptions of Haleakalā on the 

same island, and no volcano on the Island of Hawai‘i has 

reached the rejuvenated stage.

Rejuvenated-stage lava flows and vent deposits diminish 

greatly in areal extent southeastward along the island chain. 

About 35 percent of both Ni‘ihau and Kaua‘i are covered 

by rejuvenated-stage products. In contrast, at Ko‘olau 

volcano (O‘ahu) the coverage is only 6 percent. Coverageof 

East Moloka‘i is only about 2 percent, where this stage is 

represented only by the single Kalaupapa shield (Clague 

and others, 1982) and perhaps two islets off the east end of 

Moloka‘i (Stearns and Macdonald, 1947). On West Maui 

the coverage is a mere 0.6 percent, the consequence of four 

scoria cones and two lava-flow units (Tagami and others, 

2003). This areal comparison, from simple GIS calculations 

(Sherrod and others, 2007b), lacks the significance of volume 

calculations. But volume decreases likely are similar in 

magnitude or even more substantial southeastward, because 

the extensive lava sequences on the northwestern islands 

include more basin-filling, thick deposits. Erosion has 

stripped some products, and substantial areas of rejuvenatedstage lava flows are flooded in the offshore reaches of the 

islands, which further limits the comparison. The time span of 

emplacement is longer on Ni‘ihau and Kaua‘i (2.5–3.5m.y) 

than on the other three islands (only the past 1 m.y.). Even 

so, normalization for age will not compensate for these 

differences in areal extent.

Rejuvenated-stage rocks share some similarities with 

postshield-stage rocks. Both are alkalic. Eruptive products 

are scoria, cinder, and ash cones, and lava flows, chiefly ‘a‘ā. 

Vent loci are scattered without regard to the alignment of rift 

zones that were active in shield-stage time.

Does a Time Gap or Chemical Change Define 

Rejuvenated-Stage Volcanism?

Where a time gap occurs between postshield and 

rejuvenated stages, that hiatus in volcanism ranges from 

0.6 m.y (West Maui; Tagami and others, 2003) to ~2 m.y. 

(Ni‘ihau; Clague and Dalrymple, 1987). The hiatus can be 

estimated confidently where rejuvenated-stage eruptive vents 

are few, because the entire rejuvenated-stage sequence can 

be dated. Helpful, too, is adequate dating of the underlying 

late shield or postshield lava flows emplaced before the 

quiescence. It is only in the last decade that these goals have 

been reached for many of the Hawaiian volcanoes.

But is a time gap a requirement to define rejuvenatedstage volcanism? The past two decades brought an 

onslaught of geochemical analyses and isotopic dating to 

Hawaiian volcano investigations. Discoveries have brought 

surprises, not the least of which is the exceedingly brief 

time gap that separates some volcanic stratigraphic units 

that once were thought divided by substantial interludes. 

Lengthy time gaps simplify the assignment of strata to 

growth stages, whereas transitions during brief interludes 

confound the task. The definition of Hawaiian rejuvenatedstage volcanism is currently a dilemma. No single chemical 

criterion adequately distinguishes between postshield- and 

rejuvenated-stage volcanism. And if a hiatus is required, 

then what is the definitive duration? Herein we first address 

the general characteristics of rejuvenated-stage volcanism 

and then return to the question of definitions.
126 Characteristics of Hawaiian Volcanoes

0.78±0.01 m.y.

0.97±0.04 m.y.

1.10±0.05 m.y.

0.93

0.97

-30

-20

-10

0

10

20

30

40

50

60

1.05

0.6

Hāna Volcanics

Transitional

or late shield

Kula Volcanics (upper part)

Kula Volcanics (lower part)

0.7 

0.78

Stratigraphic position

Age, in millions of years (Ma)

-200

-100

0

Stratigraphic depth, in meters below

highest shield-stage lava

5

4

3

2

1

0

Age, in millions

of years ago (Ma)

Postshield

Postshield

Rejuvenated

Rejuvenated

Rejuvenated

Postshield

Late shield

Late shield

Clague and

Dalrymple (1988)

Garcia and

others (2010)

0.7028 0.7030 0.7032 0.7034 0.7036 0.7038 0.7040 0.7042 0.7044 0.7046

87Sr / 86Sr

Eruptive stage (no stratigraphic

information for samples within

their formations)

D. Ko‘olau (O‘ahu)

C. Kaua‘i

B. East Moloka‘i

A. Haleakalā

Shield
Growth and Degradation of Hawaiian Volcanoes 127

Figure 16. Graphs showing strontium isotopic ratio (87Sr/86Sr), 

relative to stratigraphic position for strata of late shield, postshield, 

and rejuvenated stages at selected Hawaiian volcanoes. A, Haleakalā 

volcano. Data from West and Leeman (1987) and Chen and others 

(1991). “Stratigraphic position” corresponds to numbering system of 

West and Leeman (1994) for lava flows of Kula Volcanics. Samples 

plotted in stratigraphic positions 0 to –6 are from underlying strata now 

considered part of the Kula (Macdonald and others, 1983), and the top 

of Honomanū Basalt is placed lower. Stratigraphic positions for Hāna 

Volcanics assigned by using location data in West (1988) and mapping 

and dating by Sherrod and McGeehin (1999). Ages, in millions of years 

ago (Ma), from isotopically analyzed samples from Chen and others 

(1991). Generalized, nonlinear time scale relies on magnetic polarity 

data and the few ages available from samples with Sr isotope data 

(chiefly from Chen and others, 1991). B, East Moloka‘i volcano. Isotopic 

data from Xu and others (2005) for samples collected along Kalaupapa 

Trail traverse along the cliffy north flank of Moloka‘i, where lava flows 

are subhorizontal in attitude. Sample altitudes, reported by Sherrod 

and others (2007b), were recalculated to stratigraphic depth by using 

datum of highest shield-stage lava. C, Kaua‘i volcano. Isotopic analyses 

and ages from Clague and Dalrymple (1988) and Garcia and others 

(2010). D, Ko‘olau volcano. Samples analyzed for isotopic ratio lack 

stratigraphic control aside from stratigraphic formation assignment. Not 

shown are four rejuvenated-stage analyses in range 0.7025–0.7027, to 

conserve space in the presentation. Discounted are three rejuvenatedstage analyses (unshaded squares) that range from 0.7038 to 0.7045 

(Lessing and Catanzaro, 1964), as high as any of the shield-stage 

results. Vague descriptions permit those three samples to be stream 

boulders or accidental lithic inclusions in tuff cones.

Youngest Rejuvenated-Stage Lava Flows Along 

the Island Chain

Sustained, sporadic rejuvenated-stage eruptions have 

continued into relatively recent time on five of the major 

islands, even though the age of the shield stage (and therefore 

of most of the lava volume) decreases from northwest 

(Ni‘ihau) to southeast (Hawai‘i). Indeed, when postshield-stage 

volcanism is added to the mix, every island except Lāna‘i and 

Kaho‘olawe has had an eruption in the past 0.4 m.y.

On Ni‘ihau, rejuvenated-stage volcanic rocks (the Ki‘eki‘e 

Basalt) are as young as 350 ka (fig. 14; G.B. Dalrymple, in 

Sherrod and others, 2007b). Four samples from offshore 

range in age from 1.39 to 0.39 Ma (table 4). On Kaua‘i, 

the rejuvenated-stage Kōloa Volcanics has ages as young 

as 150 ka, with 10 samples younger than 500 ka (Garcia 

and others, 2010). On O‘ahu’s Ko‘olau volcano, subaerial 

deposits of the Honolulu Volcanics generally yield youngest 

ages of about 100 ka (see results and discussion in Ozawa 

and others, 2005); younger ages have been reported but never 

corroborated. Offshore, submarine samples from the Koko Rift, 

also part of the Honolulu Volcanics, yielded youngest ages of 

about 140 ka (Clague and others, 2006). On East Moloka‘i, 

the Kalaupapa Volcanics have an age of about 350 ka (Clague 

and others, 1982). On West Maui, lava flows of the Lahaina 

Volcanics are clustered into two age groups: about 600 ka and 

300 ka (Tagami and others, 2003).

In summary, the youngest rejuvenated-stage lavas 

on Ni‘ihau, Kaua‘i, Ko‘olau (O‘ahu), East Moloka‘i, and 

West Maui are in the age range 350–100 ka. The duration 

of the rejuvenated stage, coupled with youngest ages of 

volcanic rocks, means that future eruptions on these islands 

are conceivable, despite their present distances from the 

Hawaiian hot spot. The infrequency of rejuvenated-stage 

eruptions, however, reduces the risk they pose to society to 

very low levels.

A Volumetrically Insignificant Part of Hawaiian 

Volcanoes

Rejuvenated-stage volcanism has long been known to 

contribute much less than 1 percent to the cumulative volume 

of a Hawaiian shield volcano (Macdonald and others, 1983; 

Clague, 1987a; Clague and Dalrymple, 1987). This estimate 

recently has been quantified for Kaua‘i, which has the most 

extensive rejuvenated-stage products among all the islands. 

About 60 km3 of rejuvenated-stage lava and tephra has 

been emplaced on the island (Garcia and others, 2010). A 

conversion to dense-rock-equivalent magma would reduce that 

volume by about 25 percent, because most of the lava is ‘a‘ā. 

Given Kaua‘i’s total shield volume of 57,600 km3 (Robinson 

and Eakins, 2006), the rejuvenated-stage lava is only about 

0.1percent of the total (Garcia and others, 2010).

Ni‘ihau is second in the abundance of rejuvenatedstage products (Holcomb and Robinson, 2004; Dixon and 

others, 2008). At Ni‘ihau, a substantial but poorly assessed 

volume of such products lies offshore (Clague and others, 

2000)—but that is still insufficient to amass even 1 percent of 

the volcano’s volume. Rejuvenated-stage lava flows are also 

widespread offshore around Ka‘ula, an islet 37 km southwest 

of Ni‘ihau (Holcomb and Robinson, 2004; Garcia and others, 

2008). Offshore lava flows and cones are rare around O‘ahu 

(Clague and others, 2006), Moloka‘i (Clague and Moore, 

2002), and Kaua‘i (Holcomb and Robinson, 2004). They are 

unknown offshore near any other Hawaiian islands.

However, volcanism resembling rejuvenated stage has 

occurred on the sea floor at somewhat greater distances from 

the islands themselves. An extensive young lava field (330 km 

along its northeast-trending axis) was emplaced on the deep 

seafloor 100–400 km north of O‘ahu (the North Arch field; 

Clague and others, 1990). These flows cover some 24,000 km2

and have an estimated volume of 1,000–1,250 km3 (Clague 

and others, 2002), about 20 times the volume of rejuvenatedstage lava on Kaua‘i. They were erupted from more than 100 

vents, whose form ranges from low shields to steep cones. 

Geochemically the flows are similar to Hawaiian rejuvenatedstage lavas (Yang and others, 2003; Hanyu and others, 

2007) by virtue of their high abundance of incompatible 

elements, depleted rare gases, higher 143Nd/144Nd, and lower 

87Sr/86Sr relative to shield-stage lava. Other examples are 

seen southwest of O‘ahu and at several places along the 

Hawaiian chain as far west as the Midway Islands (Holcomb 
128 Characteristics of Hawaiian Volcanoes

and Robinson, 2004). These submarine flows call into question 

the concept of the rejuvenated stage, because they occur at 

locations that lack prior shield-stage activity, yet they have 

geochemical characteristics indicating similar magmatic sources 

and origins as the island-mantling rejuvenated-stage volcanic 

rocks (Yang and others, 2003).

Other submarine alkalic lavas, similar in many, but not all, 

aspects to rejuvenated-stage lavas, occur on the Hawaiian Arch 

south of Hawai‘i (Lipman and others, 1989; Hanyu and others, 

2005). They are richer in H2O and have higher, more primitive 

He isotopic ratios than rejuvenated-stage lavas (Dixon and 

Clague, 2001; Hanyu and others, 2005).

Geochemical Characteristics

The products of some rejuvenated-stage volcanism have 

silica contents that range down to values as low or lower 

than rocks of the shield and postshield stages, in the range 

35–45percent SiO2. Rejuvenated and shield-stage compositions 

are nearly distinct, but the fields for rejuvenated and postshield 

Honolulu Volcanics (on Ko‘olau shield)

Kōloa Volcanics (on Kaua‘i shield)

Ki‘eki‘e Basalt (on Ni‘ihau shield)

Rejuvenated-stage rocks

35 40 45 50 55 60 0

2

4

6

8

10

12

14

Na O + K 2 O, in weight percent 2

SiO2, in weight percent

Mugearite

Benmoreite

Tephriphonolite

Phonolite

Trachyte

Phonotephrite

Basanite

Hawaiite

Foidite

Basalt Basaltic andesite Andesite Tholeiitic

Alkalic

Kaua‘i

postshield

Ko‘olau shield

Ni‘ihau postshield

Ni‘ihau shield

Kaua‘i shield

Figure 17. Alkali-silica diagram (Na2O+K2O versus SiO2) for rejuvenated-stage volcanic rocks on Ni‘ihau, 

Kaua‘i, and O‘ahu. Also shown are the fields of the shield and postshield rocks on those volcanoes. Rock 

classification grid from Le Maitre (2002); shown dashed is boundary separating tholeiitic from alkalic basalt 

(Macdonald and Katsura, 1964). Listed in appendix are the specific data sources for these data, on the basis of 

a published Hawaii-statewide whole-rock geochemistry GIS database (Sherrod and others, 2007b).

analyses overlap substantially (fig. 17). As another example, at 

high MgO content (10–15 percent), contents of incompatible 

elements, such as Rb (15–40 ppm), are similar to the enriched 

values seen in postshield rocks.

Strontium isotopic ratios for rejuvenated-stage lava are 

lower than those of late-shield rocks, but they overlap with 

those of postshield flows and tephra. On East Moloka‘i, where 

the volcanic stages are clearly demarcated, rejuvenated-stage 

lava (Kalaupapa Volcanics) has 87Sr/86Sr less than 0.7032, 

and underlying postshield and shield-stage lava has ratios 

greater than 0.7033 (fig. 16B). At Ko‘olau volcano (O‘ahu), 

where the postshield stage is lacking, 87Sr/86Sr is in the range 

0.7025–0.7036 for rejuvenated-stage lava (Honolulu Volcanics) 

and 0.7036–0.7045 for shield-stage lava (fig. 16D). On Kaua‘i, 

Sr isotopic ratios of rejuvenated-stage lava (0.7030–0.7033) are 

only slightly lower than those of several of the postshield-stage 

rocks (~0.7033; fig. 16C). Indeed, on Kaua‘i, the notable shift 

toward lower 87Sr/86Sr values occurs in the postshield stage 

before the largest time gap in activity, but with both high and low 

Sr isotopic ratios in samples erupted at nearly the same time.
Growth and Degradation of Hawaiian Volcanoes 129

Duration of Quiescence

The hiatus in volcanism, where such exists, between 

postshield and rejuvenated stages ranges from 0.6 m.y. (West 

Maui; Tagami and others, 2003) to ~2 m.y. (Ni‘ihau; Clague 

and Dalrymple, 1987). Estimating the hiatus can be done with 

some confidence where rejuvenated-stage eruptive vents are 

few, because the entire rejuvenated-stage sequence can be 

dated, and the youngest underlying late shield or postshield 

lava flows are also well dated. These conditions have been 

reached in the last decade for many of the islands’ volcanoes.

For Kaua‘i, Rejuvenated Stage Is Gradational from 

Postshield Stage

For the Island of Kaua‘i, the age and chemistry of 

earliest rejuvenated-stage volcanism suggests no hiatus in 

eruptive activity in the progression from postshield stage. 

Radiometric dating in the 1980s suggested an odd feature of 

Kauaʻi’s rejuvenated stage compared to that of other Hawaiian 

volcanoes—an exceedingly brief quiescence preceding the 

rejuvenated stage but a lengthy pause from 3.65 to 2.59 Ma 

within rejuvenated time (Clague and Dalrymple, 1988). 

This “start-pause-resume” history may be the result of an 

incompletely sampled sequence of volcanic rocks.

Recently, 40Ar/39Ar ages as old as about 3.4 Ma were 

obtained from drill cuttings from the Hanamā‘ulu well, which 

penetrates a deep basin on the east side of Kaua‘i (Izuka 

and Sherrod, 2011). The ages are 3.11±0.56, 3.22±0.26, and 

3.42±0.24 Ma at downhole depths of 160, 210, and 270 m, 

respectively. All lavas in the Hanamā‘ulu well are geochemically 

similar to rejuvenated-stage lavas (Reiners and others, 1999). 

Thus, even with their large analytical errors, these new ages fill 

much of the gap in the preexisting dating archive of rejuvenatedstage lava in the range 3.6–3.0 Ma. The hiatus disappears 

completely if three basanite lava flows—described by Garcia 

and others (2010) as postshield with ages of 3.92, 3.85, and 

3.58Ma—are assigned instead to the Kōloa Volcanics, as 

their low Sr isotopic values suggest (<0.70322 for all three). 

Regardless of how these controversial samples are finally 

assigned across the postshield-rejuvenated-stage boundary, any 

hiatus before rejuvenated-stage activity on Kaua‘i was extremely 

brief, as was any hiatus between shield and postshield activity. 

Plausibly, volcanic activity on Kaua‘i continued uninterrupted 

from shield to postshield to rejuvenated stage, and the change in 

isotopic signature (especially Sr) at about 3.9–3.8Ma marks a 

significant change in the magma source region.

Hypotheses of Rejuvenated-Stage Magma 

Generation

Rejuvenated-stage volcanism probably originates from 

lithospheric sources (Lassiter and others, 2000) that undergo 

decompression melting—for example, as the lithosphere 

rebounds from the zone of depression beneath the largest 

young volcanoes (Jackson and Wright, 1970; Bianco and 

others, 2005) or when hot mantle, dragged initially downward 

in response to plate motion, rises naturally by its lower 

density until pressure and temperature are suitable for melt 

production (Ribe and Christensen, 1999). A low degree of 

melting of recently metasomatized depleted mantle has been 

the general model for the formation of rejuvenated-stage 

magmas for nearly 30 years (Clague and Frey, 1982). As 

modeled by Garcia and others (2010), mantle in the upper 

part of the plume undergoes a low degree of partial melting, 

less than 3 percent. More recent studies have focused on the 

mantle components included in the melting or on the nature 

and amount of carbonatitic and silicate melt metasomatism 

(Dixon and others, 2008).

An alternative explanation for rejuvenated-stage magmas 

invokes lithospheric melting by conductive heating (Gurriet, 

1987), but this is inconsistent with Pb isotopic data from 

Kaua‘i and O‘ahu (Garcia and others, 2010). It appears that 

the mantle plume contributes little or no source material 

(Yang and others, 2003; Hanyu and others, 2005). In any 

model, enhanced crustal fracturing, such as might result from 

the lithospheric rebound, may enable magma to percolate 

upward more easily (Clague and others, 1990).

Degradation and Eventual Submergence 

of Hawaiian Volcanoes

A spectrum of processes work to degrade the volcanoes 

once they have formed. Primary among these are subaerial 

erosion, landslides, and subsidence.

Erosion

Hawaiian volcanoes, especially those that grow high 

above sea level, experience high rainfall, particularly on their 

northeastern sides that face the trade winds. Mean annual 

rainfall amounts locally exceed 9 m (Giambelluca and others, 

1986; Haleakala Climate Network, 2011), leading to rapid 

erosion that carves deep canyons in <1 million years, as 

seen at Kohala on the Island of Hawai‘i. Other features once 

thought to be entirely erosional, such as the pali (cliff) on the 

north side of East Moloka‘i, are now interpreted as landslide 

headwalls (on Moloka‘i, of the Wailau slide) or backstepping 

of those headwalls (Clague and Moore, 2002).

Landslides

Landslides occur at all scales in Hawai‘i, ranging from 

small slides that modify coastal cliffs and canyon walls to 

giant slides that displace large parts of islands (Denlinger 

and Morgan, this volume, chap. 4). The idea that Hawaiian 

volcanoes, particularly Ko‘olau (O‘ahu) and East Moloka‘i, 
130 Characteristics of Hawaiian Volcanoes

had been modified by giant landslides was first proposed by 

Moore (1964). The enormous scale of the landslides made 

the idea controversial at first, but the evidence became firmly 

established when the U.S. Geological Survey’s Marine 

Geology Program mapped the newly established 370-km 

(200nautical miles) Exclusive Economic Zone around 

Hawai‘i in the mid to late 1980s, using the GLORIA sidescan 

sonar system. Those surveys showed extensive debris fields 

strewn across the deep sea floor several hundred kilometers 

from the islands, as well as rotational slumps characterized by 

large segmented blocks (Moore and others, 1989). The debris 

fields were proposed to have formed during catastrophic 

debris avalanches spawned on the upper submarine flanks of 

the islands.

Seventeen discrete slides that formed in the past 5 m.y. 

were identified around the main Hawaiian Islands (Moore and 

others, 1989), and fully 70 are known along the Hawaiian 

Ridge between Midway Islands and the Island of Hawai‘i, 

formed during a 30-m.y. period of emplacement (Holcomb 

and Robinson, 2004). The corresponding rates of occurrence—

one per 300 k.y. and one per 400–450 k.y., respectively—

must be minima, because many slides likely occur along the 

leading (southeast) edge of the chain and are buried by 

growth of subsequent volcanoes. For example, the entire 

summit caldera and north rift of Ni‘ihau are missing, yet no 

corresponding seafloor deposit has been recognized (Moore 

and others, 1989). The growth of the Kaua‘i edifice has 

presumably covered the area northeast of Ni‘ihau where the 

landslide would have gone. Likewise, the summit of West 

Moloka‘i apparently slumped to the east, leaving only some 

small east-facing fault scarps, and was buried by growth of 

East Moloka‘i. The eastern half of the caldera on Kaho‘olawe 

is also missing, now presumably buried beneath Haleakalā 

and, perhaps, Māhukona and Kohala volcanoes.

Are large landslides more likely during particular growth 

stages at a volcano? The large slides seem most likely to 

occur during the most active volcano growth phases—from 

the late preshield to the early postshield stage. A zone of hot, 

deforming olivine cumulate inside the active volcanoes may 

provide some of the force needed to slide a volcano’s flank 

(Clague and Denlinger, 1994), providing a rationale for why 

the slides may occur mainly during the active growth stages 

of the volcano’s lifespan. The relation between the movement 

of Kīlauea’s south flank (and the southeast and west flanks 

of Mauna Loa) and the Hilina Pali slump identified in the 

offshore sonar and bathymetric maps (Morgan and others, 

2003) suggests that tectonics of the active volcanoes may 

trigger the slides. On the other hand, the abundance of the 

slides along the entire chain suggests that they could also 

occur long after the volcanoes become inactive. It remains 

uncertain how, why, and when the slides occur. Would we 

recognize the precursors, if there were any, should one happen 

now? The question of slide timing is not an academic one, 

as slides that could occur at any stage in the life of a volcano 

pose significantly higher risk to the populations of the older 

Hawaiian Islands than would catastrophic slides limited to 

only the active, and sparsely populated, flanks of Kīlauea and 

Mauna Loa volcanoes.

Landslides on the rainy windward sides of the islands are 

associated with deep erosional canyons (Clague and Moore, 

2002). High rainfall may lead to deep erosion, as well as high 

pore pressure on the windward (north and northeast) sides of the 

volcanoes, contributing to flank instability and failures while 

the rift and magma reservoir systems are actively spreading. 

Landslides, however, are equally common on the lee or dry 

sides of the islands. Studies of the submarine slides on the 

west flank of Mauna Loa have shown that the slides may move 

along surfaces that also serve as hot fluid pathways, forming 

greenschist facies metamorphic rocks along those surfaces 

(Morgan and Clague, 2003; Morgan and others, 2007).

Megatsunami

A significant hazard created by giant landslides at the 

Hawaiian Islands is the large displacement of seawater to 

generate catastrophic giant waves (megatsunami). The geologic 

evidence for megatsunami in the Hawaiian Islands was 

recognized first on Lāna‘i and Moloka‘i, where chaotic coral and 

lava-clast breccia is preserved as high as 155 m above sea level 

(J.G. Moore and Moore, 1984; G.W. Moore and Moore, 1988; 

Moore and others, 1994). This interpretation has been debated 

in numerous papers (for example, Grigg and Jones, 1997; Felton 

and others, 2000; Rubin and others, 2000; Keating and Helsey, 

2002), which argue that these high-stand deposits are the result 

of island uplift. No other data suggest that Lāna‘i is, or has been, 

uplifted, however, and evidence to the contrary has been gleaned 

from drowned reefs south of Lāna‘i (Webster and others, 2006, 

2007). McMurtry and others (2004) describe a similar deposit 

near the shoreline on Kohala volcano that is close in age to an 

offshore drowned reef now at nearly 400-m depth, suggesting 

that the megatsunami that produced this deposit washed up the 

slope of Kohala at least 400 m. Roughly one-third of the giant 

landslides from the Hawaiian Islands and the Hawaiian Ridge 

are debris avalanches with the capacity to generate huge tsunami 

that sweep hundreds of meters up the slopes of nearby islands. 

Such tsunami from future, though infrequent, landslides pose a 

large but unquantified risk to the State of Hawaii and possibly to 

coastal lands along the Pacific rim.

Subsidence

Subsidence of Hawaiian volcanoes occurs by two 

processes that overlap in time and space (Moore, 1987). The 

first process occurs as the increasing mass of a growing volcano 

depresses the underlying lithosphere. This rapid phase of 

subsidence, which lasts perhaps 1 m.y., submerges shorelines 

and reef complexes by more than 1 km. The rate of subsidence 

is high beneath young islands: for example, a rate of 2.6 m/k.y. 

was estimated off the northwest coast of the Island of Hawai‘i 

by dating drowned coral reefs (Ludwig and others, 1991). 

Smaller volcanoes may completely submerge during the short 
Growth and Degradation of Hawaiian Volcanoes 131

period of rapid subsidence, as did, for example, Māhukona 

volcano (Clague and Moore, 1991; Clague and Calvert, 2009) 

and many volcanoes in the northwestern Hawaiian Islands. 

Once a volcano has submerged by more than the glacioeustatic 

variations in sea level (~125 m), it is unlikely to reemerge. 

Drowned coral reefs are the primary evidence for this early 

period of rapid subsidence, and modeling has shown that 

the reefs drown during deglacial periods of rising sea level 

(Webster and others, 2009). Very high rates of sea-level rise 

during meltwater pulses of deglacial periods, coupled with 

rapid subsidence, may be required to rapidly submerge the reef 

(Webster and others, 2004). The drowned reefs surrounding 

each island formed and drowned during the period of rapid 

subsidence caused by lithospheric flexure during the active 

growth of each volcano, so reefs around Lāna‘i, for example, 

are older than reefs around Hawai‘i (Webster and others, 2010).

Concurrently, the lithosphere beneath the islands ages 

and thermally contracts. The rate of this subsidence diminishes 

exponentially as a function of lithospheric age. Along much of 

the Hawaiian Ridge and Emperor Seamounts, this subsidence 

rate is ~0.01 m/k.y. (Clague and others, 2010) and, thus, is 

insignificant for human culture.

The flexural loading of the lithosphere from the weight of 

the growing volcanic edifices causes a small amount of uplift, 

focused at a distance several hundred kilometers out from 

the center of the magma supply zone, to form the Hawaiian 

Arch (Deitz and Menard, 1953). Onshore evidence for a small 

amount of uplift comes from O‘ahu, where coral reef deposits 

of oxygen-isotope substage 5e (133–115 ka; Shackleton and 

others, 2003) are exposed around the entire island (Stearns, 

1974; Muhs and Szabo, 1994) and a ~334-ka coral reef now 

stands 21 m above sea level (McMurtry and others, 2010).

Large Hawaiian volcanoes can persist as islands through 

the rapid subsidence by building upward rapidly enough. But 

in the long run, the inexorable thermal contraction-induced 

subsidence, coupled with surface erosion, erases any volcanic 

remnant above sea level in about 15 m.y. Gardner Pinnacles, 

of that age, is the oldest surviving island in the chain with 

subaerially exposed volcanic rock. Beyond Gardner Pinnacles 

to the northwest are small sand islands and atolls, interspersed 

with smaller volcanoes whose summits submerged soon after 

the volcano formed. Many of the atolls have long and complex 

histories of carbonate deposition; at Kōkō Seamount in the 

southern Emperor Seamounts, deposition lasted from about 

50Ma to 16 Ma (Clague and others, 2010), when the last deepwater coralline algae finally submerged below their growth 

limit of –150 m. For the Hawaiian chain, a complex interplay 

of subsidence, northward movement into cooler waters, and 

rapid climate change is required to drown the reefs (Clague 

and others, 2010), because corals can grow faster than the 

slow rates of subsidence caused by thermal contraction of the 

lithosphere. The atolls and sand islands along the Hawaiian 

chain have undergone complicated growth, emergence, and 

subsidence histories related to Pleistocene and earlier sea level 

changes, but only Midway Islands atoll has been sampled 

well enough (Ladd and others, 1970) to perhaps decipher 

this history. In the Hawaiian chain, these changes conspire to 

drown the reefs about 33 m.y. after the underlying volcanoes 

formed, at which time the last of the atolls submerge to become 

guyots. Grigg (1982, 1997) called this time the Darwin Point. 

All the volcanoes west of Midway Islands and Kure Atoll are 

submerged, and most are guyots.

Future Work

In assembling the enormous amount of information about 

the geology of the Hawaiian Islands for this overview, we were 

struck by how much of our understanding of the evolution of 

Hawai‘i is built on three basic building blocks: detailed geologic 

mapping, radiometric dating, and geochemical analyses. Many 

of these framework studies were done decades ago and are 

now in dire need of updating. A fourth building block—that 

of submarine studies of the flanks and submarine rifts of the 

islands—has come into its own mainly during the past 25 years. 

If we expect to continue to develop new ideas and improve 

our knowledge of Hawaiian geology in the coming years, the 

Hawaiian Volcano Observatory should continue to foster broad 

interest and studies in the geology of all the Hawaiian Islands.

Acknowledgments

The authors thank Pete Lipman, John Sinton, Don 

Swanson, and Peter Stauffer for penetrating comments 

that greatly improved the presentation and content of the 

manuscript. We appreciate their time and effort in wading 

through such a long treatise. Jim Kauahikaua kindly provided 

the gravity data shown in figure 10. Our views on Hawaiian 

geology have developed over many years of interactions, 

collaborations, discussions, and arguments with many other 

scientists, most notably Michelle Coombs, Brian Cousens, 

Brent Dalrymple, Roger Denlinger, Jackie Dixon, Dan 

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Jim Kauahikaua, Pete Lipman, Jack Lockwood, Jim Moore, 

Julia Morgan, John Sinton, Don Swanson, Frank Trusdell, 

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JB093iB12p14856.
Growth and Degradation of Hawaiian Volcanoes 145

Appendix 

Described here are the myriad source publications used 

to prepare the chemical variation diagrams. All data were 

compiled into a geochemical database, and geographical 

coordinates were assigned successfully to about 70 percent of 

those. The geochemical database originated from a compilation 

by Kevin Johnson while at the Bishop Museum in the mid1990s6

. Our database was probably current until about 2004, 

with only sparse additions since then. For the Island of Hawai‘i 

we relied almost entirely upon an extant, major-element 

geochemical database of samples analyzed during a Big Island 

mapping project (Wolfe and Morris, 1996b). An electronic 

version of that database, including geographic coordinates, was 

published by Trusdell and others (2006).

Figure 3

Analyses for Kīlauea (456) and Mauna Loa (500) are 

from Wolfe and others (1996b). The total count of published 

analyses from Kīlauea and Mauna Loa is probably threefold 

greater, but the display of points shown is sufficient for the 

descriptive purposes of this chapter.

Late Shield-Stage Sequences (Fields)

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

Haleakalā, Honomanū Basalt, 41 analyses: West (1988); Chen 

and others (1991); Sherrod and others (2007b).

Mauna Kea, Hāmākua Volcanics, 196 analyses: Wolfe and 

Morris (1996b).

East Moloka‘i, lower member of East Moloka’i Volcanics, 

87analyses: Beeson (1976); Clague and Beeson (1980); 

Clague and Moore (2002); J.M. Sinton*; Xu and others (2005).

Kohala, Pololū Volcanics, 121 analyses: Wolfe and Morris (1996b).

Postshield-Stage Sequences

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

Wai‘anae volcano, Pālehua Member of Wai‘anae Volcanics, 

43analyses: Macdonald and Katsura (1964); Macdonald, 

(1968); Presley and others (1997); T.K. Presley*; J.M. Sinton 

and G.A. Macdonald.*

Kohala, Hāwī Volcanics, 114 analyses: Wolfe and Morris (1996b).

Mauna Kea, Laupāhoehoe Volcanics, 213 analyses: Wolfe and 

Morris (1996b).

Figure 13

Panel A, Mauna Kea Volcano (449 Analyses)

[Outcrops and drill core from depths shallower than 420 m; thus late shieldstage and postshield-stage strata] 

Rhodes (1996); Wolfe and Morris (1996b)

Panel B, Wai‘anae Volcano (196 Analyses)

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

J.M. Sinton and G.A. Macdonald*, Macdonald and Katsura 

(1964), T.K. Presley*, Presley and others (1997), Sinton 

(1987), Macdonald (1968), and Bauer and others (1973).

Panel C, Haleakalā Volcano (520 Analyses)

Macdonald and Powers (1946); Macdonald and Katsura 

(1964); Macdonald (1968); Macdonald and Powers (1968); 

Brill (1975); Horton (1977); Chen and others (1990); Chen 

and others (1991); West and Leeman (1994); Bergmanis 

(1998, with many appearing in Bergmanis and others, 2000); 

Sherrod and others (2003); D.R. Sherrod in Sherrod and 

others (2007b).

Figure 17

Shield-Stage Sequences (Fields)

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

Ni‘ihau, Pānī‘au Basalt, 20 analyses: D.A. Clague*.

Kaua‘i, Nāpali Member of Waimea Canyon Basalt, 8 analyses 

(exclusive of drill cuttings): Cross (1915); Macdonald and 

others (1960); Macdonald and Katsura (1964).

Ko‘olau volcano, Ko‘olau Basalt, 212 analyses: Wentworth 

and Winchell (1947), Yoder and Tilley (1962), Muir and Tilley 

(1963), Macdonald (1968), Jackson and Wright (1970), Frey 

and others (1994), Haskins and Garcia (2004), T.K. Presley*.

Postshield-Stage Sequences (Fields)

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

Ni‘ihau, Ka‘eo plug, 4 analyses: Washington and Keyes 

(1926); D.A. Clague*.

Kaua‘i, alkalic rocks in Olokele Member and Makaweli Member 

of Waimea Canyon Basalt, 15 analyses: Macdonald and Katsura 

(1964); Feigenson (1984); Clague and Dalrymple (1988).

6

Bishop Museum geochemical database current through about 1995. [http://

www.bishopmuseum.org/research/natsci/geology/geochem.html, accessed 

April 2012]
146 Characteristics of Hawaiian Volcanoes

Rejuvenated-Stage Data

[Asterisk indicates analyses provided originally as unpublished data to the 

Bishop Museum database]

Ni‘ihau volcano, Ki‘ei‘e Basalt, 30 analyses: Washington and Keyes 

(1926); Macdonald (1968); D.A. Clague*.

Kaua‘i volcano, Kōloa Volcanics (exclusive of Palikea Breccia 

Member), 183 analyses: Cross (1915); Washington and Keyes 

(1926); Macdonald and others (1960); Macdonald and Katsura 

(1964); Macdonald (1968); Kay and Gast (1973); Palmiter (1975); 

Feigenson (1984); Clague and Dalrymple (1988); Maaløe and 

others (1992); Reiners and Nelson (1998); Reiners and others 

(1999).

Ko‘olau volcano, Honolulu Volcanics, 142 analyses: Cross 

(1915), Winchell (1947), Tatsumoto (1966), Macdonald (1968), 

Macdonald and Powers (1968), Jackson and Wright (1970), 

Clague and Frey (1982), Wilkinson and Stolz (1983).
Next page: Coastline near Halapē, on the south flank of Kīlauea Volcano, before (top) and after (bottom) the M7.7 Kalapana earthquake, 

which resulted in several meters of seaward movement of the volcano’s flank, as well as subsidence (note the drowned trees in the bottom 

image). USGS photographs by D.A. Swanson in 1971 (top) and by J.G. Moore in 1975 (bottom), about a month after the earthquake.

On the island of O‘ahu, the rejuvenated Diamond Head (Lē‘ahi) Crater, in the foreground, and its deposits lie on top of the eroded Ko‘olau 

volcano, represented by the dissected ridge in the background. Waikiki is at the far left of the image. USGS photograph by P.G. Okubo, 

October 3, 2014.

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

1

U.S. Geological Survey.

2

Rice University.

Abstract

Hawaiian volcanoes build long rift zones and some 

of the largest volcanic edifices on Earth. For the active 

volcanoes on the Island of Hawai‘i, the growth of these rift 

zones is upward and seaward and occurs through a repetitive 

process of decades-long buildup of a magma-system 

head along the rift zones, followed by rapid large-scale 

displacement of the seaward flank in seconds to minutes. 

This large-scale flank movement, which may be rapid enough 

to generate a large earthquake and tsunami, always causes 

subsidence along the coast, opening of the rift zone, and 

collapse of the magma-system head. If magma continues 

to flow into the conduit and out into the rift system, then 

the cycle of growth and collapse begins again. This pattern 

characterizes currently active Kīlauea Volcano, where 

periods of upward and seaward growth along rift zones were 

punctuated by large (>10 m) and rapid flank displacements in 

1823, 1868, 1924, and 1975. At the much larger Mauna Loa 

volcano, rapid flank movements have occurred only twice in 

the past 200 years, in 1868 and 1951.

All seaward flank movement occurs along a detachment 

fault, or décollement, that forms within the mixture of pelagic 

clays and volcaniclastic deposits on the old seafloor and 

pushes up a bench of debris along the distal margin of the 

flank. The offshore uplift that builds this bench is generated by 

décollement slip that terminates upward into the overburden 

along thrust faults. Finite strain and finite strength models 

for volcano growth on a low-friction décollement reproduce 

this bench structure, as well as much of the morphology 

and patterns of faulting observed on the actively growing 

volcanoes of Mauna Loa and Kīlauea. These models show 

how stress is stored within growing volcano flanks, but not 

how rapid, potentially seismic slip is triggered along their 

décollements. The imbalance of forces that triggers large, 

rapid seaward displacement of the flank after decades of creep 

may result either from driving forces that change rapidly, such 

as magma pressure gradients; from resisting forces that rapidly 

diminish with slip, such as those arising from coupling of 

pore pressure and dilatancy within décollement sediment; or, 

from some interplay between driving and resisting forces that 

produces flank motion. Our understanding of the processes of 

flank motion is limited by available data, though recent studies 

have increased our ability to quantitatively address flank 

instability and associated hazards.

Introduction

The southern end of the Hawaiian-Emperor volcanic chain 

(fig. 1) includes some of the largest volcanoes on Earth. The 

sheer size of these volcanoes is enabled by the development 

of long rift zones that extend outward from the summits for 

100 km or more (Fiske and Jackson, 1972). These rift zones 

grow preferentially upward and outward to the seaward side, 

spreading laterally on a décollement formed along the interface 

between the volcanic edifice and the seafloor (Swanson and 

others, 1976). From studies of the active volcanoes Mauna Loa 

and Kīlauea, we know that integral to the growth of these long 

rift zones is persistent seismicity within the flanks and along 

their décollements (Koyanagi and others, 1972), punctuated by 

large earthquakes and, occasionally, by catastrophic collapse as 

a large avalanche or landslide (Moore, 1964; Moore and others, 

1989; Moore and Clague, 1992).

The evidence for catastrophic collapse is clear in maps of 

debris on the deep seafloor surrounding the Hawaiian Islands 

(figs. 1 and 2). The debris forms fanlike aprons connected to 

landslide and slump structures that form large cliffs (or pali) 

along the coasts of each island in the Hawaiian chain (Moore 

and others, 1989). Islands as large as Oʻahu and Molokaʻi are 

inferred to have been entirely dissected, with huge landslide 

scars forming the prominent pali for which the islands are 

justly famous. The debris scattered across thousands of square 

kilometers of ocean floor are the missing portions of these 

islands, and some debris consists of subaerially deposited 

Chapter 4

Instability of Hawaiian Volcanoes

By Roger P. Denlinger1 and Julia K. Morgan2
150 Characteristics of Hawaiian Volcanoes

lava (Moore and others 1994a,b). Dating of this seafloor 

debris, as well as dissected lava flows on the islands, provides 

constraints for studies that show that these landslides occur 

episodically during growth of Hawaiian volcanoes and rift 

zones (Moore and others, 1989), with the largest landslides 

reserved for the latter stages of growth (Moore and Clague, 

1992; Clague and Sherrod, this volume, chap. 3).

Indications of flank instability on active Hawaiian 

volcanoes come from both geologic structures and historical 

observations. On the Island of Hawai‘i, large, rapid flank 

movements (often occurring with large earthquakes) were 

observed four times during the 19th and 20th centuries, each 

spaced about 50 years apart. The most spectacular examples 

are the great 1868 Ka‘ū earthquakes, which opened rift 

zones on Kīlauea and Mauna Loa and moved the southeast 

flanks of both Kīlauea and Mauna Loa seaward in two great 

earthquakes spaced days apart (Wyss, 1988). Large flank 

movements occurred again in 1924 and 1975, preceded by 

decades of slow growth of summit and rift magma systems 

(Swanson and others, 1976). Current observations of gradual 

seaward flank growth on Kīlauea support the interpretation 

that the East Rift Zone (ERZ) episodically opens as the flank is 

gradually compressed by deep magma and cumulate intrusion 

and that this compression also triggers décollement slip, which 

propagates the south flank seaward (Montgomery-Brown and 

others, 2011). Similar patterns of summit and rift growth, 

followed by flank spreading, are displayed by detailed models 

of progressive volcano growth. These models simulate many of 

the geologic structures observed on the south flank of Kīlauea, 

Moloka‘i

Lāna‘i

O‘ahu

Hawai‘i

Maui Maui Fracture Zone

Moloka‘i Fracture Zone

Hawaiian

Trench

160° 158° 156° 154° 152° 150°

19°

18°

21°

20°

23°

22°

24°

Kaua‘i

Elevation,

in kilometers

EXPLANATION

Mauna Loa

Mauna Kea

0

50

50

0

100 150 200 KILOMETERS

100 MILES

3

2

1

0

-1

-2

-3

-4

-5

Kaho‘olawe

Lō‘ihi

Kīlauea

but they do not detail the triggers of rapid flank movement or 

wholesale flank collapse. Our understanding of the mechanics 

of rapid flank movements and what drives them is incomplete.

What conditions result in flank instability, or an imbalance 

between forces driving seaward flank motion, relative to forces 

resisting flank motion, as islands grow? Is rift-generated flank 

instability the cause of landslides large enough to dissect 

entire islands? How can large landslides or avalanches occur 

catastrophically when subaerial slopes are moderate and the 

seafloor actually slopes inward towards the center of each island 

as the weight of the island depresses the oceanic lithosphere? 

These questions have motivated research into Hawaiian Island 

stability since discovery of large landslide deposits on the 

seafloor more than 20 years ago (Moore and others, 1989), but, 

despite exhaustive studies, remain largely unanswered. These 

questions are important for hazards, as well as research, because 

we know that active volcanoes on the Island of Hawai‘i are 

deforming and potentially could one day produce a landslide 

large enough to dissect the Island of Hawaiʻi. However, we do 

not know what to look for, or where we are in the evolution of 

what appears to be, given abundant evidence for flank failure 

along the Hawaiian chain, a characteristic and ubiquitous 

pattern of volcano growth and decay (Clague and Sherrod, this 

volume, chap. 3).

We show, in this chapter, that the observed pattern of 

decades-long subaerial flank compression and uplift, followed 

by rapid flank extension and summit subsidence, is part of a 

systemic pattern of growth of Hawaiian volcanoes. We investigate 

these growth processes by drawing on onshore and offshore field 

Figure 1. Bathymetric map of the Hawaiian Islands, showing debris scattered around pediments on which islands sit. All islands 

have significant aprons of landslide debris associated with previously active volcanoes. Significant deposits are also associated 

with currently active Mauna Loa volcano on the Island of Hawai‘i.
Instability of Hawaiian Volcanoes 151

studies of Kīlauea and Mauna Loa, mechanical models that are 

analogs for flank motion, and sophisticated models for volcano 

growth, flank failure, and décollement resistance. We find that 

the processes of volcano growth migrate rift-zone flanks seaward 

and are associated with a cyclic evolution of flank instability. The 

mechanisms of seaward flank migration involve all of the basic 

structural elements of Hawaiian volcanoes and provide clues to 

promising paths for future research.

Geologic Evidence for Flank Instability

Subaerial Evidence

Although evidence that the flanks of Hawaiian volcanoes 

could be unstable and collapse catastrophically was present 

in subaerial morphology and structures (that is, fault-bounded 

steep cliffs, or pali, along the coasts), the significance of these 

structures was not fully appreciated until the late 20th century. 

In their prescient model for motion of the south flank of Kīlauea, 

Moore and Krivoy (1964) (1) attributed the ERZ eruptive 

fissures formed during the 1962 eruption to be a continuation 

of a listric Koa‘e Fault System, (2) identified antithetic faulting 

-6 -4 -2 2 4

Elevation, in kilometers above sea level

0

Kohala

Mauna Kea

Hualālai

South

Kona

slide

South

Kona

slump

‘Ālika 2

slide

‘Ālika 1

slide

North

Kona

slump

SWRZ

Cape

Kumukahi

‘Āpua

Point

Mauna Loa

Kahuku

Fault

Kealakekua

Fault

Maunaiki

Kīlauea

Lō‘ihi

Punalu‘u

slump

Papau 

Seamount

Hilina

slump

ERZ

Hilina

Pali

Mauna Koa‘e Ulu

19°

19°30'

20°

18°30'

156°20' 156° 155°30' 155° 154°30' 154°

20°30'

EXPLANATION

0

0

25 50 KILOMETERS

10 20 30 MILES

in the Koa‘e system, (3) suggested that the ERZ was fed by the 

same source of magma that fed the summit, and (4) went on to 

suggest that the Hilina Fault System is also a seaward-dipping 

listric fault. They interpreted all seaward-facing subaerial flank 

scarps as listric faults connected to a décollement under the 

south flank at depths well below sea level. They further 

suggested that southward, gravity-driven flank motion opened 

the ERZ and allowed it to be fed from summit magma chambers. 

Although not all of these assertions have withstood continued 

study, many aspects of their model are contained in modern 

concepts of south-flank deformation (fig. 3).

The concept of a magmatically driven rift and flank was 

developed by Swanson and others (1976) using trilateration, 

triangulation, and spirit leveling to constrain compression and 

uplift of the south flank of Kīlauea between the summit and 

the coast, using geodetic data collected from 1896 to 1975. 

They then correlated these changes to observed changes in 

the magmatic system of Kīlauea. Between 1924 and 1968, the 

lava level in Halema‘uma‘u Crater, within Kīlauea’s summit 

caldera, gradually increased from 600 to 100 m below the 

rim (Macdonald and others, 1983). The high lava level was 

associated with a voluminous eruptive output at Kīlauea in 

the 1960s and early 1970s, including an active lava lake at the 

summit during 1967–68, frequent summit eruptions, eruptions 

combined with flank movement that opened both the Southwest 

Figure 2. Map 

showing the Island of 

Hawai‘i, surrounding 

seafloor, structures 

off the south flank of 

Kīlauea, and locations 

of features indicated 

in the text. ERZ, East 

Rift Zone; SWRZ, 

Southwest Rift Zone.
152 Characteristics of Hawaiian Volcanoes

Rift Zone (SWRZ) and ERZ, growth of the Mauna Ulu eruptive 

vent on the ERZ during 1969–74, and rates of summit and 

rift-zone extension unseen today (Swanson and others, 1976). 

This led Swanson and others (1976) to predict that a large flank 

movement would occur, and their intuition was rewarded when 

their paper had passed review but was still in press. Repeating 

a pattern established during the previous two centuries, the 

elevated level of summit eruptive activity and extrusive growth 

ended abruptly in late 1975 with a great earthquake and flank 

movement. The 1975 Kalapana earthquake (M7.7; Nettles and 

Ekström, 2004) produced a maximum of 10 m of seaward flank 

motion (Lipman and others, 1985) and a tsunami (Ma and others, 

1999). In the next 7.2 years, until the eruption of Pu‘u ‘Ō‘ō in 

January 1983, only 3 small eruptive events occurred, whereas 

14 intrusive events occurred at the summit and along the ERZ. 

This ratio of eruptive to intrusive events was the inverse of that 

established in the decades-long increase of magma-system head 

in the summit region (Macdonald and others, 1983; Dzurisin 

and others, 1984), demonstrating that a significant change in 

behavior had occurred.

On Kīlauea, this sequence of gradual growth and inflation 

of the summit, followed by rapid flank movement that collapses 

the summit, opens the rift zones, and drains the magma system, 

is a pattern characterizing large, rapid flank movements in 

1823, 1868, 1924, and 1975 (Swanson and others, 1976; 

Macdonald and others, 1983) and probably also occurred on 

Kīlauea sometime between 1750 and 1790 (Macdonald and 

others, 1983). The great 1868 M7.1 and M7.9 earthquakes 

and associated flank movements affected both Kīlauea and 

Mauna Loa, opened both rift zones on Kīlauea and the SWRZ 

of Mauna Loa, and produced coastal extension and subsidence 

from Punalu‘u to Cape Kumukahi, as well as a tsunami that 

killed 46 people along the island’s southeast coast near Punalu‘u 

(Macdonald and others, 1983). The 1924 flank movement on 

Kīlauea was associated with abundant seismicity and strong 

ground motion that progressed rapidly from the summit 

eastward, along the ERZ to Cape Kumukahi, without a large 

earthquake (Macdonald and others, 1983) but caused extension 

0 100

-8

-4

0

4

-12

Kīlauea

Koa‘e

FS

Ka‘ōiki

FS

Outer Sea level

bench

NORTHWEST SOUTHEAST

Earthquakes and slow slip

Mauna

Loa

Flowing

cumulate

Shallow

landslide

Hilina

Dikes

FS

20 40 60 80

Lō‘ihi

Elevation, in kilometers

Distance along section, in kilometers

across the entire ERZ, as well as coastal subsidence. Subsidence 

of Cape Kumukahi and encroachment of the sea, a kilometer 

up the rift, combined with the 60° dip of the faults flanking the 

opening of a 1-km-wide graben at Cape Kumukahi, suggests 

that about 4.6 m of seaward (southward) migration of the south 

flank of Kīlauea occurred at that time. With the opening of 

this graben and the ERZ, the level of lava at Kīlauea’s summit 

dropped more than 500 m, well below the water table, resulting 

in violent phreatic eruptions several weeks later in May 1924 

(Swanson and others, 2012). It was not until the 1960s that 

subsequent summit growth regained the magma levels last seen 

in 1924 (Swanson and others, 1976), and summit eruptions 

again became commonplace. As in 1924, high magma levels 

and extrusive activity did not last, as rapid extension of the flank 

associated with the 1975 M7.2 earthquake dropped the magma 

level again and ushered in a period of intrusive growth.

Submarine Evidence

The realization in 1975 that previously identified landslidelike fault scarps on the flanks of Hawaiian volcanoes could 

slip suddenly during large flank movements, first proposed 

by Moore (1964), helped to motivate mapping of the seafloor 

surrounding the islands. In fall 1988, side-scan GLORIA 

(Geologic LOng-Range Inclined Asdic) surveys were completed 

in a zone extending from south of the Island of Hawai‘i to 

north of Kaua‘i. In this region, deposits of 17 well-defined large 

landslides were identified (Moore and others, 1989), with many 

overlapping adjacent deposits. Of particular note, the surveys 

near O‘ahu and Moloka‘i revealed deposits from some of the 

largest landslides on Earth. Sampling of these deposits indicated 

that landslide debris had moved as far as 200 km from its 

subaerial source, had crossed a trough in the Hawaiian Trench 

(the depression of the oceanic plate as it sags under the weight 

of the islands), and had traveled upslope on the opposite side. 

In other slide deposits, curved scars in ocean-floor sediments 

showed the track of slide debris. In a landslide track off the 

west coast of the Island of Hawai‘i, the track curves and points 

Figure 3. Cross section diagram showing interpreted internal structure for the south flank of Kīlauea Volcano, modified from Morgan (2006). High-angle 

normal faults underlie summit region and accommodate subsidence and axial extension. Intrusion and subsidence along rift axis and seaward flow of 

ductile cumulates precipitated from magma chambers help to drive unbuttressed flank seaward. Stress transfer from the magma system to the distal end 

occurs through flank slip along a basal décollement, producing earthquakes and slow slip events. This basal slip terminates by thrust faulting, building an 

overthrust structure within the outer bench. Shallow faults on seaward slopes define surficial slump features, such as the Hilina Fault System. 
Instability of Hawaiian Volcanoes 153

toward Lāna‘i and Kaho‘olawe, where cobbles of coral reef 

are found stranded hundreds of meters above sea level (Moore 

and Moore, 1984). This and other evidence are explained by 

catastrophic failure of the flanks of Hawaiian volcanoes in 

which the momentum of landslide debris not only carries it far 

from the volcano, but also generates tsunamis large enough 

to break up coral reefs and deposit fragments above sea level 

(Moore and others, 1992). The GLORIA surveys show that both 

submarine and subaerial evidence are required to interpret the 

genesis of the morphology and structure of Hawaiian volcanoes.

Beginning in the 1990s, various studies of the submarine 

flanks of the Hawaiian Islands gradually tested and refined 

previous interpretations for volcano-flank deformation and 

evolution. Submersible surveys and dredge hauls of submarine 

landslide blocks west of the Island of Hawai‘i confirmed the 

subaerial origin of these blocks and provided the first age 

constraints on the deposits (Moore and others, 1995). Seismic 

surveys reported by Smith and others (1999) revealed the 

presence of thick, landward-tilted sedimentary deposits on 

Kīlauea’s offshore bench, consistent with thrust faulting at the 

toe of the flank. High-resolution bathymetric data collected 

by the University of Hawai‘i, Mānoa, around the Island of 

Hawai‘i, clarified the morphology of the submarine flanks 

(Moore and Chadwick, 1995), leading to morphotectonic 

interpretations that helped extend surficial structures offshore 

(for example, the Punalu‘u slide, Hilina slump, and South 

Kona landslide complex); however, these interpretations were 

limited by a lack of subsurface imaging, as well as few direct 

samples. This situation changed in the late 1990s as a result of a 

comprehensive multichannel seismic reflection survey offshore 

the south flanks of Kīlauea and Mauna Loa (Morgan and others, 

2000, 2003a,b; Hills and others, 2002), seafloor mapping surveys 

carried out by the Japan Agency for Marine-Earth Science and 

Technology (Smith and others, 2002; Eakins and others, 2003), 

and multiple manned and unmanned submersible dives around 

all of the islands (Lipman and others, 1988; Lipman and others, 

2000; Naka and others, 2000; Takahashi and others, 2002; 

Coombs and others, 2004; Morgan and others, 2007).

The morphologic features revealed by the new seafloor 

maps clarified the distribution and geometry of landslides 

around the islands, leading to reinterpretations of the nature 

and origin of several landslides previously based primarily on 

GLORIA imagery (Takahashi and others, 2002). In particular, 

the large Nu‘uanu and Wailau landslides, which broke away 

from the north flanks of O‘ahu and Moloka‘i, respectively, and 

the South Kona landslide complex, west of Mauna Loa, were 

shown to consist of coherent, dispersed megablocks that could 

be reconstructed back onto adjacent, broken flanks (Moore and 

Clague, 2002; Yokose and Lipman, 2004). The more proximal 

portions of the deformed flanks exhibited broad benches 

similar to the bench offshore from Kīlauea.

These characteristics reaffirmed, for several different 

settings, the flank model proposed by Denlinger and Okubo 

(1995). In their model, the south flank of Kīlauea moves 

seaward by gradual, long-term creep that builds broad 

submarine benches braced by imbricate thrusts that build an 

outer bench along the distal end. They proposed that formation 

of this structure is accompanied by intermittent seismicity and 

is occasionally punctuated by rapid catastrophic slip of the 

seaward flank that generates large earthquakes.

This general picture was further confirmed by seismicreflection profiles obtained along the south flanks of Kīlauea 

and Mauna Loa and the north flank of O‘ahu (Moore and 

others, 1997; Morgan and others, 1998, 2010; Hills and others, 

2002). All of these volcano flanks are underlain by a deep 

reflective surface, interpreted to define a décollement coincident 

with the top of the oceanic crust (ten Brink, 1987; Morgan 

and others, 2000; Leslie and others, 2004). The best imaging 

comes from Kīlauea’s south flank, which exhibits seawardvergent thrust faults rising beneath the prominent midslope 

bench (Morgan and others, 2000, 2003a). Evidence that the 

broad benches have undergone progressive uplift and rotation 

indicates that they are probably large overthrust structures 

(Smith and others, 1999; Morgan and others, 2000; Hills and 

others, 2002). These structures have accommodated tens of 

kilometers of lateral displacement (Morgan and others, 2000, 

2003a). Thus, Kīlauea’s offshore bench has formed slowly, 

while accumulating a thick infill of sediment in the closed 

midslope basin formed by extension on its landward side (fig. 

4). This model is more consistent with contraction at the toe of 

the volcano flank in response to its outward displacement above 

a continuous décollement (fig. 4), as originally suggested by 

Denlinger and Okubo (1995), and less compatible with more 

modest displacements proposed by models for gravitational 

slumping. A similar model may hold for several benches that 

compose the Kona landslide complex along the west flank of 

Mauna Loa (Morgan and Clague, 2003; Morgan and others, 

2007), although controversy persists, owing to the lack of 

high-quality seismic-reflection data and sparse geochemical and 

stratigraphic data (Lipman and others, 2006).

New observations also confirm that the offshore benches 

at Kīlauea and elsewhere are constructed largely of fragmental 

debris shed as lava deltas build and then break off at the 

shoreline (Moore and Chadwick, 1995; Moore and others, 

1995; Clague and Moore, 2002; Clague and others, 2002; Naka 

and others, 2002; Morgan and others, 2007). Submersible 

surveys of the outer benches revealed thick accumulations of 

indurated volcaniclastic materials composed of subaerially 

derived fragmental lavas, shallow landslide debris, or both. 

Seismic-reflection profiles over Kīlauea’s midslope basin 

and upper flanks revealed buried normal faults and intensely 

folded layering—a juxtaposition of normal and reverse faulting 

that may denote past landslide structures or reveal complex 

deformation within the currently mobile flank (fig. 4; Hills 

and others, 2002; Morgan and others, 2003b). Volcaniclastic 

deposits of mixed compositions and diagenetic states make up 

the incised bench of Mauna Loa’s Kona landslide complex, 

suggesting a long history of accumulation and burial, followed 

by exhumation and uplift (Morgan and others, 2007). Mauna 

Loa’s upper submarine slopes, in contrast, are commonly 

composed of subaerially derived lava flows, as well as 

volcaniclastic deposits (Garcia and Davis, 2001), forming an 
154 Characteristics of Hawaiian Volcanoes

Figure 4. Maps and diagrams showing structural interpretations of Kīlauea’s south flank, based on seismic-reflection profiles (compiled from 

Morgan and others, 2003a). A, Shaded slope and bathymetric map of south flank of Kīlauea Volcano, showing locations of seismic-reflection 

lines 14, 15, 21, and 22 from 1998 survey. Southwest edge of Kīlauea’s mobile flank is bounded by an offshore lineament defined by ridges and 

scarps, including Papau Seamount (PS) and Nali‘ikakani Ridge (NR). Upper part of submerged flank is marked by two embayments: Halapē Bay 

(HB) shoreline reentrant, between Nali‘ikakani Point (NP) and ‘Āpua Point (AP); and a central flank embayment above a midslope basin. A broad 

outer bench fronts the midslope basin, which is partly filled with volcaniclastic sediment. Steep slope of outer bench is incised by several arcuate 

scarps. Bathymetry gridded at 1,000 m from Smith and others (1994). B, C, Interpreted depth sections for dip-parallel seismic-reflection lines 15 

and line 21. Reflections: D, décollement; G, internal glide plane; OC, top of oceanic crust; T, thrust fault. Transects show contrasting structure: line 

15 shows the more coherent Hilina slump (pink), whereas disrupted strata (yellow) underlie bedding-parallel slope and basin sediment (green) on 

line 21. Blue shows common imbricated stack of accreted volcaniclastic debris. D, E, Cutaway views through Kīlauea’s south flank (looking north) 

showing subsurface structures compiled from seismic lines mapped in part A. Intersection of lines 14 and 15 (part D) reveals structure of west 

flank, detachment G1, and the Hilina slump (pink). Intersection of lines 22 and 21 (part E) shows uplift and westward thrusting of Papau Seamount 

(PS) due to oblique convergence of the Hilina slump on western boundary fault. Transition to region of central flank failure (yellow) is marked by 

an arcuate scarp at seafloor and listric G2 detachment at depth. Imbricate thrust sheets within outer bench (blue) front central flank embayment 

with ponding sediment (green) within midslope basin. D, décollement; ERZ, East Rift Zone; L, left-hand boundary and western boundary fault; 

OC, top of oceanic crust; SWRZ, Southwest Rift Zone; T, thrust fault.

Outer bench

Upper flank

Toe

0

2

4

6

8

Depth (km)

Décollement Ocean crust

Hilina

slump

Landslide

block

Accreted

volcaniclastic 

strata

Detachment

B Line 15

G1

D OC

T1

T2

T3

T4

G2

D OC

T1

T2 T’3

T4

T3

G1

D OC

T1

T2

T3

T4

Basin axis

Upper flank Outer bench

Toe

Old Midslope basin

landslide

deposit

Hawaiian

Trench Detachment

Accreted

volcaniclastic 

strata

Décollement Ocean crust

Line 21

Depth (km) 4

6

8

0

2

C

G2

D OC

T1

T2 T3

T4

T3

G2

D OC

T1

T2 T3

T4

T3

-4000

-5000

154 ̊40'

-5250

-5000

155 ̊20' 155 ̊

19 ̊20'

19 ̊

-1000

-2000

-3000

-5250

-2000

-5000

-1000

Kīlauea

summit

NP Midslope basin

Landslide

blocks

Outer bench

Lō‘ihi

Upper flank

Line 21

Line 15

AP

ERZ

SWRZ

HB

PS

Toe

NR

A

Hilina FaultSystem

Line 14

Line 22

0

0

2.5 5 7.5 10 12.5 KILOMETERS

1 2 3 4 5 6 7 8 MILES

0

0

5 10 KILOMETERS

2 4 6 MILES

Projected basin

axis strata
Instability of Hawaiian Volcanoes 155

interbedded package that appears to drape the distal offshore 

bench (Morgan and Clague, 2003; Morgan and others, 2007).

An integrated view of the evolution of the mobile flanks 

of Hawaiian volcanoes, best exemplified by Kīlauea’s south 

flank, implies a complex but repeatable history (Denlinger and 

Okubo, 1995; Morgan and others, 2003a,b). Debris derived 

from volcano growth is shed into a moat along the volcanoesʼ 

seaward flank, and then pushed seaward by continued flank 

motion. Flank motion plows the sediment, shortening and 

uplifting a sediment pile along the distal margin through 

formation of thrust faults that originate wherever décollement 

slip terminates. It is this process that builds an outer bench. 

Flank motion through décollement slip is very efficient and 

extends the flank more rapidly than sediment can accumulate; 

thus, a basin forms behind the bench within which subaerially 

derived volcaniclastic sediment is trapped before it is shoved 

seaward and deformed, recording a history of deposition and 

deformation that can be read in seismic-reflection profiles and 

stratigraphic records (fig. 4).

This general model for submarine flank deformation and 

evolution may extrapolate well to other volcanoes with less 

complete datasets than Kīlauea. Mauna Loa’s west flank, in 

154°40'

0

2

4

6

8

Depth (km)

0

2

4

6

8

Depth (km)

5 km

5 km

Hilina Fault System

Kīlauea

summit Ka‘ōiki Fault System

ERZ

Midslope basin

Mauna

Loa

Puna peninsula

Outer bench

ERZ

Landslide block

E

Line 22 Line 21

Northeastern

flank Central flank

PS

embayment

Shelf

Western

flank

G1 G2 L

D D

T1

T2

OC

OC

T

Mauna Loa flank T

T3

T4

T?

WEST

SOUTH Not to scale

Not to scale

EAST

WEST

EAST

154 40'W

0

2

4

6

8

Depth (km)

0

2

4

6

8

Depth (km)

West boundary fault

West boundary fault

OC

D 5 km D

Line 14 5 km

D

Line 15

Western

flank

Ka‘ōiki Fault System

Hilina Fault System ERZ

Midslope basin

Mauna

Loa

Puna peninsula

Outer bench

Mauna Loa

flank

ERZ

Landslide block

Northeastern

flank Central flank

embayment

Shelf Nali‘ikakani

Ridge

NORTH

Kīlauea

summit

G1 L

T1 T2

OC

T

T3

T4

G1

T

Primary volcanic edifice

Active Hilina slump

Central flank failure

Accreted volcaniclastics

Post-failure sediments

EXPLANATION

155°20' 155°10' 155° 154°50'

155°20' 155°10' 155° 154°50'

19°30'

19°20'

19°10'

19°

18°50'

19°30'

19°20'

19°10'

19°

18°50'

particular, has been the focus of intense study (Lipman and 

others, 1988; Garcia and Davis, 2001; Morgan and Clague, 

2003; Yokose and others, 2004) in efforts to constrain the 

origin and relative timing of multiple interpreted landslide 

deposits (for example, the South Kona and ‘Ālika landslides, 

fig. 2). The presence of subaerially derived lava flows deep 

on the submarine flanks of western Mauna Loa and the 

radiometric ages of subaerial lavas point to phasing of flank 

activity similar to that at Kīlauea (Morgan and others, 2007). 

In particular, it is hypothesized that a large-scale subaerial to 

submarine flank failure contributed the large blocks that make 

up the South Kona debris field (fig. 2), preconditioning Mauna 

Loa’s west flank to spread outward and bulldozing the newly 

deposited debris to build the South Kona bench. This bench 

was subsequently breached by a smaller landslide, leaving 

behind a narrow scar and debris track that leads to the lobate 

‘Ālika 2 debris field to the northwest of the source region. A 

similar sequence of events may have resulted in the present 

benchlike morphologies of O‘ahu and Moloka‘i’s north and 

south flanks (Nu‘uanu and Wailau structures, respectively; 

Moore and others, 1997). Limited sampling of these features 

during submersible dives suggests that they are also composed 
156 Characteristics of Hawaiian Volcanoes

of heterogeneous volcaniclastic materials (Clague and others, 

2002) that may have been reworked during flank spreading and 

now sit high on submarine benches. Both Nu‘uanu and Wailau 

may have been subjected to subsequent breakup and local 

slumping to form the blocky debris-avalanche deposits, similar 

to, but on a much larger scale than, the late-stage incision of 

Mauna Loa’s western submarine bench.

Geophysical Evidence for Flank 

Instability

Seismic Studies

The first comprehensive discussion of the seismicity 

of Kīlauea and Mauna Loa volcanoes by Klein and others 

(1987) segregated the seismicity of these active volcanoes 

into three parts, based upon the locations of earthquakes in the 

edifice. Shallow (<5 km below ground level [bgl]) seismicity 

is associated with summit and rift-zone eruptions and shallow 

intrusions (fig. 5A), deeper (>5 km and < 10 km bgl) seismicity 

is associated with magma storage and flank deformation 

(fig.5B), and very deep (15–40 km bgl) seismicity is associated 

with magma supply through the mantle and oceanic crust to the 

volcanic edifice (Wright and Klein, 2006). Seismicity produced 

by flank deformation in the depth range of 5–10 km bgl within 

Kīlauea and Mauna Loa volcanoes is offset from their rift 

zones (Wolfe and others, 2007). As shown in figure 5B, this 

seismicity within the flank of Kīlauea is distributed 2 to 8 km 

away from the ERZ. This distributed activity results from both 

internal deformation of the flank (including movement along 

fault systems forming scarps, such as the Hilina Pali), as well 

as slip along a décollement defining the base of the mobile 

south flank of the volcano (Delaney and others, 1998).

Shallow (<5 km bgl) seismicity along the summits and 

rift zones of Mauna Loa and Kīlauea volcanoes is associated 

with shallow intrusive or eruptive activity (Klein and others, 

1987). In contrast to deeper flank seismicity, interpretation of 

the causes of this seismicity is usually unequivocal because 

the occurrence of these events at the summit or along the 

rift zones is directly associated with eruptive or shallowintrusive activity that cracks and faults the surface. The effect 

of the opening of a rift zone at shallow (<5 km bgl) levels 

on stability or deformation of the adjacent flank, however, is 

less clear (Wolfe and others, 2007), as studies of long-term 

deformation and seismicity at Kīlauea demonstrate (Delaney 

and others, 1998; Owen and others, 2000).

There is a complex interplay between the opening of rift 

zones and south flank deformation that was first interpreted on 

the basis of seismic observations (Koyanagi and others, 1972), 

before deformation could be measured on short time scales 

(Swanson and others, 1976). Those authors noticed that riftzone intrusions or eruptions were commonly associated with 

deeper flank seismicity, though their speculation as to cause 

and effect was hampered by lack of knowledge of the detailed 

deformation field that accompanied these earthquakes. More 

detailed studies of seismicity in the past decade (Brooks and 

others, 2006; Wolfe and others, 2007) have confirmed earlier 

suspicions of links between the rift zone and flank seismicity 

(Dvorak and others, 1986), but the addition of detailed 

deformation measurements and combined modeling of 

seismicity and deformation are required to adequately address 

this coupling.

Deformation Studies

Macdonald and Eaton (1957) realized that significant 

ground deformation accompanied the movement of magma 

within Kīlauea Volcano, and both Stearns and Clark (1930) 

and Stearns and Macdonald (1946) recognized that the Hilina 

Fault System on the flank of Kīlauea resembled the headwall 

of a large landslide. Other studies have constrained coastal 

deformation associated with eruptions or with earthquakes. 

Following the great 1868 Kā‘u earthquakes, Titus Coan 

(Brigham, 1909) reported that the entire coast from Cape 

Kumukahi to ‘Āpua Point subsided 1 to 2 m and that the 

volumetric eruptive output from Mauna Loa was halved. 

Additionally, aside from a brief summit eruption in 1877 and 

a persistent lava lake in Halema‘uma‘u, Kīlauea also stopped 

erupting until 1919—repeating a pattern established in the 

19th century. However, the true significance of the intriguing 

landslide-like fault structures along Kīlauea’s south flank 

that contributed to this subsidence, and the relation between 

flank movement and volcanic activity, were unrecognized. 

Distributed flank deformation had occurred on such a vast 

scale (tens to hundreds of kilometers of coastline, involving 

much of the south half of the Island of Hawai‘i) that it was 

impossible, with existing measurement capabilities in the 

19th and early 20th centuries, to detect any links between 

coastal warping and nearby volcanic activity. Consequently, 

the relation between large flank motions and great 

earthquakes went unrecognized until 1975.

The beginning of geodetic constraints on both volcano 

and flank deformation at Kīlauea began with regional 

geodetic-control networks set up in the late 19th and early 

20th centuries to connect Hilo to the summit of Kīlauea, 

to Pāhoa (on Kīlauea’s lower ERZ), and to the south coast 

to survey for road construction and property boundaries 

(Swanson and others, 1976). These surveys were initially 

conducted by the U.S. Coast and Geodetic Survey, but 

beginning in the 1950s, the U.S. Geological Survey began 

reoccupying and extending these geodetic networks for use in 

volcano monitoring. Despite the short time scale (50 years) of 

comprehensive long-term monitoring and the coarse standards 

by which changes in distances could be measured with 

triangulation (at best, 1 part in 105), the huge displacement 

rates of the south flank resulted in large signal-to-noise ratios 

and revealed a wealth of information regarding variations 
Instability of Hawaiian Volcanoes 157

19°30 ́

19°

155°30 ́ 155° 154°30 ́

19°30 ́

19°

AB Earthquakes 1970–89

Depth 5–15 km below sea level

Earthquakes 1970–89

Depth <5 km below sea level

EXPLANATION

0025 50 KILOMETERS

10 20 30 MILES

-20 -15 -10 -505

Surface of 6.8 km/s P-wave velocity, 

in km above sea level

Countour interval

Heavy black countour

= 500 m

= sea level

Figure 5. Contour and 

seismicity maps of the Island of 

Hawai‘i. A, Shallow seismicity 

(<5 km below sea level [bsl]) 

at Kīlauea Volcano summit 

aligns with trends of Southwest 

and East Rift Zones and is 

associated with extension and 

eruptions from 1970 to 1989. 

Earthquakes (red dots) are 

commonly associated with 

surface faulting connected with 

opening of summit and rift zones 

during extension associated 

with shallow dike emplacement. 

Also shown is top surface 

of internal P-wave velocities 

>6.8 km/s (from model of Park 

and others, 2009). Where this 

surface intrudes volcanic edifice, 

it is most likely composed of 

dense dunite accumulated 

during volcano growth. Areas 

of high gravity (see discussion 

below and figure 8) that lack 

high P-wave velocities are hot. 

Summit of Kīlauea is also hot, 

but the volume of accumulated 

dunite is apparently large 

enough to create a P-wavevelocity anomaly. B, Same 

P-wave distribution as in figure 5A but showing deeper (>5 

km bsl) seismicity associated 

with seaward flank movement. 

Most of this seismicity is 

associated with sliding of flank 

on a décollement between 

8- and 10-km depth and is 

near or within this décollement. 

Horizontal separations between 

shallow seismicity in part A

and deep seismicity here are 

significant because in gap 

between them, observations 

of dike opening and flank 

movement show that flank 

moves easily on its décollement 

(Montgomery-Brown and others, 

2011). Stress transmission 

apparently occurs easily from 

rift zone to crest of pali overlying 

décollement seismicity.
158 Characteristics of Hawaiian Volcanoes

in flank movement during the 20th century (Swanson and 

others, 1976; Owen and Bürgmann, 2006). Application of 

reasonable geologic constraints to network reduction produces 

a general pattern of flank motion (Swanson and others, 1976; 

Denlinger and Okubo, 1995; Owen and Bürgmann, 2006) that 

is remarkably consistent with that obtained since 1990, using 

sophisticated satellite geodesy (Owen and others, 2000).

These studies broadly constrain distinct regions of 

deformation on Kīlauea. The volcano is buttressed to the westsouthwest by Mauna Loa. Although the flank north of the ERZ is 

pushed northward episodically during shallow rift-zone intrusions 

(for example, Montgomery-Brown and others, 2010, 2011), 

the flank generally moves southeastward and seaward from the 

summit and from the ERZ to the coast. The summit and ERZ, 

though subject to ephemeral eruptive activity, have undergone 

average long-term subsidence and extension since 1975 (Delaney 

and others, 1998). This extension is associated with seaward 

movement of the remainder of the south flank of the volcano, 

accommodated by slip along a décollement underlying the south 

flank and its offshore bench (Owen and Bürgmann, 2006).

As the décollement slips, the summit and ERZ extend. 

The summit occupies a roughly triangular region that has been 

subsiding and extending for most of the past century (Swanson 

and others, 1976), with only a few brief periods of uplift 

associated with rapid summit growth. Extension and subsidence 

abruptly terminate along an eastern margin between Chain of 

Craters Road and Escape Road, where the ERZ trends east from 

the summit to Mauna Ulu. The ERZ at Mauna Ulu, if extended 

westward to the SWRZ near Maunaiki, would form the south 

boundary of extension and subsidence within the largely 

antithetic (north-dipping) Koa‘e Fault System (fig. 2; Duffield, 

1975). Extension and subsidence of the summit region thus 

continue seamlessly southward into the Koa‘e Fault System 

and southeastward into the ERZ (Duffield, 1975), forming a 

continuously extending region whose south boundary extends 

from Maunaiki (SWRZ) through Mauna Ulu (ERZ) to Cape 

Kumukahi (ERZ) (see Poland and others, this volume, chap. 5).

Despite extension along its crest, growth of the ERZ is 

upward through accumulation of lava and seaward through 

flank displacement and rift extension (fig. 6). This onesided growth pattern, first proposed by Swanson and others 

(1976), is likely to be the generic pattern for many oceanflanked rifts on Hawaiian volcanoes (Leslie and others, 

2004) undergoing seaward migration of the rift system 

over time (fig. 6B). Here, the “ridge-push” mechanism for 

seaward motion of the flank (Swanson and others, 1976) is 

an integral part of Kīlauea deformation: ERZ dikes wedge 

the flank seaward, compressing it and stimulating slip on a 

décollement near the base of the edifice at ~8–10-km depth 

(Dieterich and others, 2003; Syracuse and others, 2010; 

Montgomery-Brown and others, 2011), as illustrated in 

figure 6. Using deformation monitoring to estimate strain 

and then comparing the strain with concurrent seismicity, 

Montgomery-Brown and others (2009) showed that the 

link between rift opening and dike injection and flank 

movement by slip along a décollement is complex and not 

N

Early stage

Summit magma

reservoir

Intermediate stage

A ́

Present stage Older dikes

Recent dikes

A

Lava

flows

Lateral growth of East Rift Zone

NORTH SOUTH

1 2 3 4 5

Dike

A A ́

A

B

Figure 6. Schematic diagrams of rift zone/flank movement, from Swanson and others (1976). A, Cross section showing upward and outward growth of 

Hawaiian rift zones. This model and repeated measurements of flank deformation led to forecast of 1975 M7.7 Kalapana earthquake. B, Map view showing 

how migration of a rift zone seaward, relative to its source at summit, leads to bending of rift zone, as illustrated here and by East Rift Zone of Kīlauea.
Instability of Hawaiian Volcanoes 159

Sea level 0

-2

-4

-6

-8

-10

-12

Altitude, in kilometers

Coulomb stress,

in MPa

3

2

4

1

0

-1

-2

-3

-4

M6.1, 1989 M7.7, 1975 Creep

Locked

0 5 10 15 KILOMETERS

0 5 10 MILES

Oceanic crust

0

-2

NORTH SOUTH

NORTH SOUTH

-4

-6

-8

-10

-12

-14

East Rift Zone

Mantle 

Altitude, in kilometers

Mauna Loa

Kīlauea

PACIFIC OCEAN

A

Sea level

Inactive

Inactive Active

B

310

300

290

280

270

Complete Bouguer gravity

anomaly, in mGals

precisely predictable on the basis of time or average slip, and 

illustrates the complexity of décollement slip mechanisms.

Deformation of the entire flank south of the Koa‘e 

Fault System and ERZ undergoes alternating contraction and 

extension in conjunction with slip triggered on patches of 

Kīlauea’s south-flank décollement. Décollement slip occurs 

with large flank earthquakes, such as the 1975 Kalapana event 

(for example, Tilling, 1976; Lipman and others, 1985), during 

persistent microseismicity (for example, Got and others, 1994; 

Wolfe and others, 2004, 2007), and during transient slow slip 

events (Brooks and others, 2006; Montgomery-Brown and 

others, 2009, 2010). The association of décollement slip with 

dike-enhanced rift-zone extension (Delaney and Denlinger, 

1999) is illustrated in detail by combined geodetic and seismic 

studies (Brooks and others, 2006; Montgomery-Brown and 

others, 2009, 2011) that demonstrate a delicate balance among 

rift-zone extension, dike injection, and flank deformation 

coupled to décollement slip.

Seaward of the summit and ERZ, the relation between slip 

on the Hilina and Hōlei Pali Fault Systems and décollement 

slip has been questioned because the 1975 M7.7 earthquake 

produced slip on both (Lipman and others, 1985; Cannon 

and others, 2001). Kinematic modeling of Global Positioning 

System (GPS) displacements suggests that the Hilina Fault 

System has shallow roots, connecting to a low-angle thrust at no 

more than 4-km depth (Cannon and others, 2001; Cervelli and 

others, 2002). This same conclusion appears to be supported 

by consideration of landward surface tilt produced by antithetic 

faulting on the footwalls (Swanson and others, 1976), as well 

as by a force balance for the south flank, based on the method 

of slices (Okubo, 2004) and detailed studies of prehistoric fault 

slip along the Hilina Pali fault scarp, using elastic-dislocation 

models (Cannon and Bürgmann, 2001). Finite-strength, particlebased modeling (Morgan and McGovern, 2005a) yields similar 

results, supporting a shallow faulting interpretation for the 

Hilina and Hōlei Pali Fault Systems.

Figure 7. Schematic diagrams showing 

estimated stresses generated in south 

flank by intrusion along the East Rift 

Zone. A, Bouguer gravity anomaly over 

East Rift Zone and vertical cross section 

of the Kīlauea south flank (from Clague 

and Sherrod, this volume, chap. 3). B, 

Potential for coulomb yield with observed 

seismicity along décollement (from Cayol 

and others, 2000).
160 Characteristics of Hawaiian Volcanoes

Geophysical Evidence for Magma 

Systems Within Hawaiian Volcanoes and 

Their Role in Flank Instability

Studies of the deep seismicity below the Island of Hawai‘i 

suggest that the magma supply to both Kīlauea and Mauna Loa 

volcanoes may have a common source at great depths (see Okubo 

and others, 1997a; Poland and others, this volume, chap. 5). From 

its origin as a mantle hot spot (Klein and others, 1987; Okubo and 

others, 1997a; Wolfe, 1998; Wright and Klein, 2006), the magma 

supply diverges beneath the island into separate magmatic 

plumbing systems (as shown by studies of magma chemistry 

and deep seismicity), although the relation of deep seismicity 

to magma transport beneath the Island of Hawai‘i remains 

controversial (Okubo and others, 1997a; Wolfe and others, 2003, 

2004; Pritchard and others, 2006). Within the island edifice, 

melt storage generates little seismicity (relative to eruptions or 

flank and rift-zone deformation) and so is hard to image (Ryan, 

1988). Thus, the magma plumbing systems for Kīlauea and 

Mauna Loa have been interpreted mainly from gravity studies 

(Kauahikaua and others, 2000) and seismic tomography studies 

(Okubo and others, 1997b; Benz and others, 2002; Baher and 

others, 2003; Park and others, 2007a, 2009), inferred from 

event-relocation studies (Okubo and others, 1997b), and more 

recently, modeled by analyses of surface deformation (see Poland 

and others, this volume, chap. 5). The segregation of seismicity 

into that related only to magma transport and that related only to 

flank deformation is determined by the proximity of earthquake 

locations to known pathways of magma transport beneath the 

summit and rift zones and by the coincidence of these paths to 

sources of gravity and seismic anomalies.

Gravity studies in the Hawaiian Islands (Strange and others, 

1965), as well as more detailed studies over the Island of Hawai‘i 

(fig. 8; Kinoshita and others, 1963; Kauahikaua and others, 2000) 

indicate that high gravity anomalies are commonly associated 

with the summits and rift zones of the volcanoes. Where detailed 

seismic tomography is available, high seismic velocities coincide 

with dense subsurface regions interpreted as the sources of high 

gravity anomalies (Benz and others, 2002; Park and others, 2009), 

as shown in figure 9. On the Island of Hawai‘i, the association 

of high seismic velocity and high density beneath the summits 

and along rift zones is hypothesized to result from the continuous 

precipitation and settling of olivine crystals to form dunite bodies 

of cumulate within or beneath the tholeiitic magma that resides 

within each volcano (Clague and Denlinger, 1994). Geologic 

support for this hypothesis is provided by eruption of picrites, 

commonly containing numerous dunite nodules, 0.05 to 0.1 m 

in size, from deep within the magma systems (Wilkenson and 

Hensel, 1988). A critical attribute of such deep dunite bodies, 

when maintained at high temperatures and replenished by 

magmatic intrusions, is their ability to flow viscously under 

load. As argued by Clague and Denlinger (1994), these ductile 

dunite bodies may play a key role in mobilizing the south flank 

of Kīlauea, helping stress transfer to the distal flank (fig. 7) 

while extending the near-summit regions, such as the Koa‘e 

Fault System. Evidence for efficient stress transfer through the 

creeping dunite is that décollement-related flank seismicity begins 

seaward of these dunite bodies (figs. 5B and 7), whose location is 

interpreted from Bouguer gravity anomalies (figs. 7 and 8).

High gravity anomalies in summit regions and along 

rift zones, which indicate increasing mass at depth, are also 

a measure of gradients in gravitational body forces acting 

within the flank of each volcano. The shape of the gravity 

field at Kīlauea indicates the presence of material (presumably 

cumulates) beneath the summit, Koa‘e Fault System, and ERZ 

that is much denser than in the surrounding flank (Kauahikaua 

and others, 2000). Gradients in the gravitational field show 

directions of body forces within the flank, and, for the south 

flank, indicate whether these forces would help drive flank 

motion seaward. If flank motion is partly driven by its weight, 

then the deformation field of the flank and associated gradients 

in the gravity field should correlate (as shown for southflank movement only in figure 8B). In addition, the pattern of 

subsidence of Kīlauea’s summit within the past few decades 

during prolonged extension of the flank has roughly the same 

shape as the summit gravity anomaly. The steepest gradient, 

along the southern margin of the gravity high over the summit, 

coincides with the south boundary of the Koa‘e Fault System 

and its extension into the ERZ, and the steepest gradient to 

the east and west corresponds to the upper ERZ and SWRZ, 

respectively (fig. 8).

Correlations between seismic-velocity inversions and 

gravity anomalies provide important insights into velocitydensity relations within these volcanoes. Different velocitydensity relations apply to active and inactive rift zones and 

summit areas (Park and others, 2007b). For example, a massive 

seismic-velocity and gravity anomaly detected within Mauna 

Loa’s south flank (fig. 9C) corresponds to a typical highseismic-velocity, high-density relation. In contrast, the high 

seismic velocities beneath Kīlauea’s magmatically active 

summit and ERZ (fig. 9) are associated with lower, but still 

anomalous, gravity highs. One interpretation of this discrepancy 

between the two settings is that regions with typical seismicvelocity/density relations denote solidified dunite cumulate, 

whereas those with lower velocity-density relations indicate the 

presence of interstitial melt within the cumulate, maintained 

by higher temperatures and capable of producing magma that 

can feed eruptions (Park and others, 2007b). If this is true, 

then the combined analysis of seismic-velocity and gravity 

anomalies provides a means of distinguishing old, cold magma 

bodies from young, hot magma bodies, providing a first-order 

constraint on the mobility of these bodies and their potential 

contributions to flank instability.

In summary, geophysical evidence for the distributions of 

high seismic velocity and high density supports the interpretation 

that seismic velocity and density anomalies are associated with 

magma pathways within the volcanoes, whether these pathways 

are active or inactive. The correlation between seismic velocity 

and density is also useful in defining which areas are still hot 

(low velocity-density ratio) and potentially active, even without 
Instability of Hawaiian Volcanoes 161

250

270

250

230

170

210

210

230

190

290

210

230

270

250

230

290

210

230

250

270

230

210

210

190

230

230

210

210

250 230

230

190

210

210

230

250

270

230

230

250

230

Kaho‘olawe

Lāna‘i

158° 157° 156°

21°

21°30'

20°30'

157°30' 156°30'

A

B

19°

19°30'

20°

156° 155°30' 155°

South

Point

Mauna

Loa Kīlauea

Hualālai

Mauna

Kea

Lō‘ihi

Cape

Kumukahi

Bouguer gravity,

in milligals

Elevation contours

at 500 m intervals

Accumulated 

displacement 

from 1970–1989; 

10 m (length) 

EXPLANATION

0

0

10 KILOMETERS

10 MILES

20

PACIFIC OCEAN

Kauaʻi

Niʻihau Oʻahu Maui

Molokaʻi

Kahoʻolawe

Lānaʻi

Hawaiʻi

360

340

320

300

280

260

240

220

O‘ahu

Moloka‘i

Maui

Figure 8. Gravity anomaly maps for the Hawaiian Islands and Island of Hawai‘i. A, Gravity anomalies (in milligals) associated with old rift zones 

on the Islands of O‘ahu, Moloka‘i, and Maui, indicating that these rift zones have a dense core, most likely composed of dunite accumulated during 

their growth. B, Current gravity anomalies associated with active and inactive volcanoes on the Island of Hawai‘i. Dunite accumulates in deep core 

of summits and rift zones when magma circulates through these regions. Magnitude of anomaly is related to length of time that magma circulated, 

and presence of anomalies in other areas (for example, southeast of summit of Mauna Loa) may indicate old rift zones. Vectors on south flank of 

Kīlauea Volcano denote the accumulated displacement from 1970 to 1989, indicating displacement of flank in response to magma pressure, as well 

as the gravitational load of dunite. Directions of gradients of gravity anomalies are indicated by gradients in colors used to display gravity.
162 Characteristics of Hawaiian Volcanoes

Depth (km) Depth (km)

5

0

5

10

15

20

25

Kink Kink Kink Kink Kink

0 10 20 30 40 50 60 70 80 90 100 110

Distance, in kilometers

Distance, in kilometers

KSWRZ KLS KERZ PR

5

0

5

10

15

20

25

MNERZ KERZ HFS

OB

MK-EF

0 10 20 30 40 50 60 70 80 90 100

B B ́

A A ́

-5000

-4000

-3000

-2000

-2000

1000

-1000

2000

3000

3000

-3000

-4000

2000

A

A ́ ML

ML

HL MK

Kil

L

156° 155°30' 155° 154°30'

19°

19°30'

20 ̊

Outer bench

-5000

-4000

-3000

-2000

-2000

-1000

1000

2000

3000

3000

-3000

-4000

2000

1000

B ́

B

HL MK

Kil

Outer bench L

156° 155°30' 155° 154°30'

19°

19°30'

20°

A

B

0

5

10

15

20

25

30

35

Hypocenter depth (km)

EXPLANATION

155°30' 155° 154°30'

19°

19°30'

C

High-velocity anomalies

Ground displacement (2 cm/yr)
Instability of Hawaiian Volcanoes 163

obvious current deformation or seismicity. The correlation 

of gradients in gravity anomalies with active flank velocities 

on the south flank of Kīlauea (fig. 8B) is evidence that the 

flow of ductile dunite cumulates contributes to seaward flank 

motion, because these anomalies record a force imbalance that 

potentially drives flank motion. In this case, the scaling of the 

gravitational forces obtained from the flank density distribution 

is useful for scaling forces driving flank motion and constraining 

mechanical models for flank instability.

Mechanics of Flank Instability

To explore the mechanical constraints on flank 

deformation at Kīlauea, we consider several models for flank 

instability. The simplest model is a rigid wedge sliding on a 

décollement, in which a force balance is derived for the entire 

domain. Adding elasticity to the wedge allows for storage and 

release of elastic strain energy, providing a mechanism for 

coupling between décollement slip and wedge deformation. 

Adding finite strength to the wedge localizes deformation 

within it and limits stress through faulting, which is most 

notably associated with large earthquakes and slow slip 

events. Each model scales the forces driving flank instability 

differently, and we consider each in order of complexity.

The models are constrained by the morphologic, 

structural, and geophysical evidence for flank instability 

outlined above. This evidence indicates phenomena that 

either contribute to, or are a consequence of, the instability 

of Hawaiian volcano flanks (fig. 3), including (1) magmatic 

intrusion, which increases mean stress beneath summits and 

along rift zones, building stress gradients that episodically 

are transferred seaward by slow slip events; (2) episodic 

transfer of stress and deformation seaward by slow slip events 

and small earthquakes, migrating the volcanic pile seaward 

on a persistent, low friction décollement; (3) termination 

of décollement slip in thrust faults that build a distal bench 

from accreted volcanic debris; (4) viscoelastic creep of dunite 

cumulate, adding to slow progressive loading of flanks beneath 

summits and along rift zones between episodes of rapid slip; 

and (5) rapid, large-scale flank movements associated with 

large earthquakes (for example, the 1975 M7.7 earthquake). 

Flank models show that each of these phenomena plays a 

different role in altering the forces affecting flank stability.

A rough estimate of the static balance of forces governing 

flank stability can be obtained from an idealized rigid-wedge 

model derived from the south-flank geometry (Dieterich, 1988; 

Iverson, 1995). For the south flank of Kīlauea, we can assume 

that the back of the wedge approximately corresponds to the 

south boundary of the Koa‘e Fault System and its extension 

along the ERZ, as outlined by Duffield (1975) and illustrated 

in figure 3. The lateral extent of rift extension defines the 

lateral margins of the wedge, and the distal end of the wedge 

corresponds to the distal end of the offshore bench. The base of 

the wedge slides on a décollement, interpreted to coincide with 

a pelagic sedimentary layer sandwiched between the volcanic 

edifice and the Cretaceous seafloor (Nakamura, 1982). This 

layer has been cored and imaged just south of the Island of 

Hawai‘i and determined to be as much as 200 m thick (Leslie 

and others, 2002). The wedge model for the south flank is 

partly submerged, which influences its specific weight and 

applied forces (Iverson, 1995; Dieterich, 1988), as shown 

schematically in figure 3.

The driving forces for a rigid-wedge model, with the 

geometry illustrated in figure 3, are produced by (1) the 

weight of magma and cumulates pressurizing the back of the 

wedge along the rift zone, (2) excess magma pressure, and 

(3) the component of weight acting tangentially along the 

décollement (where the surface slope exceeds the décollement 

slope) (Iverson, 1995). Resisting forces are dominated by (1) 

the component of weight acting normal to the décollement, 

resisting wedge motion through fault friction, and (2) excess 

pore pressures or seepage forces within the décollement that 

offset fault friction.

Limiting equilibrium provides a framework for assessing 

the relative importance of each of the driving and resisting 

forces outlined above. The weight of the flank, combined 

with a wedge-shaped cross section, produces both a shear 

stress and a normal stress across the décollement. In the 

absence of internal deformation, forces driving flank motion 

are resisted only by friction along the décollement (Dieterich, 

1988; Iverson, 1995) formed in the sedimentary layer that 

is overridden by volcano growth. As the flank overrides 

and deforms this sediment, effective friction along the 

décollement is reduced by excess pore pressures generated 

by burial, shearing, and consolidation. For the south flank 

of Kīlauea, rigid-wedge models indicate that zero excess 

magma pressure along the ERZ requires the average effective 

Figure 9. Vertical profiles through P-wave-velocity model derived 

from first-arrival tomographic inversion around Kīlauea and surrounding 

volcanoes (Park and others, 2009). Accompanying maps show profile 

locations, island coastline and 1,000-m bathymetric contours, and main 

geologic features. Abbreviations are: HL, Hualālai; L, Lō‘ihi; MK, Mauna 

Kea; ML, Mauna Loa; Kil, Kīlauea. A, Profile A–A′ along Kīlauea’s 

Southwest Rift Zone (KSWRZ), summit (KLS), East Rift Zone (KERZ), 

and Puna ridge (PR). B, Profile B–B′ crossing Mauna Kea’s east flank 

(MK-EF) and northeast rift zone (MNERZ), KERZ, Hilina Fracture 

System (HFS), and offshore outer bench (OB). White areas, unsampled 

regions of model in vertical profiles. Contour interval, 0.5 km/s, index 

contours labeled in white. Vertical exaggeration 2x. Earthquakes located 

within 2 km on either side of each profile are projected onto profiles 

(black dots). Dashed black lines, interpreted top of oceanic crust and 

lithospheric Moho. C, Maps of southeastern part of the Island of Hawai‘i 

(adapted from Park and others, 2007a), showing ground displacements, 

seismicity, and interpreted extent of high seismic-velocity (~6.5 km/s) 

anomalies near base of volcanic edifice. Red arrows, average horizontal 

surface velocities, with 95 percent confidence ellipses, measured by 

campaign Global Positioning System (GPS) between 1997 and 2002 

(Miklius and others, 2005). Colored dots are M>2.0 earthquakes with 

hypocentral depths of 5–13 km from 1995 to 2005.
164 Characteristics of Hawaiian Volcanoes

internal friction angle along the décollement to be <16° for 

any flank motion to occur (Iverson, 1995). Given typical 

values of internal friction angles for pelagic clays (30°–40°), 

a zero excess magma pressure condition along the ERZ places 

stringent limits on the magnitude of excess pore pressure 

required to produce flank motion.

In contrast to rigid-wedge models, an elastic-wedge 

model for south-flank deformation deforms internally and 

can accommodate limited décollement slip adjacent to the 

rift without requiring either large magma pressures or low 

friction along the entire décollement. At Kīlauea, even small 

dike injections move the south flank, as demonstrated by 

Montgomery-Brown and others (2010). Elastic models (Yin 

and Kelty, 2000) show that excess magma pressures of ~10 

MPa, typical of dike injections along the ERZ (Rubin and 

Gillard, 1997), can produce limited slip on a ~3-km-wide 

swath next to the rift zone, comparable to the region of the 

south flank where rift-zone extension initiates slow slip events 

(Montgomery-Brown and others, 2011). However, we lack 

observations that constrain how stress accumulates along the 

rift zone and propagates seaward through the flank. How do 

repeated dike injections move a 50-km-wide flank seaward?

Scaling provided by wedge models shows that the primary 

resistance to flank motion in response to loading along the rift 

is décollement friction (Iverson, 1995; Yin and Kelty, 2000), 

and the décollement is where most south-flank earthquakes, 

including the largest events, occur. We assume that the excess 

pore pressure, which reduces friction within décollement 

sediment, builds as the sediment is overridden and compressed 

by the flank. Pore pressure within pelagic clays will increase 

through compaction and may be enhanced by fluid circulation 

within the oceanic crust (Christiansen and Garven, 2004). 

Although wedge models idealize décollement resistance in 

terms of some constant average, décollement friction during 

consolidation is likely to vary broadly in space and time, just 

as Iverson (1986) describes for landslide motion. In particular, 

coupling between pore pressure and dilatancy during shear 

deformation can strongly modulate décollement resistance 

to flank motion, just as it does along the base of landslides 

(Iverson, 2005), and the transient nature of this coupling can 

lead to rapid, unstable landslide motion after decades of creep. 

Similar mechanisms have been proposed to act along shallow 

subduction zones (Beeler, 2007).

The time scales for instability obtained by coupling 

dilatancy and pore pressure depend on the material fabric 

within the décollement. With reference to Iverson’s (2005) 

model for landslides, the coupling between south-flank motion, 

shear-induced décollement-zone volume change (dilatancy), 

and décollement pore-pressure change depends on the rate 

of volume change with shear strain and shear strain-rate and 

the intrinsic time scales for pore-pressure generation and 

dissipation within the décollement. Dilatancy produces a 

negative pore pressure feedback that can regulate south-flank 

motion for years. However, if dilatancy decays over time (or 

with increasing shear strain), then pore pressure can suddenly 

increase rapidly after years of creep, generating rapid slip. The 

material properties of the subaerial Minor Creek landslide in 

northwestern California, reported by Iverson (2005), happen 

to be consistent with a mixture of volcaniclastic sediment and 

pelagic clays, giving a wide range of potential time constants for 

instability. The addition of volcaniclastic sediment to clay-rich 

sediment produces a fabric in which dilatancy induced by shear 

will decrease over time. For the values in table 1 of Iverson 

(2005), but with a décollement thickness of 300 m for Kīlauea’s 

south flank rather than 6 m for Minor Creek, continued southflank motion will generate rapid pore pressure growth within 

the décollement after 2 years of steady creep—an intriguing 

prospect, given the observed 2-year frequency of south-flank 

slow slip events (Brooks and others, 2006; MontgomeryBrown and others, 2009). And, just as with large landslides, 

observations of repetitive-slip events support a mechanism 

that either resets fabric dilatancy along the old slip surface or 

generates a new slip surface within décollement sediments.

 Given these considerations and recent observations of 

triggered seismicity associated with slow slip events (Segall 

and others, 2006; Wolfe and others, 2007), such events 

may be an endemic process indicative of flank instability at 

Kīlauea (Cervelli and others, 2002; Brooks and others, 2008; 

Montgomery-Brown and others, 2009, 2011). Slow slip events 

occur updip of documented south-flank earthquakes and are 

presumed to correlate with the transition between seismic and 

aseismic slip, similar to subduction zones (Liu and Rice; 2007), 

where slow slip events may occur both downdip and updip of 

the seismogenic zone (Dragert and others, 2001; Rogers and 

Dragert, 2003; Norabuena and others, 2004; Obara and others, 

2004). At Kīlauea Volcano, slow slip events help translate the 

south flank seaward, that is, in the same direction as overall 

flank motion, and are a record of stress transfer.

Elastic modeling of the locations of slow slip along 

Kīlauea’s décollement (Brooks and others, 2006; MontgomeryBrown and others, 2011) shows termination of slip landward 

of the uplifted outer bench. The bench may restrict further 

transfer of stress seaward, at least intermittently. Indeed, 

mechanical models show that décollement slip will terminate 

by transforming into a thrust fault that curves up into the 

overburden (Mandl, 1988), similar to the thrust faults 

recognized within the outer bench at Kīlauea (fig. 4; Morgan 

and others, 2000, 2003a). The associated bench uplift thickens 

the overburden, increasing resistance to slip along the 

décollement (Mandl, 1988) and further enhancing the tendency 

for thrust faulting and bench growth. New submarine geodetic 

studies (Phillips and others, 2008) support uplift by thrust 

faults in the area where geodetic studies show that décollement 

slip terminates (Montgomery-Brown and others, 2011). Slow 

slip events may thus be delimited by the outer bench along 

Hawaiian volcanoes and may be the manifestation of long-term 

flank-creep-driven upper-flank loading.

The mechanical models discussed above are focused on 

coherent mobile flanks driven by loading along the rift zones; 

however, the structure of the volcano flanks, particularly 

Kīlauea’s south flank, reveals both extensional and contractional 

faults, confirming the importance of internal deformation 
Instability of Hawaiian Volcanoes 165

during flank instability. The relations between flank structure, 

décollement strength, and internal strength can be explored 

during volcano growth by using particle-based numerical 

models. These models, in which volcanoes are constructed 

by the addition of granular material near the summits above 

a frictional décollement, result in outward spreading of the 

volcano flanks, accompanied by faulting (fig. 10). The balance 

of stresses within the flank, governed by the relative strengths 

of the flank and the underlying décollement, defines the final 

geometry and internal structure of the mobile flank (Morgan 

and McGovern, 2005b). In the examples shown here (Zivney, 

2010), low décollement friction (fig. 10A) results in deep-seated 

faulting centered beneath the summit, causing predominantly 

lateral displacement of the volcano flanks, whereas increasing 

basal friction (figs. 10B–10D) results in progressively shallower 

flank deformation, culminating in shallow landslides that cut the 

flanks. The addition of a dense, ductile core (shown in purple in 

figs. 10E–10H), analogous to the dunite cores interpreted within 

Hawaiian volcanoes (Clague and Denlinger, 1994), concentrates 

extensional deformation beneath the summit and is independent 

of décollement strength. Slip becomes localized along the 

décollement fault and is enhanced by summit subsidence. The 

volcano flanks undergo modest extension as they are displaced. 

An important element of these simple simulations is the 

incremental nature of décollement slip and flank deformation 

as stress builds and is released during volcano growth, resulting 

in discrete slip and deformation events (Morgan, 2006). These 

examples offer great potential for constructing more elaborate, 

mechanical-based models to better understand the unique 

processes currently underway at Hawaiian volcanoes.

Discussion

For the past 200 years, observations of the active volcanoes 

Mauna Loa and Kīlauea have revealed repeated episodes in 

which steady growth of the volcano at the summit and along 

the rift zones was facilitated by gradual seaward spreading of 

the flanks and punctuated by episodes of rapid seaward slip. For 

Kīlauea, such large-scale rapid slip occurred on the south flank 

in the second half of the 18th century and in 1823, 1868, 1924, 

and 1975, whereas for Mauna Loa it has occurred only twice, 

in 1868 and 1951 (Macdonald and Wentworth, 1954). Although 

these flank movements were sometimes associated with a large 

earthquake and tsunami, they were always associated with 

extension and subsidence of the coast during seaward movement, 

opening of the rift zones, and, with extension, a subsequent drop 

in magma-system head by hundreds of meters.

During the 1975 M7.7 earthquake on Kīlauea, downward 

and seaward movement of the south flank reversed the coastal 

compression and uplift that had been recorded during growth 

of the volcano since the last large flank movement in 1924, 

when the lava level had dropped by hundreds of meters at the 

summit. As it had in 1924, the lava level dropped again during 

the 1975 earthquake. Large flank movements and sudden drops 

in magma-system head, followed by subsequent refilling of 

the magma system, form a pattern of growth and collapse 

observed at Kīlauea after each large flank movement since 1823. 

During each decades-long increase in magma-system head, 

many smaller episodes of flank deformation occur in which 

stress is transmitted from intrusions along the ERZ to the south 

flank. Before each episode of slip, the summit and rift zones 

are gradually wedged open by magma pressure, presumably 

facilitated by dunite creep at depth, and the adjacent flank is 

compressed by shallow magma intrusion. As the rift is wedged 

open, compression on the seaward flank progressively increases 

until seismic or aseismic slip is triggered on the underlying 

décollement. At Kīlauea, seismicity associated with décollement 

slip begins 2 to 3 km seaward of the rift zone (compare figs. 

5A and 5B; see Montgomery-Brown and others, 2011) and is 

manifested in the deep seismicity mapped in figure 5B.

The offset between the shallow rift-zone seismicity and 

deeper décollement seismicity indicates separate responses of 

the flank to push along the rift zone. This response is intriguing, 

because analogous mechanical models of thrust sheets in which 

stress is measured (Blay and others, 1977; Mandl, 1988) can 

explain this offset. Experiments show that if a thrust sheet rides 

on a frictional base but is underlain by a short section of viscous 

material at the proximal end, as shown in figure 11, then a push 

on the proximal end rapidly transmits stress to the distal end of 

the viscous underlayment. Stress accumulates at the distal end 

of the viscous underlayment (which separates it from where the 

push is applied) until frictional slip begins there. As slip begins, 

stick-slip events then migrate stress along the frictional base 

toward the distal end. Relating this experiment to the south flank 

of Kīlauea, the viscous underlayment is an analog for hot dunite, 

and the frictional base is an analog for the décollement; thus, the 

offset between shallow rift seismicity and deeper décollement 

seismicity appears to be consistent with the mechanical indication 

of a viscous cumulate body adjacent to the rift zone and more 

distal slip along a décollement. A comparison between the gravity 

data and seismicity mapped in figure 8 and the velocity data and 

seismicity mapped in figure 5 also supports this conclusion. The 

shallow rift seismicity overlies one side of the summit gravity and 

velocity anomalies, whereas the deeper décollement seismicity 

forms the seaward boundary of the other side of the gravity and 

velocity anomalies. Slip events initiate at the distal end of the 

hot dunite body and migrate the flank seaward, opening the rift 

zone (Montgomery-Brown and others, 2010). If the intervening 

aseismic zone is composed of hot dunite, then its viscous response 

will not only separate the epicenters of shallow rift seismicity 

from deeper décollement seismicity, but also delay the response 

of the rift zone to slow slip events or the triggering of slow slip 

events by rift-zone loading.

At the distal end of décollement slip, each modeled slow 

slip event (Montgomery-Brown and others, 2010) is likely to 

terminate beneath the offshore portion of the flank undergoing 

uplift (Phillips and others, 2008), consistent with mechanical analysis showing that décollement slip on model thrust 

sheets terminates where thrust faults penetrate into the overburden (Mandl, 1988). A corollary of this model is that if the 
166 Characteristics of Hawaiian Volcanoes Height, in kilometers Height, in kilometers

100

0

100

0

100

0

100

0

100

0

100

0

100

0

100

0

400 300 200 100 0 100 200 300 400 300 200 100 0 100 200 300 400

Distance from model center, in kilometers

A.

B.

C.

D.

E.

F.

G.

H.

Basal friction: 0.0

Basal friction: 0.1

Basal friction: 0.2

Basal friction: 0.3

Basal friction: 0.0

Basal friction: 0.1

Basal friction: 0.2

Basal friction: 0.3

Homogeneous models (internal friction: 0.3)

Cumulate-cored models (internal friction: 0.3)

Particle configuration Incremental displacements

Increasing décollement friction Increasing décollement friction

Figure 10. Cross-sectional diagrams showing particle-dynamics simulations of symmetrical volcanic spreading with and 

without weak cumulate cores (adapted from Zivney, 2010, after Morgan and McGovern, 2005a,b). Interparticle friction noted 

here differs from bulk friction, which is quantified by Morgan and McGovern (2005a). A–D, Effects of basal-friction variation with 

no cumulate (that is, homogeneous models), showing final particle configuration (left) and incremental distortional strain (right). 

Red, right-lateral shear; blue, left-lateral shear. Basal friction increases from A to D, resulting in decreased edifice spreading and 

steeper flanks. Deformation transitions from slip along base for low basal friction, through enhanced normal faulting, to surficial 

avalanching at high basal friction. E–H, Effects of basal friction variation for cumulate-cored (purple) edifices, otherwise similar to 

homogeneous volcanoes. Basal friction increases from figure E to figure H, causing decreased edifice spreading; weak cumulate 

material spreads laterally, pushing flanks outward, relative to homogeneous edifices. Deformation is characterized by enhanced 

slip along base of edifice and within weak cumulate; normal faulting and summit subsidence occur above spreading cumulate, and 

modest surficial landslides still occur. Cumulate-cored volcanoes exhibit dome-shaped morphology with inward-dipping strata.
Instability of Hawaiian Volcanoes 167

thickness of the overburden increases seaward (as in the toe of 

the mobile flank), then the locus of imbricate thrusting from 

the décollement will move landward. If, over decades, many 

repeated slow slip events occur (Brooks and others, 2006) 

and each event builds and uplifts the offshore bench, then, 

as slip events terminate closer to shore, the distal bench will 

grow progressively and inexorably landward over time. Evidence for this process exists in the marine seismic data over 

Kīlauea’s south flank but shows that thrusts may also form on 

the seaward side of the outer bench (Morgan and others, 2000, 

2003a), presumably when décollement slip undercuts the 

bench and terminates there.

The transition of basal slip to thrusts at the distal end of 

the flank is just one component of flank growth over time that is 

represented in particle-based models in which stress is limited 

by a finite-strength criterion and finite strain is allowed (Morgan 

and McGovern, 2005b). These models show that many of the 

structures and fault patterns we observe on Kīlauea and Mauna 

Loa volcanoes evolve over long time periods during volcano 

growth as the flanks spread over a weak décollement (Morgan, 

2006). The correspondence of these models to volcano shape, 

formation of pali on the subaerial flank, shallow landslide 

activity, and formation of a distal bench contributes to our 

understanding of volcano growth.

Migration of lateral stress

Low High

Onshore and offshore studies of the structure of Hawaiian 

volcanoes of different ages, from young (Lō‘ihi) through 

middle-aged (Kīlauea and Mauna Loa) to old (Moloka‘i and 

O‘ahu), and observations of the behavior of Kīlauea and Mauna 

Loa volcanoes, can be incorporated into an inclusive model 

representing the volcanotectonic stages of volcano growth and 

degradation (fig. 12). During the earliest submerged stages 

of shield building (stage 1), the rift zones grow upward and 

outward on their seaward side(s). A modest cumulate core lies 

within the edifice, and the low gravity anomaly over small 

edifices like Lō‘ihi indicates that the push from the cumulate 

at this stage of growth is too small to drive outward spreading 

of the volcano flanks. Steep slopes produce shallow landslides 

along the volcano flanks, building an apron of debris at the distal 

edge of the flank. In stage 2, an enlarged cumulate core can push 

the flank outward, in concert with dike intrusions along the rift 

axis. A décollement fault develops near the base of the edifice, 

allowing the mobile flank to be pushed into the apron of debris, 

which deforms by imbricate thrusting at the distal edge. In stage 

3, the upper parts of the volcano are subaerial, and the seaward 

flank is large and thick enough to sustain shallow landslide 

activity that does not penetrate to the décollement, similar to 

that interpreted for Kīlauea’s south flank (Morgan and others, 

2003a). Décollement slip continues, further building the outer 

Figure 11. Series of 

photographs showing 

stresses in gelatin, capturing 

migration of lateral stress 

(higher stress=lighter color) 

within laboratory model of 

a thrust sheet pushed from 

behind (left edge). Base is 

lubricated from left edge 

to point X. Applied load 

on left edge increases in 

each photograph from top 

to bottom, which are all 

lined up, relative to point X. 

Stress migrates quickly over 

lubricated zone between 

this edge and point X and 

is stored at its distal end, 

where it accumulates, until 

sheet can decouple from its 

base at point X, transferring 

stress farther toward distal 

end in a series of small 

stick-slip events along 

décollement.
168 Characteristics of Hawaiian Volcanoes

bench. Décollement slip is accomplished both by intermittent 

large earthquakes and by slow slip events (Brooks and others, 

2006). Finally, in stage 4, catastrophic failure of a sector of the 

flank may occur, possibly including parts of the outer bench, 

dispersing landslide debris far across the seafloor and leaving 

telltale scars along the volcano flanks. If growth continues at 

the summit and along the rift zones, the cycles begin again.

Many of the processes outlined in the evolutionary stages 

we present here are captured in the particle-based models 

introduced above, although such models are not yet refined 

enough to reproduce specific events or structures. Further 

work using cohesive models with geometries designed to 

match Hawaiian volcanoes could provide additional insights 

into these behaviors. Importantly, however, the specific 

triggers for rapid flank movement are poorly represented in 

such models because these triggers occur on different spatial 

and temporal scales from those represented in the models. 

Despite this deficiency, the particle-based models do show 

how stress is stored and transmitted within the volcano flanks 

and the critical importance of décollement resistance to 

stabilizing the forces that drive flank motion. We therefore 

argue that flank instability is directly related to décollement 

processes and properties that are capable of the rapid changes 

that will influence the resistance to flank motion. Preexisting 

landslide detachments within the edifice will respond to 

similar changes. One process that is expected to occur as 

saturated sediment is sheared is the tight coupling between 

dilatancy and pore pressure. Using the analysis of Iverson 

(2005) as a model for the décollement beneath the south flank 

of Kīlauea, and using his parameters and a thickness (200 

m to 300 m) appropriate for a mixture of turbidites (sandy 

silt) and clay for the pelagic layer overridden by the south 

flank, we obtain time constants for slow creep, terminated 

by rupture and rapid slip. For this mixture and thickness, the 

rapid slip occurs at intervals of one to three years, consistent 

with the observed range of recurrence of slow slip events at 

Kīlauea (Brooks and others, 2006). As a consequence, the 

coupling of induced pore pressure and dilatancy within the 

décollement represents a promising and intriguing prospect 

for future research.

Finally, the near-periodic occurrence of large, rapid flank 

displacements on Kīlauea and Mauna Loa raises the question 

of whether these events are driven by the same mechanisms 

that produce sector collapse, such as those documented along 

the northeast shore of O‘ahu and the north shore of Moloka‘i 

(fig.1), as well as along Mauna Loa’s west flank (Lipman and 

others, 1988; Morgan and others, 2007). Our understanding of 

the scaling driving flank motion, though limited, suggests that 

each large rapid flank movement is just a smaller version of a 

much larger sector collapse. The processes affecting resistance 

to flank motion, such as pore-pressure/dilatancy coupling 

along the décollement, can drastically and rapidly reduce the 

resistance to sliding on time scales comparable to those of the 

slow slip events observed on Kīlauea, but it is unknown how 

this scales up to areas comparable to a large flank earthquake 

or scales by orders of magnitude in size to form a sector 

collapse. In addition, special conditions and circumstances 

are needed to rebuild pore pressure/dilatancy coupling along 

the décollement after each slow slip event, as well as after 

each large catastrophic flank movement. Alternatively, or 

perhaps in association with pore-pressure/dilatancy coupling, 

other processes may contribute to instability through rate- and 

state-dependent variations in friction along the décollement 

(Dieterich and others, 2003). Finally, explosive volcanic 

eruptions, which are known to have occurred in all of these 

settings, provide a potential and unconstrained trigger for largescale sector collapses (McMurtry and others, 1999; Morgan 

and others, 2003a). Further study is needed to determine how 

to constrain these potential mechanisms of flank instability.

Conclusions

Hawaiian volcanoes of different ages exhibit a wealth 

of evidence for past mobility and instability, including large 

debris-avalanche deposits scattered across the seafloor, 

uplifted outer benches at the distal edges of growing volcano 

flanks, proximal fault scarps, and seafloor scars related to 

rapid landslide motion. In recent decades, we have gained 

tremendous insights about the processes responsible for these 

features through improved seafloor mapping, submersible 

surveys of the incised and exposed submarine flanks, marine 

seismic surveys, and seismic-velocity and density modeling. 

The knowledge gained about the internal structures of 

Hawaiian volcanoes has revealed the cumulative effects 

of volcano deformation throughout the growth of these 

volcanoes, suggesting an evolution of processes during the 

growth and collapse of Hawaiian volcanoes. Some of these 

processes promote flank instability, producing episodic and 

rapid flank movements that ultimately may be the cause of the 

huge landslides that dissect entire islands as they grow.

 The active volcanoes Kīlauea and Mauna Loa on the 

Island of Hawai‘i have shown separate volcanotectonic stages 

of growth, but both volcanoes exhibit cyclic variations in rapid 

flank movement during growth. Decades-long growth across 

their summits and along their rift zones is punctuated by brief 

periods during which the seaward flank of the growing rift 

zone rapidly moves seaward in seconds to minutes. In this 

fashion, the growth of rift zones is one-sided, always upward 

through lava accretion and seaward through flank motion, and 

it is to the seaward side that rapid movement and associated 

large-scale gravitational collapse occurs. Sometimes this rapid 

motion is coseismic, generating a great earthquake and tsunami, 

but it always results in opening of the rift zone, subsidence of 

the coast, and collapse of the summit magma-system head by 

hundreds of meters. Whereas idealized rigid and elastic models 

of rift flanks shed light on the relative importance of driving 

and resisting forces that move the flank seaward, successful 

interpretation of flank evolution requires continual balancing 

of forces in a flank with finite strength as the volcano grows. 

Incorporating finite strength and finite strain into models for 

volcano growth provides considerable insight into the growth of 
Instability of Hawaiian Volcanoes 169

Subaerial

flows

Oceanic crust

Overthrusting

Oceanic crust

Sea level

Sea level

Sea level

Sea level

Growth of 

volcanic apron

Slope failure,

avalanching

?

?

Slumping

Oceanic crust

Volcaniclastic apron

Bench

uplift

Uplift of

slump toe

Seaward 

sliding 

Catastrophic

slope failure

Oceanic crust

Landslide blocks Bench collapse

Surface flows

Surface flows

Older edifice

Older edifice

Older edifice

Older edifice

Fragmental

basalt

Seaward 

sliding 

Rift zone

Rift zone

Rift zone

Rift zone

**

* *

Basal earthquake

Basal earthquakes

Edifice cools;

spreading rate declines

Stage 1

Stage 2

Stage 3

Stage 4

Figure 12. Cross-sectional diagrams showing proposed stages of island growth and degradation, based on integration of 

morphology, structural geology, and geophysical observations of Hawaiian volcanoes of different ages.
170 Characteristics of Hawaiian Volcanoes

Hawaiian volcanoes and into the mechanisms by which stress 

may be stored within volcano flanks as they grow. Despite their 

sophistication, however, these broad-scale models provide few 

constraints on the mechanisms that suddenly trigger rapid slip 

of the flank seaward along its décollement and are believed to 

promote flank collapse.

Instability, as evidenced by the occurrences of rapid slip 

and the ubiquitous geologic deposits indicating sector collapse, 

results from rapid changes in an evolving force imbalance 

driving and resisting flank motion. These forces develop in 

response to processes on temporal and spatial scales that are 

challenging to measure and constrain, particularly within and 

along a deeply buried décollement. Processes that likely occur 

within the décollement layer, most notably coupling of induced 

pore-pressure and dilatancy, are capable of inducing rapid and 

catastrophic changes in the resistance to flank motion after 

decades of nearly steady or episodic creep. The mechanics 

of pore-pressure/dilatancy coupling provide an explanation 

for slow slip events, as well, since décollement models using 

known properties for saturated mixtures of turbidites and clays 

200 to 300 m thick give timescales for slip events tantalizingly 

close to observed timing of sequences of events on the south 

flank of Kīlauea. Yet our understanding of these processes is 

woefully incomplete. Given our inability to precisely scale the 

forces affecting flank instability, the knowledge we have gained 

through recent studies only allows us to qualitatively make 

an educated guess as to the frequency and magnitude of flank 

instability during the growth of Hawaiian volcanoes.

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Global Positioning System receiver atop Pu‘ukapukapu on Kīlauea’s south flank. The cliff in the background—a normal fault that is part of 

the Hilina Pali Fault System—is an indicator of the flank’s instability. GPS is one method used by the Hawaiian Volcano Observatory for 

measuring flank motion. USGS photograph by M.P. Poland, April 16, 2008.
T. Smith operates a transit at Kīlauea on the rim of Halema‘uma‘u Crater to measure lava levels and the locations of features within the 

lava lake. USGS photograph by T.A. Jaggar, Jr., November 3, 1916. 
1

U.S. Geological Survey.

2

University of Wisconsin.

*

now at U.S. Geological Survey.

Abstract

The characteristics of magma supply, storage, and transport 

are among the most critical parameters governing volcanic 

activity, yet they remain largely unconstrained because all three 

processes are hidden beneath the surface. Hawaiian volcanoes, 

particularly Kīlauea and Mauna Loa, offer excellent prospects 

for studying subsurface magmatic processes, owing to their 

accessibility and frequent eruptive and intrusive activity. 

In addition, the Hawaiian Volcano Observatory, founded in 

1912, maintains long records of geological, geophysical, and 

geochemical data. As a result, Hawaiian volcanoes have served 

as both a model for basaltic volcanism in general and a starting 

point for many studies of volcanic processes. 

Magma supply to Hawaiian volcanoes has varied over 

millions of years but is presently at a high level. Supply to 

Kīlauea’s shallow magmatic system averages about 0.1 km3/yr 

and fluctuates on timescales of months to years due to changes 

in pressure within the summit reservoir system, as well as in 

the volume of melt supplied by the source hot spot. Magma 

plumbing systems beneath Kīlauea and Mauna Loa are complex 

and are best constrained at Kīlauea. Multiple regions of magma 

storage characterize Kīlauea’s summit, and two pairs of rift 

zones, one providing a shallow magma pathway and the other 

forming a structural boundary within the volcano, radiate from 

the summit to carry magma to intrusion/eruption sites located 

nearby or tens of kilometers from the caldera. Whether or not 

magma is present within the deep rift zone, which extends 

beneath the structural rift zones at ~3-km depth to the base of 

the volcano at ~9-km depth, remains an open question, but we 

suggest that most magma entering Kīlauea must pass through the 

summit reservoir system before entering the rift zones. Mauna 

Loa’s summit magma storage system includes at least two 

interconnected reservoirs, with one centered beneath the south 

margin of the caldera and the other elongated along the axis of 

the caldera. Transport of magma within shield-stage Hawaiian 

volcanoes occurs through dikes that can evolve into longlived pipe-like pathways. The ratio of eruptive to noneruptive 

dikes is large in Hawai‘i, compared to other basaltic volcanoes 

(in Iceland, for example), because Hawaiian dikes tend to be 

intruded with high driving pressures. Passive dike intrusions 

also occur, motivated at Kīlauea by rift opening in response to 

seaward slip of the volcano’s south flank.

Introduction

Perhaps the most fundamental, yet least understood, aspects 

of volcanic activity involve the supply, storage, and transport 

of magma. Is magma supplied to a volcano episodically or 

continuously (and if the latter, what is the rate of supply)? What 

determines the location and geometry of subsurface magma 

plumbing? How does magma move between source reservoirs 

and eruptive vents? Despite significant advances in volcanology, 

especially since the 1950s, these questions continue to motivate 

much of the research into volcanic and magmatic processes.

Magma supply to subsurface reservoirs is possibly the 

most important control on eruptive and intrusive activity at 

a volcano (Wadge, 1982; Dvorak and Dzurisin, 1993). In 

many cases, supply is episodic, and the arrival of magma 

batches is frequently invoked as a trigger for eruption. For 

instance, an intrusion of fresh basalt into a dacite magma body 

is interpreted as the trigger for the 1991 eruption of Mount 

Pinatubo, Philippines (Pallister and others, 1992). Similarly, 

repeated episodes of magma intrusion and accumulation 

beneath Eyjafjallajökull, Iceland, over at least 18 years 

culminated in an eruption in 2010 (Sigmundsson and others, 

2010). Episodes of magma accumulation are not always 

associated with eruption, however, as exemplified by the 

sudden onset and gradual cessation of inflation, inferred to be 

the result of magma accumulation, near South Sister, Oregon 

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 5

Magma Supply, Storage, and Transport at Shield-Stage 

Hawaiian Volcanoes

By Michael P. Poland1, Asta Miklius1, and Emily K. Montgomery-Brown2*
180 Characteristics of Hawaiian Volcanoes

(Dzurisin and others, 2009). At other volcanoes, like Kīlauea, 

magma supply is inferred to be nearly continuous over 

decadal time scales, with the supply rate roughly proportional 

to the eruption rate (for example, Swanson, 1972; Dvorak and 

Dzurisin, 1993; Cayol and others, 2000).

Magma supplied to a volcano is typically stored in 

reservoirs below the surface prior to intrusion or eruption. Such 

reservoirs occupy a wide depth range. For example, seismic 

and geodetic data place New Mexico’s Socorro magma body 

at a depth of ~20 km (Reilinger and others, 1980; Fialko 

and Simons, 2001), but a zone of magma accumulation at 

Kīlauea may be only 1–2 km beneath the surface (Poland and 

others, 2009; Anderson and others, in press). Reservoir depth 

is controlled by (1) density, as magma ascent stalls where 

magma and host rock density are equal (for example, Ryan, 

1987; Pinel and Jaupart, 2000; Burov and others, 2003), and 

(2) mechanical resistance, where stress barriers due to the 

thermal or mechanical properties of the host rock, relative to 

the magma, inhibit magma flow (for example, Gudmundsson, 

1987; Martel, 2000; Burov and others, 2003). Magma reservoir 

volumes also vary widely, from hundreds of cubic kilometers, 

such as those that fed silicic ash-flow eruptions, like that from 

the Yellowstone Caldera about 0.64 Ma (Christiansen, 2001), 

to about 1 cubic kilometer for the shallow magma reservoir at 

Kīlauea (Poland and others, 2009).

From storage areas, magma is transported toward the 

surface via porous flow through partially molten rock, diapiric 

ascent, or dike and sill intrusion (Rubin, 1995). Dikes and 

sills are perhaps the most common means of magma transport, 

as suggested by their abundance in eroded volcanoes (for 

example, Gudmundsson, 1983; Walker, 1987), as well as 

monitoring data from active volcanoes (for example, Pollard 

and others, 1983). Dike propagation is triggered by both 

source overpressure (for example, Rubin, 1990) and tectonic 

extension (Bursik and Sieh, 1989) and can be either vertical 

or lateral, possibly resulting in eruptions far from the source 

magma reservoir (for example, Sigurdsson and Sparks, 1978). 

Despite research into the mechanics of magma transport, 

important problems remain—for instance, distinguishing 

between intrusions that will erupt versus those that will not, 

which is especially important during volcanic crises (for 

example, Geshi and others, 2010).

Kīlauea and Mauna Loa volcanoes, in Hawai‘i, offer an 

outstanding opportunity to explore magma supply, plumbing, 

and transport, owing to their accessibility, frequent eruptive 

activity, and excellent long-term monitoring record made 

possible by over 100 years of study by the Hawaiian Volcano 

Observatory (HVO) and collaborators. This chapter synthesizes 

past research and new insights from recent volcanic activity 

at Kīlauea and Mauna Loa to develop refined models for 

magma supply, magma storage, and magma transport at shieldstage Hawaiian volcanoes. We discuss those three topics, 

in turn, beginning each discussion with a review of current 

understanding, based on studies of volcanoes around the world, 

before focusing on insights provided by Hawaiian volcanoes. 

Our goal is to derive new knowledge of magmatic processes 

in Hawai‘i that may serve as an example for other volcanoes, 

following the well-established tradition that models based on 

Hawaiian volcanism provide a basis for understanding volcanic 

activity on the Earth and on other planets.

A comprehensive review of even just one aspect of 

Hawaiian volcanism (surface deformation or seismology, for 

example) is impossible, owing to the vast literature on such 

topics. Instead, we seek to mark HVO’s 2012 centennial and 

beyond by (1) building on previous studies (the examination 

of supply, storage, and eruption by Decker [1987] from 

HVO’s 75th-anniversary volume, for example) and (2) 

by incorporating insights from new datasets (including 

Interferometric Synthetic Aperture Radar [InSAR] and the 

Global Positioning System [GPS]) and from recent volcanic 

activity to develop our conceptual models. A more historically 

comprehensive analysis of the eruptive and intrusive history of 

Kīlauea, with special emphasis on magma supply, storage, and 

transport, is provided by Wright and Klein (2014).

Magma Supply

Magma supply from a source region to the base of a 

volcano can be estimated only indirectly. Over long time 

periods, magma supply can be approximated by modeling 

the volume of melt needed to sustain a measured heat flow at 

the surface. For example, Bacon (1982) calculated a basaltic 

magma supply rate of 570 km3/m.y. for the Coso volcanic 

field, California, on the basis of heat flow measurements. 

Similarly, Guffanti and others (1996) used a petrologic 

model of basalt-driven volcanism to determine the basalt 

influx per unit area necessary to account for erupted volumes 

and heat flow at two volcanic centers in northern California: 

the Caribou volcanic field (minimum basalt influx calculated 

to be 0.3 km3 per km2 per m.y.) and the Lassen volcanic 

center (minimum basalt influx of 1.6 km3 per km2 per m.y.). 

The same principle can also be applied over shorter time 

scales. For instance, Francis and others (1993) inferred 

heat loss from the Halema‘uma‘u lava lake that occupied 

Kīlauea’s summit caldera before 1924 and calculated that 

the magma supply was much higher than the eruption rate, 

implying that the volcano grew endogenously through 

subsurface intrusion and cumulate formation. 

Calculations of contemporary rates of magma supply 

have been made at only a few volcanoes on Earth and are 

most often accomplished using eruption rate, gas emission, 

and deformation data. For example, Wadge (1982) calculated 

the supply rates to several volcanoes on the basis of eruption 

volumes over time. Burt and others (1994) and Wadge and 

Burt (2011) inferred a 2.5-fold increase in magma supply to 

Nyamulagira, Democratic Republic of the Congo, starting 

in 1977–80, based on an increase in erupted volumes over 

time. Lowenstern and Hurwitz (2008) used CO2 degassing to 

infer an influx of 0.3 km3/yr of mantle-derived basalt beneath 

Yellowstone Caldera, which is consistent with measured heat 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 181

flow. Allard (1997) and Allard and others (2006) determined 

that about 3–4 times more magma degassed than erupted at 

Mount Etna during 1975–2005, which suggests a supply rate 

to the shallow volcanic system of less than 0.1 km3/yr.

Owing to infrequent eruptions at many volcanoes and the 

challenge in collecting comprehensive gas emission datasets 

in general, magma supply rates are commonly estimated 

from deformation data. At Okmok volcano, Alaska, Lu and 

others (2005) and Fournier and others (2009) inferred variable 

supply to the volcano’s shallow magma reservoir on the basis 

of fluctuations in inflation rate prior to eruptions in 1997 

and 2008. Similarly, deformation measurements at Sierra 

Negra volcano, in the Galápagos Islands, indicated variable 

rates of subsurface magma accumulation, and supply by 

inference, prior to an eruption in 2005 (Chadwick and others, 

2006). After eruptions at many volcanoes (including Okmok 

and Sierra Negra), reinflation is rapid but gradually wanes, 

suggesting that an initially large pressure difference between a 

deep magma source and the shallow magma storage reservoir 

drives magma ascent. The pressure difference decreases as 

the shallow reservoir fills, causing the supply rate to decrease. 

This concept explains the exponential decay of deformation 

rates at Kīlauea (Dvorak and Okamura, 1987) and several 

other basaltic volcanoes (for example, Lu and others, 2005, 

2010; Fournier and others, 2009; Lu and Dzurisin, 2010) and 

has been used to argue for a “top-down” influence on magma 

supply—in other words, the pressure state of the shallow 

reservoir influences the rate of magma ascent from depth 

(Dvorak and Dzurisin, 1993).

Unfortunately, studies based on deformation data alone 

do not account for magma compressibility, which, in the case 

of strong host rock and the presence of exsolved volatiles 

within the magma, may cause 80 percent or more of the magma 

supplied or withdrawn from a reservoir to be accommodated 

without measurable surface inflation or deflation (Johnson, 

1992; Johnson and others, 2000). Better constraints on magma 

supply are achieved by combining gas emissions, deformation, 

and effusion rate (in the case of ongoing eruptions). Using such 

a combination of data, the magma supply rate to Soufrière Hills 

volcano, Montserrat, was estimated to be about 0.06 km3/yr 

and steady over several years, with much of that volume 

accommodated by compression of magma stored in a crustal 

reservoir (Voight and others, 2010). Physics-based models 

that utilize multidisciplinary datasets also hold promise for 

elucidating magma supply characteristics, as demonstrated by 

Anderson and Segall’s (2011) determination that little or no 

recharge was associated with the 2004–08 eruption of Mount St. 

Helens, Washington. 

Magma Supply to Hawaiian Volcanoes

The Hawaiian-Emperor chain of seamounts and islands 

extends over 6,000 km across the Pacific Ocean (fig. 1),tracing 

a record of volcanic activity over at least the past 70 m.y., with 

the age of volcanism progressively younger to the southeast 

(Clague and Dalrymple, 1987; Tilling and Dvorak, 1993). 

Volcanism is attributed to a mantle melting anomaly, termed 

a “hot spot” (Wilson, 1963). The origin of the hot spot is 

generally assumed to be a plume of high-temperature material 

upwelling from the deep mantle, as originally proposed by 

Morgan (1971), and supported by geochemical evidence of a 

primitive source region (for example, Kurz and others, 1983). 

Paleomagnetic evidence suggests that the plume is not fixed in 

the deep mantle but drifts over time, like during the formation 

of the Emperor Seamounts (Clague and Dalrymple, 1987; 

Tarduno and others, 2003), but plate motion models argue that 

Hawaiian hot spot motion may be relatively small (Gripp and 

Gordon, 2002; Wessel and others, 2006; Wessel and Kroenke, 

2008). The plume hypothesis has been challenged as dogma 

that fails to explain numerous geochemical and geophysical 

datasets (see Hamilton, 2011, and references therein, for 

alternatives to a plume origin), but seismic tomography 

studies indicate a deep mantle source for the melting anomaly, 

providing evidence in favor of the presence of a mantle plume 

(Wolfe and others, 2009, 2011). Regardless of its origin, the 

Hawaiian hot spot is the most productive on Earth, based on 

buoyancy fluxes (for example, Sleep, 1990) and eruption rates 

(Wadge, 1982). Tracking magma supply from the hot spot to 

the crust thus provides important constraints on the production 

of melt over time.

Estimates of the Hawaiian hot spot’s magma supply 

rate—the flux of magma from the source region to the 

surface—over millions of years are based on the volumes and 

ages of the islands and seamounts of the Hawaiian-Emperor 

chain. Early calculations used bathymetric and topographic 

maps to determine the volumes of individual volcanoes (Vogt, 

1972; Bargar and Jackson, 1974). Later, more sophisticated 

studies incorporated seismic, gravity, and crustal loading 

data to quantify mass flux (White, 1993; Van Ark and Lin, 

2004; Vidal and Bonneville, 2004; Robinson and Eakins, 

2006). Regardless of the method used, all long-term magma 

supply calculations have found an increase in the activity of 

the Hawaiian hot spot over the past 30 m.y. after relatively 

constant melt production during the formation of the Emperor 

chain, with the highest rates occurring at present (fig. 2). The 

current high supply may be reflected in the geochemistry 

of erupted lavas. Weis and others (2011) suggested that the 

strong appearance of the “Loa” geochemical composition (see 

Clague and Sherrod, this volume, chap. 3; Helz and others, 

this volume, chap. 6) may coincide with increased magma 

supply over the past 5 m.y. Variations in the supply rate on the 

scale of millions of years have been attributed to changes in 

lithospheric thickness (especially where the hot spot crossed 

fracture zones), temperature, and age, as well as pulsation of 

the hot spot itself (Vogt, 1972; White, 1993; Van Ark and Lin, 

2004; Vidal and Bonneville, 2004). 

Although the magma supply rates calculated by various 

studies differ by several times (fig. 2), they are similar in two 

important respects. First, activity of the Hawaiian hot spot 

(manifested primarily by erupted volume) varies over millions 

of years, with the highest values occurring in recent times 

(Clague and Dalrymple, 1987). Second, a dominant influence 
182 Characteristics of Hawaiian Volcanoes

200° 190° 180° 170° 160° 150°

20°

30°

40°

50°

60°

Emperor chain

Bend

Aleutian arc

Kamchatka

Kīlauea

Mauna Loa

Mauna Kea

Hualālai

Kohala

Hawaiian Islands

Island of Hawaiʻi

85 mm/yr

−10,000

4,000

0

Bathymetry/

topography

0

0

500 1,000 KILOMETERS

250 500 MILES

Hawaiian chain

0 10 20 MILES

0 20 40 KILOMETERS

Figure 1. Bathymetric map 

showing the Hawaiian and Emperor 

seamount chains in the northwest 

Pacific Ocean basin, as well as 

the Hawaiian Islands. Inset shows 

topography of Island of Hawai‘i, with 

the five volcanoes that make up the 

island indicated. Arrow indicates plate 

motion velocity in millimeters per year 

(Simkin and others, 2006).

0.3

0.2

0.1

0

80 70 60 50 40 30 20 10 0

0.4

0.5

0.6

Volume flux of magma,

in cubic kilometers per year

Age of volcanism, in millions of years

Bargar and Jackson, 1974

White, 1993

Vidal and Bonneville, 2004 Bend in HawaiianEmperor chain

Van Ark and Lin, 2004

Vogt, 1972 Figure 2. Plot of the volume flux of the Hawaiian hot spot over 

time, as calculated by various authors. The estimate of Bargar and 

Jackson (1974) is based on the volume of seamounts and islands in 

the Hawaiian-Emperor chain, with the timing established by Robinson 

and Eakins (2006). Vogt (1972) used a similar volume-age relation. 

White (1993) based his volume flux curve on crustal thickening 

determined seismically, coupled with the volume of seamounts and 

volcanoes and their underplated roots, assuming Airy isostasy. Vidal 

and Bonneville (2004) distinguished topography/bathymetry created 

by the Hawaiian Swell (caused by uplift driven by the mantle plume) 

from that caused by volcanism, which is a better measure of the 

magma production rate. Finally, Van Ark and Lin (2004) calculated 

crustal thicknesses from gravity data to compute hot-spot volume 

flux over time. Dashed vertical line is the location of the bend in the 

Hawaiian-Emperor chain of seamounts and islands (see fig. 1).
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 183

on hot-spot activity is the rate of melt production. The volume 

of melt provided to a volcano by the hot spot apparently varies 

over millions of years, and perhaps on even shorter timescales. 

The composition of Mauna Kea basalt accessed through 

drilling, for example, suggests waxing and waning supply 

within the shield-building stage of that volcano (less than 

1 m.y. in duration), instead of a smoother increase and then 

decrease in supply as the volcano grew and was then carried 

away from the hot spot (Stolper and others, 2004; Rhodes 

and others, 2012). The possibility thus exists that shorter-term 

variations in melt production occur in addition to longer-term 

changes documented by studies of the Hawaiian hot-spot track.

Contemporary Estimates

Determinations of the volume of erupted material 

over time at Mauna Loa and Kīlauea provide estimates of 

the minimum contemporary magma supply. Stearns and 

Macdonald (1946, p. 113–114), for example, noted that 

the total volume of erupted products at Kīlauea during 

1823–1945 was approximately 3.2 km3, corresponding to an 

annual rate of about 0.03 km3/yr. Moore (1970) calculated 

a combined average annual eruption rate of 0.05 km3/yr for 

Kīlauea and Mauna Loa during 1823–1969 and noted that the 

volume was less by a factor of 5 than the annual volume of 

subsidence caused by loading. He attributed the discrepancy 

to an unknown volume of magma added by intrusion and 

submarine eruption. Both of these estimates are consistent 

with that of Quane and others (2000), who determined an 

eruption rate of 0.05 km3/yr over the past 350,000 years at 

Kīlauea, based on lava ages from a drill core from the lower 

East Rift Zone (ERZ).

Swanson (1972) made the first estimate of magma supply 

rate to Kīlauea by accounting for volumes of both stored 

magma and erupted lava. His calculation of 0.11 km3/yr (dense 

rock equivalent) during 1952–71 was based on the volumes 

of three effusive eruptions (each of which lasted at least 4.5 

months), during which there was little change in the amount of 

magma stored beneath the summit. In addition, he speculated 

that this supply rate was constant even between eruptions, 

when Kīlauea inflated to accommodate magma influx. He 

interpreted the supply rate to reflect the total amount of 

magma produced by the hot spot, given the lack of Mauna 

Loa eruptions during 1952–71. Swanson’s (1972) supply rate 

eliminates the discrepancy between the volumes of subsidence 

and eruption rate cited by Moore (1970) when simple isostatic 

compensation is assumed.

Subsequently, numerous authors built on Swanson’s 

(1972) work by combining models of surface deformation 

with extrusion volumes to calculate supply rates (table 

1). For example, Dzurisin and others (1984) calculated 

an average supply rate of 0.09 km3/yr during 1956–1983, 

although the period included short-term fluctuations 

associated with changing eruptive activity. Dvorak and 

Dzurisin (1993) found that magma supply between the 1960s 

and 1990s varied over days to years between 0.02 km3/yr and 

0.18 km3/yr, with changes in rate controlled by the pressure 

in the summit magma reservoir. Estimates of the supply rate 

during the sustained Pu‘u ‘Ō‘ō ERZ eruption, which started 

in 1983, fall within this range. Wolfe and others (1987) 

determined a supply rate of 0.12 km3/yr from the first 20 

episodes of the eruption in 1983–84, while Denlinger (1997) 

inferred a supply rate of 0.08 km3/yr based on the extrusion 

rate during a period of zero summit elevation change. 

Heliker and others (2003) calculated an average effusion rate 

of 0.12 km3/yr over 19 years that were dominated by summit 

deflation (implying that nearly all of the magma entering the 

volcano’s summit reservoir system was erupted), although 

Garcia and others (1996) suggested that minor magma 

supply variations sometimes influenced the composition of 

erupted lava during the Pu‘u ‘Ō‘ō activity. 

The average magma supply rate in published reports 

covering the period 1952–2002 (table 1) is 0.1 km3/yr. Wright 

and Klein (2014) presented a comprehensive analysis of magma 

supply since 1790, independent of the studies summarized in 

table 1, and also found that the supply rate since the 1960s was 

relatively steady. Short-term increases or decreases in rate are 

probably caused by fluctuations in pressure within Kīlauea’s 

shallow summit magma reservoir system that promote or 

suppress supply from depth (for instance, Cervelli and Miklius, 

2003). There is some evidence that the supply rate has been 

steady, at least since the 1800s. Francis and others (1993) 

determined that a magma supply rate of about 0.12 km3/yr was 

needed to account for heat loss from the lava lake occupying 

Kīlauea’s summit prior to the 1924 collapse, even though the 

eruption rate—which was up to an order of magnitude lower 

(Francis and others, 1993; Wright and Klein, 2014)—indicated 

that most of the supplied magma was stored, not erupted.

Models of geodetic data from Kīlauea suggest deep 

(~4–10 km) dilation beneath the East and Southwest Rift 

Zones (for example, Dieterich, 1988; Delaney and others, 

1990; Owen and others, 1995, 2000a; Cayol and others, 

2000). If this opening volume is filled by magma, as is 

interpreted from seismic data (Ryan, 1988; Wyss and 

others, 2001), the magma supply rate to Kīlauea could be 

almost twice that estimated from eruption and shallow 

magma storage alone. Cayol and others (2000) suggested 

that the magma supply rate to Kīlauea during 1961–91 was 

relatively steady, at 0.18 km3/yr, based on deformation 

modeling, consistent with the analysis of Wright and Klein 

(2014). Such a high value implies that the CO2 content of 

the primary magma supplied to Kīlauea is ~0.70 percent, 

which agrees with previous estimates, while a supply rate 

of 0.1 km3/yr would require 1.21 percent, which may be 

unrealistic (Gerlach and others, 2002). The higher supply 

rate calculated by Cayol and others (2000) and Wright and 

Klein (2014) does not invalidate the previous estimates 

(summarized in table 1) but, instead, adds a new dimension 

to the discussion—is magma supplied to, and stored within, 

Kīlauea’s deep rift zones? We address this question below 

in the sections entitled “Magma Supply Dynamics Between 

Kīlauea and Mauna Loa” and “Deep Rift Zones.”
184 Characteristics of Hawaiian Volcanoes

Time period Supply (km3/yr) Method Reference

1918–1979 0.08 Ratio between repose times and erupted volumes Klein, 1982

1919–1990 0.09 Effusion rate of several sustained eruptions Dvorak and Dzurisin, 1993

1952–1971 0.11 Effusion rate of three sustained eruptions Swanson, 1972

1956–1983 0.09 Average summit and rift deformation and eruption 

volumes Dzurisin and others, 1984

1959–1990 0.06 Average based on deformation and eruption volumes Dvorak and Dzurisin, 1993

1960–1967 0.02–0.18 Deformation-inferred refilling of summit reservoir Dvorak and Dzurisin, 1993

1961–1970 0.18 Eruption rate and deformation volumes Cayol and others, 2000

1966–1970 0.07 Deformation and eruption volumes Dvorak and others, 1983

1967–1975 0.05–0.18 Deformation and eruption volumes Wright and Klein, 2008

1971–1972 0.08 Deformation and erupted volumes Duffield and others, 1982

1975–1977 0.07–0.16 Microgravity and deformation Dzurisin and others, 1980

1976–1982 0.19 Eruption rate and deformation volumes Cayol and others, 2000

1983–1984 0.12 First 20 episodes of Puʻu ʻŌʻō eruption Wolfe and others, 1987

1983–1991 0.14–0.18 Eruption rate and deformation volumes Cayol and others, 2000

1983–2002 0.12 Puʻu ʻŌʻō eruption volumes Heliker and Mattox, 2003

1983–2002 0.13 SO2 emissions from ERZ Sutton and others, 2003

1991 0.08 Deformation and effusion rates Denlinger, 1997

Table 1. Estimates of the magma supply rate during various times in Kīlauea’s recent history. Shaded estimates include the volume created by 

deep rift opening. Table does not include estimates by Wright and Klein (2014), which span all time periods.
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 185

Only sparse attempts have been made to document magma 

supply to Mauna Loa. This lack of attention is unsurprising, 

given the paucity of eruptive activity since 1950 (compared to 

Kīlauea) and the dearth of deformation data (in time and space). 

Lipman (1995) calculated a magma supply rate for Mauna Loa 

of 0.028 km3/yr over the past 4 k.y. by summing the eruption 

rate over that time period (0.02 km3/yr)—an estimate of the 

volume added by shallow dike emplacement in the rift zones 

(0.004 km3/yr) and an estimate of magma accumulation at depth 

(0.004 km3/yr)—and by assuming that cumulative summit 

inflation over the time period was negligible (based on analogy 

with Kīlauea). Such a supply rate, too small to account for the 

size attained by Mauna Loa (80,000 km3) over its lifetime, led 

Lipman (1995) to speculate that magma supply rates must have 

been higher in the past—perhaps similar to the 0.1 km3/yr rate 

that has characterized Kīlauea over the last several decades—

and decayed as Mauna Loa was carried away from the hot spot 

locus by Pacific Plate motion. 

155°30' 155°

19°30'

155°20' 155°10' 154°50'

19°20'

19°10'

UWEV

AHUP

WAOP

NUPM

HVO7

SWRZ

ERZ

Summit

caldera

0

0

5 10 KILOMETERS

2.5 5 MILES

NW of caldera

KOSM

May 24, 2007

M4.7

May 24, 2007, M4.1

Mauna Ulu

Makaopuhi

Crater Puʻu ʻŌʻō

Kupaianaha

Nāpau

Crater Summit

eruptive

vent

June 17–19, 2007, dikes

June 19, 2007,

eruptive vent July 21, 2007,

eruptive fissure

July 21, 2007 – 

March 5, 2011

eruptive vent

Halema Kīlauea Iki ʻumaʻu

Crater

Keanakākoʻi

Crater

Island of Hawai‘i

Mauna

Loa Kīlauea

Great Crack

Maunaiki

2003–07 Increase in Magma Supply to Kīlauea

Deformation, seismic, gas, and geologic data indicate 

an increase in the rate of magma supply from the mantle 

to Kīlauea during 2003–07. In late 2003, deformation 

at Kīlauea’s summit (fig. 3) transitioned from long-term 

deflation to inflation, as indicated by the cessation of 

subsidence at GPS station AHUP and increasing distance 

between GPS stations AHUP and UWEV (fig. 4A). SO2

emission rates from Kīlauea’s ERZ remained nearly constant 

at about 1,370 metric tons per day (t/d) during this period 

(fig. 4B), however, indicating no change in the amount of 

magma being transported from the summit to the ERZ—

about 0.13 km3/yr (Sutton and others, 2003). The inflation 

was therefore not caused by a decrease in effusion rate and 

subsequent backup in Kīlauea’s magma plumbing system, 

as has occurred at other times during the Pu‘u ‘Ō‘ō eruption 

(Kauahikaua and others, 1996; Miklius and Cervelli, 2003). 

Figure 3. Shaded 

topographic map of Kīlauea 

Volcano. Labeled red dots 

give the locations and 

site names of continuous 

Global Positioning System 

stations mentioned in the 

text and other figures. 

ERZ, East Rift Zone; 

SWRZ, Southwest Rift 

Zone. Bottom image shows 

detail of Kīlauea’s summit 

and ERZ, with major 

features discussed in the 

text labeled. Red diamonds 

give the locations of 

eruptive vents discussed 

in the text. White stars are 

earthquakes that occurred 

on May 24, 2007 (Wauthier 

and others, 2013). 

Dashed red lines give the 

geometry of a pair of en 

echelon dike segments 

that were emplaced during 

June 17–19, 2007 (from 

Montgomery-Brown and 

others, 2010). Solid red 

line shows location of the 

July 21, 2007, eruptive 

fissure.
186 Characteristics of Hawaiian Volcanoes

2000 2002 2004 2006 2008 2010

0

1,000

2,000

3,000

4,000

5,000

6,000

7,000

8,000

9,000 B

SO2 emissions,

in metric tons per day

0

10,000

20,000

30,000

40,000

1995-09-201998-10-201999-05-06

E

CO2 emissions,

in metric tons per day

0

0.2

0.4

0.6

0.8

1.0

1.2

1.4 A

AHUP vertical

KOSM vertical

Displacement, in meters

AHUP-UWEV distance change

0

500

1,000

1,500

2,000 C

Upper SWRZ

NW of caldera

Number of

earthquakes

Upper ERZ

6.5

7.0

7.5

8.0

9.0

8.5

D

Bulk lava MgO content,

in weight percent

June 17-19 

intrusion/eruption

1,130

1,135

1,140

1,145

1,150

1,155

1,160

1,165

glass geothermometry,

Lava temperature from

in degrees Celsius

Glass temperature

MgO content

Year

Figure 4. Geophysical and 

geochemical time-series data from 

Kīlauea Volcano during 2000–10. 

Dotted red vertical line indicates 

June 17–19, 2007, East Rift Zone 

(ERZ) intrusion and eruption. A, 

Global Positioning System data 

showing distance change between 

the AHUP and UWEV stations 

(black; positive change indicates 

extension), vertical change at 

station AHUP (blue; positive change 

indicates uplift), and vertical change 

at station KOSM (red). Station 

locations are given in figure 3. 

Uplift and extension are indicative 

of surface inflation. B, ERZ SO2

emissions. C, Cumulative numbers 

of located earthquakes in the upper 

ERZ (black), upper Southwest Rift 

Zone (SWRZ; blue), and northwest 

of the caldera (red) (see fig. 3 for 

locations). D, MgO weight percent 

(left axis) and eruption temperatures 

from glass geothermometry (right 

axis; Helz and Thornber, 1987), 

measured in lavas erupted from the 

ERZ. E, Summit CO2 emissions. 

The first three measurements were 

made before 2000 but are shown 

here to indicate the consistency of 

pre-2004 emission rates.
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 187

2000 2002 2004 2006 2008 2010

0.15

0.1

0.05

0.0

-0.05

-0.10

-0.15

WAOP

A

B

NUPM

Vertical deformation, in meters

1975 1980 1985 1990 1995 2000 2005 2010

0.2

0.15

0.1

0.0

-0.05

-0.10

0.05

0.3

0.25

Leveling

GPS

HVO7

Year

Since there was no precursory deflation event that would 

have created a pressure imbalance and a “top-down” change 

in magma supply, the most likely explanation for the 

inflation is increased magma flux from the hot spot. In early 

2005, a pulse of magma from the summit to the ERZ resulted 

in both temporarily heightened rift SO2 emissions (fig. 4B), 

increased lava effusion from the Pu‘u ‘Ō‘ō eruptive vent, and 

a several-week period of summit deflation as magma drained 

from the summit to feed the East Rift Zone effusive surge 

(fig. 4A). By mid-2005, summit inflation had resumed (fig. 

4A), and rift SO2 emissions had increased to a new average 

rate of ~1,900 t/d (fig. 4B), implying that about 0.18 km3/yr 

of magma was then being delivered from the summit to the 

ERZ.

The summit continued to inflate, following the 2005 

surge in ERZ effusion, indicating that the ERZ conduit could 

not accommodate the elevated amount of magma being 

supplied to Kīlauea and that magma was accumulating beneath 

the summit. The rate of summit inflation increased rapidly in 

early 2006 (fig. 4A), and by the middle of that year, uplift was 

also occurring beneath the Southwest Rift Zone (which had 

been subsiding since 1982), suggesting magma accumulation 

beneath that part of the volcano, as well (fig. 4A, station 

KOSM). Heightened stress in the summit region caused by the 

inflation was indicated by swarms of short-period earthquakes 

in the upper parts of the East and Southwest Rift Zones, as 

well as northwest of the caldera (fig. 4C).

Modeled volume changes suggest that at least 0.01 km3

of magma accumulated beneath the summit and Southwest 

Rift Zone in 2006 (Poland and others, 2012) which, when 

combined with ERZ effusion, implies a minimum supply rate 

of 0.19 km3/yr to Kīlauea’s shallow magma system during 

that year. In addition, GPS data reveal that the normal pattern 

of ERZ subsidence stopped or switched to uplift at several 

places during 2003–07 (fig. 5), indicating magma storage in 

the ERZ and that the magma supply rate to Kīlauea’s shallow 

magmatic system must have been even higher. No change 

in the rate of deep opening of the ERZ was detected during 

2003–07, as suggested by steady motion of Kīlauea’s south 

flank during this time period (fig. 6), so the increased magma 

supply apparently affected only Kīlauea’s shallow magma 

plumbing system. 

Petrologic changes to lava erupted from the ERZ also 

reflect the magma supply increase. Lava erupted from Pu‘u 

‘Ō‘ō is a hybrid between a high-MgO, high-temperature 

magma supplied from the hot spot and a partially degassed, 

lower-MgO, lower-temperature resident magma that has 

partially crystallized due to storage at shallow levels within 

the volcano (Thornber, 2003; Thornber and others, 2003). 

MgO values and lava temperature had been declining since 

2000 (fig. 4D), reflecting an increase in the proportion 

of lower-temperature magma being erupted. In addition, 

the mineralogical composition of the lava included a 

nonequilibrium assemblage of high-temperature olivine plus 

low-temperature clinopyroxene, olivine, and plagioclase 

(Thornber and others, 2010). MgO content stabilized at 

Figure 5. Vertical deformation at sites along Kīlauea’s East Rift Zone. 

Station locations are given in figure 3. Dashed lines indicate June 17–

19, 2007, intrusion and eruption. A, Plot of elevation change at Global 

Positioning System (GPS) stations NUPM and WAOP. Both sites show 

a transition from subsidence to uplift in late 2003 at the same time that 

summit inflation commenced. No data are available from WAOP after 

August 24, 2007, because the station was overrun by advancing lava. 

B, Time series of vertical elevation change at site HVO7 from leveling 

(red circles) and campaign GPS (blue circles). The trend of long-term 

subsidence at the site was interrupted during 2003–07, coincident with 

the period of increased magma supply. The station is located 6.5 km 

from a geothermal power plant, which is distant enough to be free from 

deformation associated with geothermal production.
188 Characteristics of Hawaiian Volcanoes

-0.2

-0.1

0

0.1

0.2

East

-0.2

-0.1

0

0.1

0.2

North

2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010

-0.1

-0.05

0

0.05

0.1 Vertical

KAEP

Deformation, in meters

Year

Figure 6. Plots showing east (top), north (middle), and vertical (bottom) components of deformation at Global Positioning System station 

KAEP, located on the south flank of Kīlauea Volcano (location given in inset). Linear deformation suggests no change in the opening rate of the 

deep rift zone during 2000–10. Small offsets in the time series are related to aseismic slip events on the south flank (Brooks and others, 2006; 

Montgomery-Brown and others, 2009).

~7.0 percent in early 2004 and, along with lava temperature, 

remained steady until mid-2007 (fig. 4D)—a sign of shallow 

mixing between hotter recharge magma and shallow resident 

magma, suggesting that cooler parts of the magma storage 

and transport pathways in Kīlauea’s summit and ERZ were 

being stirred by an influx of new magma. Lava erupted on 

June 19, 2007, near Makaopuhi Crater (see section entitled 

“Consequences of Changes in Supply”), was the hottest 

and most primitive since 1998, representing the high-MgO, 

high-temperature magma component. After that eruption, 

MgO content and temperature increased steadily in lava 

erupted from Pu‘u ‘Ō‘ō until mid-2008 (fig. 4D), suggesting 

that heightened magma supply introduced more of the hightemperature magmatic component to Kīlauea.

CO2 emissions from Kīlauea are an independent indication 

of deep magma supply, because gas begins to exsolve from 

ascending magma at a depth of ~30 km and is therefore mostly 

insensitive to shallow processes (Gerlach, 1986; Gerlach 

and others, 2002). Before 2005, sporadic CO2 emission 

measurements yielded constant values of about 8,000 t/d, but a 

measurement in July 2004 was 18,200±2,500 t/d (fig. 4E), and 

more frequent measurements, starting in 2005, averaged about 

20,000 t/d for several months. The increase in CO2 emissions 

began sometime between mid-2003 and mid-2004—the same 

period of time that summit inflation commenced. Interestingly, 

the CO2 emission rate did not track summit inflation but, 

instead, reached a maximum more than a year before the 2005 

increase in lava effusion and three years prior to the highest 

rates of summit inflation, emphasizing the importance of CO2

monitoring for forecasting changes in volcano behavior.

By the end of 2008, geological, geophysical, and 

geochemical data indicated a return to pre-2003 rates of magma 

supply. CO2 emission rates, which began to decline in 2006, 

reached pre-2003 levels (fig. 4E), deflation dominated the 

summit after formation of a new ERZ vent on July 21, 2007 

(fig. 4A), and SO2 emissions from the ERZ had decayed towards 

pre-2003 rates by late 2008 (fig. 4B), indicating a decrease in 

the quantity of magma being transported from the summit to the 

ERZ eruption site. Summit seismicity returned to background 

levels following the intrusion and eruption of June 2007 

(fig. 4C), although tremor levels increased in the months prior to 

the start of the 2008 summit eruption (Poland and others, 2009). 

Finally, the MgO content of lavas erupted from the ERZ began 

to decline in mid-2008 (fig. 4D), indicating a decrease in the 

proportion of newly supplied magma.

The long-term Pu‘u ‘Ō‘ō eruption on Kīlauea’s ERZ, 

which started in 1983 and continued unabated through 

2003–07, suggests that the increased supply was not caused 

by pressure fluctuations in the shallow summit magma system, 

as has been proposed for previous changes in supply rate 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 189

(Dvorak and Dzurisin, 1993). Instead, the increase must have 

been driven by a greater flux of magma from the mantle, 

which is also supported by heightened CO2 emissions. This 

documentation of a short-term change in hot-spot activity 

demonstrates that supply of magma from the mantle to 

Hawaiian volcanoes can vary on time scales of years.

Consequences of Changes in Supply

Variations in magma supply to Kīlauea and Mauna 

Loa have been associated with major changes in the style 

of activity at the two volcanoes, as well as compositional 

variations in erupted lavas (Garcia and others, 1996, 2003). 

For example, Dzurisin and others (1984) calculated high 

supply rates to Kīlauea in the years before and during the large 

East Rift Zone eruptions that occurred in 1960 (Kapoho) and 

1969–74 (Mauna Ulu). Dvorak and Dzurisin (1993) proposed 

that magma supply rates to Kīlauea were low during the 1800s 

and early 1900s but higher starting in the 1950s, implying 

that summit eruptive activity was favored during periods of 

low supply, and that ERZ activity and sustained rift eruptions 

were a consequence of high supply—similar to the model of 

Wright and Klein (2014) for magma supply during the 20th 

and early 21st centuries. Geochemical variations in lava flows 

erupted during the early 19th to middle 20th centuries suggest 

decreased partial melting of the source region, also indicative 

of a lower magma supply rate (Pietruszka and Garcia, 1999a). 

In contrast, modeled heat loss from the summit lava lake prior 

to 1924 argues for higher supply (Francis and others, 1993), 

exemplifying the ambiguity in magma supply estimates.

Changes in volcanic activity as a consequence of 

increased magma supply were best documented following 

the 2003–07 magma supply surge to Kīlauea, described 

above and by Poland and others (2012). Rapid summit 

inflation due to the supply surge caused an increase in 

stress on the caldera-bounding normal faults and resulted 

in a pair of M4+ earthquakes on the southeast margin of 

the caldera on May 24, 2007 (fig. 3; Wauthier and others, 

2013). Continued accumulation of pressure at the summit 

drove an intrusion from the summit into the ERZ during 

June 17–19, 2007, which resulted in a small eruption of 

high-MgO lava (fig. 4D) just north of Makaopuhi Crater 

(fig. 3) on June 19 (Poland and others, 2008; Fee and others, 

2011). Increased compressive stress on Kīlauea’s south 

flank due to ERZ opening during June 17–19 apparently 

triggered an aseismic slip event on the basal detachment 

fault that underlies Kīlauea shortly after the onset of the 

intrusion (Brooks and others, 2008; Montgomery-Brown 

and others, 2010, 2011). Eruptive activity resumed in Pu‘u 

‘Ō‘ō crater on July 1, but pressure buildup beneath the vent 

caused a fissure eruption on the east flank of the cone on 

July 21, leading to the formation of a new long-term eruptive 

vent about 2 km downrift of Pu‘u ‘Ō‘ō (fig. 3; Poland and 

others, 2008). The formation of an eruptive vent at Kīlauea’s 

summit (fig. 3) on March 19, 2008 (Wilson and others, 2008; 

Houghton and others, 2011; Patrick and others, 2011), has 

also been attributed indirectly to the magma supply increase. 

Poland and others (2009) suggested that decompression of 

the inflated summit due to formation of the July 21, 2007, 

East Rift Zone eruptive vent caused exsolution of volatiles, 

which ascended through existing fractures and reached 

the surface along the southeast margin of Halema‘uma‘u 

Crater. Increasing volatile pressure dilated the pathways and 

caused the March 19 explosion that formed the new eruptive 

vent, allowing magma to passively rise toward the surface 

without breaking rock and without causing earthquakes or 

deformation. 

Variable magma supply has also been cited as a cause 

for changes in eruptive activity at Mauna Loa and even the 

structural evolution of that volcano. The eruption rate between 

1843 (the first recorded eruption of Mauna Loa) and 1877 

was more than twice that of the post-1877 rate (Lockwood 

and Lipman, 1987; Lipman, 1995). Decreasing incompatibleelement abundances in erupted lavas during 1843–77, coupled 

with the heightened eruption rate, might be a sign of increased 

magma supply (Rhodes and Hart, 1995). Riker and others 

(2009) proposed that an eruption of MgO-rich lava in 1859 from 

a radial vent on the northwest flank of the volcano (outside a rift 

zone) was a consequence of high magma supply based on the 

primitive composition and vent location.

Deformation and gravity data spanning Mauna Loa’s 1984 

eruption indicate that the volume of extruded lava was much 

greater than the volume removed from the summit magma 

reservoir. This discrepancy, coupled with the observation that 

CO2 emissions were at their highest during the later part of the 

eruption, prompted Johnson (1995a) to suggest that magma 

supply to Mauna Loa is episodic, with most influx occurring 

concurrently with eruptions—akin to the “top-down” model 

of Dvorak and Dzurisin (1993) for Kīlauea. Unfortunately, 

testing these hypotheses will be difficult without more frequent 

eruptive activity at Mauna Loa than characterized the latter half 

of the 20th century and the start of the 21st century.

Magma Supply Dynamics Between Kīlauea and 

Mauna Loa

The possibility of a connection between the magmatic 

systems of Kīlauea and Mauna Loa has been a source of 

debate since the volcanoes were first described scientifically 

in the mid-1800s (see discussion in Stearns and Macdonald, 

1946, p. 132–135). In general, petrologic differences in lava 

erupted from the two volcanoes provide convincing evidence 

that their magma sources are geochemically distinct and their 

magma plumbing systems are independent (for example, 

Wright, 1971; Rhodes and others, 1989; Frey and Rhodes, 

1993). Mauna Loa-like magmas have erupted from Kīlauea 

(Rhodes and others, 1989), however, and lavas erupted at 

Kīlauea during 1998–2003 showed signs of an increasing 

proportion of a Mauna Loa component (Marske and others, 

2008). In addition, small-scale compositional heterogeneities 

have been seen to affect both volcanoes at roughly the 
190 Characteristics of Hawaiian Volcanoes

same time, suggesting that the source regions for the two 

volcanoes are not widely separated (Marske and others, 

2007). In fact, seismic evidence supports the possibility of an 

interconnected melt zone between the two volcanoes below 

about 30-km depth (Ellsworth and Koyanagi, 1977). 

Observations of eruptive activity at Kīlauea and Mauna 

Loa highlight correlations that argue for some sort of shared 

magma supply, with several authors noting an inverse relation 

between their eruptions (for example, Moore, 1970; Klein, 

1982). During 1934–52, for instance, Kīlauea was dormant 

but Mauna Loa erupted six times, while during 1952–2014, 

Mauna Loa erupted only twice and Kīlauea was frequently 

(and often continuously) active. Klein (1982) found this anticorrelation to be statistically significant and suggested that the 

volcanoes were competing for a common supply of magma. 

Kīlauea and Mauna Loa have also behaved sympathetically. In 

May 2002, an effusive surge from Kīlauea’s ERZ occurred at 

the same time as the onset of inflation at Mauna Loa. Miklius 

and Cervelli (2003) suggested that input of magma into 

Mauna Loa increased the pressure in Kīlauea’s magma system, 

triggering the ERZ effusive episode. 

A more direct case of sympathetic behavior between the 

two volcanoes is suggested by activity spanning 2002–07. 

Mauna Loa began to inflate in 2002 after nearly a decade 

of deflation (Miklius and Cervelli, 2003; Amelung and 

others, 2007). The inflation rate increased in late 2004 and 

was accompanied by a swarm of thousands of long-period 

earthquakes at >30-km depth (Okubo and Wolfe, 2008), but 

inflation waned and ceased by the end of 2009 (fig. 7). The 

fact that Mauna Loa’s inflation occurred at approximately the 

same time as the 2003–07 surge in magma supply to Kīlauea 

is an unlikely coincidence and suggests the possibility that 

an increase in magma supplied from the mantle affected both 

volcanoes. Dzurisin and others (1984) proposed that a similar 

relation existed in the late 1970s, when their model suggested 

a period of increasing magma supply to Kīlauea during an 

episode of inflation at Mauna Loa. Similarly, a simultaneous 

change in the composition of erupted lavas at both volcanoes 

occurred during 250–1400 C.E. (Marske and others, 2007), 

indicating that Kīlauea and Mauna Loa have previously 

been affected by mantle source processes at the same time. 

Gonnermann and others (2012) argued that the two volcanoes 

may be dynamically linked through an asthenospheric porous 

melt zone located tens of kilometers deep. In their model, 

shallow magma storage at each volcano is distinct, and the 

magma feeding systems tap different parts of the mantle 

source (thereby explaining the overall petrologic differences 

in erupted lavas). Changes in magma pressure, however, can 

be transmitted between crustal storage reservoirs through the 

asthenospheric melt zone on time scales of less than a year 

without requiring direct melt transport between volcanoes. 

The model therefore provides a mechanism by which Kīlauea 

and Mauna Loa display complementary modes of behavior 

without requiring a shallow connection.

An asthenospheric melt zone that links Mauna Loa and 

Kīlauea might also explain observations of CO2 emissions at 

the two volcanoes. As discussed above, Gerlach and others 

(2002) favored the 0.18 km3/yr supply rate of Cayol and others (2000), because lower rates imply that the CO2

 content 

of the primary magma supplied to Kīlauea is unrealistically 

high (1.21 percent versus 0.70 percent). This higher supply 

rate is based on the assumption that nearly half of the magma 

fed to Kīlauea is stored in the deep ERZ, which is shallow 

enough that CO2 would degas, but deeper than the exsolution 

pressures of other volatile species. If no magma is stored in 

this region (see “Deep Rift Zones” section below), however, 

how can the lower supply rate be reconciled with the favored 

CO2 content of the magma?

A possible explanation is that CO2 from most magma 

supplied by the hot spot degasses through Kīlauea’s summit. 

Deep seismicity (primarily tremor and long-period earthquakes) 

at ~40-km depth offshore of Kīlauea’s south flank has been 

interpreted as the magma source that feeds the active volcanoes 

of the Island of Hawai‘i (Aki and Koyanagi, 1981; Wright 

and Klein, 2006). Wright and Klein (2006) further proposed 

that the deep feeder is linked to Kīlauea via a subhorizontal 

zone of magma transport at about 30-km depth, similar to the 

asthenospheric melt zone envisioned by Gonnermann and 

others (2012). CO2 starts to exsolve at about this depth (Gerlach 

and others, 2002), so CO2 bubbles might ascend along the path 

closest to the deep conduit from the source—the nearby conduit 

to Kīlauea’s summit—regardless of the ultimate destination 

of the magma, be it Kīlauea, Mauna Loa, or possibly Lō‘ihi. 

Such a model (fig. 8) has the potential to explain not only the 

CO2 content discrepancy raised by Gerlach and others (2002) 

at Kīlauea, but also why so little CO2 has been emitted from 

Mauna Loa despite periods of magma accumulation and 

eruption (Ryan, 1995). The recognition that Mauna Loa magmas 

have infiltrated Kīlauea’s magma plumbing system (Rhodes and 

others, 1989) lends additional support to this hypothesis. This 

proposal is speculative but provides an interesting alternative 

to requiring higher average rates of magma supply to Kīlauea 

(~0.18 km3/yr, based on CO2 emissions, versus ~0.1 km3/yr 

from deformation and effusion-rate data) and storing that excess 

magma in Kīlauea’s deep rift zones.

Summary

Volcanism from the Hawaiian hot spot is driven by upwelling of high-temperature material, probably originating deep 

within Earth’s mantle. The rate of magma supplied by the plume 

over millions of years has varied by an order of magnitude but 

has been increasing since about 30 Ma and is currently high, 

with the contemporary rate of magma supply to Kīlauea’s shallow magmatic system best approximated by the eruption rate 

of long-term effusions—at least 0.1 km3/yr. Historical changes 

in the rate of magma supply have been driven both by pressure fluctuations in shallow crustal reservoirs (“top-down”) and 

variations in the volume of magma supplied from the mantle 

(“bottom-up”). Bottom-up control of magma supply affects both 

volcanoes, as demonstrated by 2002–07 activity at Kīlauea and 

Mauna Loa. Supply also varies between volcanoes, even when 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 191

Mantle

magma

supply

30-km-deep

melt zone

CO2

bubbles

Mauna Loa

Kīlauea

Lōʻihi

CO2-rich plume

N

Figure 7. Plot of vertical elevation change at Global Positioning Station station MLSP, located on the south side of Mauna Loa’s summit caldera 

(location given in inset). Uplift and inflation began in 2002, accelerated in 2004–05, and gradually waned to no deformation by the end of 2009 

(Miklius and Cervelli, 2003; Amelung and others, 2007).

Figure 8. Schematic diagram of possible magma supply pathways 

beneath the Island of Hawai‘i, based in part on figure 2 of Gonnermann 

and others (2012). Black vertical arrow represents magma supplied 

from the mantle hot spot to a subhorizontal melt zone at about 30-km 

depth (red plane), located beneath Kīlauea’s Southwest Rift Zone, 

based on long-period seismicity and tremor. Colored arrows depict 

magma transport paths to the most active volcanoes of Hawai‘i. Most 

exsolved CO2 might ascend with magma that is fed to Kīlauea, since 

that volcano is closest to the source of mantle supply, resulting in a 

CO2-rich plume from Kīlauea’s summit. 

magma ascends from the hot spot at a steady rate, causing periods of few eruptions at one volcano and vigorous activity at the 

other. For example, since 1952 and through 2014, Mauna Loa 

erupted only twice (in 1975 and 1984), while Kīlauea erupted 

quasi-continuously. Short-term (days to weeks) fluctuations in 

magma supply are superimposed on the overall rate and are “top 

down” processes that are driven, for instance, by pressure in a 

shallow magma reservoir that is low enough to promote magma 

ascent from depth. Kīlauea and Mauna Loa share magma supply 

from the hot spot, which explains observations of both inverse 

patterns of eruptive activity and sympathetic patterns of magma 

accumulation. Differences in mantle source composition for 

Kīlauea and Mauna Loa reflect compositional heterogeneity in 

the asthenospheric melt region that links the two volcanoes and 

through which pressure is transmitted without requiring direct 

melt transport. A model of shared magma supply may also 

explain why CO2 emissions from Kīlauea imply a higher rate 

of supply than is suggested by long-term eruptive activity. CO2

from all hot-spot magma may degas through Kīlauea’s summit, 

regardless of whether that magma ultimately feeds the shallow 

systems of Kīlauea, Mauna Loa, or even Lō‘ihi. 

Magma Storage

Before erupting at the surface, magma is often (although 

not always) stored in subsurface reservoirs, where chemical 

differentiation produces a variety of compositional products. 

Petrologists and structural geologists have spent considerable 

effort attempting to understand the formation of such reservoirs 

(for example, Glazner and Bartley, 2006) and their physical and 

chemical evolution (for example, Marsh, 1989). 

The identification of active magma storage areas is 

most often accomplished by geophysical studies, especially 

2000 2001 2002 2003 2004 2005 2006 2007 2008 2009 2010

-0.2

-0.1

0

0.1

0.2

Vertical deformation, in meters

MLSP

HAWAI‘I

Year
192 Characteristics of Hawaiian Volcanoes

seismology and geodesy (Dvorak and Dzurisin, 1997). For 

instance, earthquake hypocenters in the years following the 

May 18, 1980, eruption of Mount St. Helens, Washington, 

defined a vertically elongate aseismic zone between depths 

of about 7 and 13 km, interpreted to delineate a magma body 

(Scandone and Malone, 1985). Models based on 2004–08 

co-eruptive GPS displacements favored magma storage in the 

same area (Lisowski and others, 2008). Surface inflation near 

South Sister volcano, Oregon, detected by Interferometric 

Synthetic Aperture Radar (InSAR) and corroborated by GPS 

and leveling, suggested magma accumulation at a depth of 

approximately 5 km located about 5 km west of the volcano’s 

summit (Dzurisin and others, 2009). Continuous GPS data 

from Mount Etna, Italy, have been used to map a multilevel 

magma plumbing system, with magma storage areas at 6.5, 

2.0, and 0.0 km below sea level (Aloisi and others, 2011). 

Magma reservoir location and geometry are best determined when a variety of data are available. Deformation, 

seismicity, and geology indicate the presence of at least three 

regions of magma storage beneath Piton de la Fournaise 

volcano, Réunion Island, at about 2.3 km, 7.5 km, and 15 km 

beneath the surface (Peltier and others, 2009). Physics-based 

models incorporating diverse datasets offer hope for constraining not just reservoir depths and locations but also such elusive 

parameters as magma reservoir volume, overpressure, and 

volatile content (Anderson and Segall, 2013).

Mapped regions of magma accumulation and storage 

span a range of geometries, depths, and sizes. Reservoir 

geometry is generally depicted as spherical or ellipsoidal, 

primarily because such shapes are simple to visualize and 

model (for example, Davis, 1986; Delaney and McTigue, 

1994; Ohminato and others, 1998; Yang and others, 1992) 

and are favored by thermal and mechanical considerations 

(Gudmundsson, 2012). High-spatial-resolution deformation 

data sometimes suggest more complex geometries (for 

example, Masterlark and Lu, 2004), and magma reservoirs 

have been described as amalgamations of small bodies 

separated by screens of solid rock or semisolid magma mush 

(for example, Fiske and Kinoshita, 1969; Gudmundsson, 

2012). Kühn and Dahm (2008) pointed out that the stress 

field induced by a plexus of dikes and sills would resemble 

that of a simpler ellipsoidal source. Indeed, the very 

definition of a magma chamber is vague, as magma storage 

areas are not homogenous but are, instead, mush zones 

of melt, crystals, and exsolved volatiles that will have 

different behaviors, effective sizes, and shapes over different 

time scales and strain rates (for example, Johnson, 1992; 

Gudmundsson, 2012).

Depths of magma accumulation vary widely among 

volcanoes. Magma accumulation can occur tens of kilometers 

beneath the surface, as is apparently the case at Hekla, Iceland 

(Ofeigsson and others, 2011), or only 1–2-km deep, as modeled 

at Kīlauea (Poland and other, 2009; Anderson and others, in 

press). At many volcanoes, like Shishaldin, Alaska, no magma 

chamber has been detected in spite of frequent eruptive activity 

(Moran and others, 2006), while other volcanoes have a series 

of vertically stacked reservoirs, like Soufrière Hills volcano in 

Montserrat (for example, Voight and others, 2010), Piton de 

la Fournaise (Peltier and others, 2009), and Etna (Aloisi and 

others, 2011). The depth of magma accumulation depends on 

the density contrast between magma and country rock, as well 

as on the presence of structural discontinuities that may inhibit 

upward magma flow (Ryan, 1987; Burov and others, 2003; 

Peltier and others, 2009). Magma chamber volumes can exceed 

100 km3, as indicated by the volume of ignimbrite sheets that 

represent single eruptive events (for example, Christiansen, 

2001). The largest eruptions aside, magma chambers are more 

generally in the range of tens of cubic kilometers (for example, 

Mastin and others, 2009; Paulatto and others, 2012) to about 

1 km3 (for example, Poland and others, 2009).

The magma plumbing systems of Kīlauea and Mauna 

Loa are well known, compared with most other volcanoes 

on Earth. Magma pathways in Hawaiian volcanoes are 

established early in the shield stage (Clague and Sherrod, 

this volume, chap. 3) but evolve over time, along with the 

volcano, for example, as calderas form and fill (Swanson and 

others, 2012a) and rift zones migrate (Swanson and others 

1976a). Some datasets, like magnetics and gravity, reflect the 

cumulative development of a volcano’s magmatic system. 

We use constraints provided by such data in combination 

with monitoring results (especially deformation and seismic), 

geologic studies, and observations of eruptions collected 

during the 100+ years of HVO’s existence to investigate 

the current magma plumbing configurations at Kīlauea 

and Mauna Loa. Below, we present refined models of the 

magmatic systems at both volcanoes.

Kīlauea

The general model for Kīlauea’s magma plumbing 

system, first proposed by Eaton and Murata (1960) and 

refined by Tilling and Dvorak (1993), is simple: magma 

generated in the mantle ascends and is stored in reservoirs 

that are one to a few kilometers beneath the summit, from 

which it may eventually erupt within the caldera or be 

transported laterally into the East or Southwest Rift Zones 

as intrusions that may feed eruptions far from the summit 

(fig. 9A). While this overall depiction remains largely 

unchanged, the characteristics of specific parts of the magma 

plumbing system have been the focus of numerous studies. 

Ryan (1988) used seismic data to define areas of subsurface 

magma transport and storage, including magma storage at 

2–4-km depth beneath the summit and almost wholly molten 

rift zones at 3–10 km beneath the surface (fig. 9B), although 

the conduit he proposed to 30-km depth beneath Kīlauea’s 

summit was later interpreted to be a mantle fault zone (Wolfe 

and others, 2003). Other studies have employed deformation 

measurements to map the locations of magma reservoirs 

beneath the summit (for example, Cervelli and Miklius, 

2003; Baker and Amelung, 2012; Anderson and others, in 

press), constrain the relation between summit and rift zone 

magma storage (for example, Dzurisin and others, 1980, 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 193

0 20 40 60 80 100 120 140 160 180 200 220 240

Distance, in kilometers

A

B

Depth, in kilometers

0

5

10

20

30

40

Kīlauea Caldera

Southwest Rift Zone East Rift Zone

5

0

-5

-10

-20

-30

-40

-50

-60

Vertical exaggeration 2x

3.0

4.0

5.3

6.8

8.25

Density

2.77 g/cm3

Crust

Mantle Mohorovičić discontinuity

VP

km/s

Density

g/cm3

1.03

4.3 2.3

6.6 2.84

7.9 3.27

Elevation, in kilometers

VP

km/s

Magma source region

Sea level

Pacific

Ocean

Figure 9. Previously published interpretations 

of Kīlauea’s magma plumbing system. A, 

Schematic cross section through an idealized 

Hawaiian volcano, as envisioned by Eaton 

and Murata (1960). Magma ascends from 

about 60-km depth through open conduits and 

collects beneath the caldera. Seismic velocities 

(VP) and densities are indicated for various 

depths. B, Northward-directed view of Kīlauea’s 

internal structure, modified from Ryan (1988). 

The preferred magma pathway within the core 

region of the primary vertical conduit is shaded 

yellow, and lateral magma injection occurs 

along a level of neutral buoyancy 2–4 km 

beneath the surface (horizontal arrows). Magma 

occupies both the shallow and deep parts of the 

rift zones. Model is based on seismicity patterns 

and subsurface density structure. 
194 Characteristics of Hawaiian Volcanoes

1984; Johnson, 1995b), and propose the presence of magma 

in the deep rift zones (Delaney and others, 1990; Owen and 

others, 1995, 2000a; Cayol and others, 2000). Petrologic 

arguments have also been utilized to develop models of 

Kīlauea’s magma plumbing system. Geochemical data from 

historical summit eruptions favor a simple magma storage 

zone of a few cubic kilometers at 2–4 km beneath the summit 

(Pietruszka and Garcia, 1999b; Garcia and others, 2003). 

Primitive compositions and deformed olivine phenocrysts 

derived from cumulate rocks have been interpreted as 

evidence that magma may bypass summit storage on its way 

to erupting on the ERZ (for example, Trusdell, 1991; Wright 

and Helz, 1996; Vinet and Higgins, 2010). 

Geodetic data collected in the 1990s and 2000s, combined 

with seismic and petrologic data, provided improved resolution 

of the geometry of Kīlauea’s magma storage zones and 

transport pathways. This improvement is largely due to the 

excellent temporal resolution of continuous ground-based 

sensors, including GPS stations and borehole tiltmeters, and 

outstanding spatial resolution from InSAR. We synthesize 

these data, especially those acquired during the 2003–07 

magma supply increase to Kīlauea (see “2003–07 Increase in 

Magma Supply to Kīlauea” section and fig. 4), with previously 

published studies to propose a refined model for the geometry 

of magma storage and transport at Kīlauea that is consistent 

with petrologic, geophysical, and geologic data (fig. 10).

Our model contains several elements: (1) a summit 

magma system consisting of two long-term reservoirs, one at 

~3 km beneath the south caldera (SC in fig. 10) and a second 

at 1–2 km beneath the caldera center (H in fig. 10), as well 

as occasional magma storage beneath Keanakāko‘i Crater 

(K in fig. 10); (2) a seismic Southwest Rift Zone (so called 

because it is easily recognized from earthquake hypocenters) 

at ~3-km depth that is connected to the south caldera magma 

reservoir and that stores magma during times of heightened 

summit magma pressure; (3) an East Rift Zone with a molten 

core (from the summit to at least its distal subaerial extent) 

at ~3-km depth (with compositionally isolated pods of stored 

magma at that depth and shallower), which is also connected 

to the south caldera magma reservoir and has been active 

continuously since 1983; (4) a volcanic Southwest Rift Zone 

(so named because it hosts more eruptive vents and fissures 

than the seismic Southwest Rift Zone) within ~1 km of the 

surface that is connected to the shallower summit magma 

reservoir and extends southwest from Halema‘uma‘u Crater; 

and (5) a “Halema‘uma‘u-Kīlauea Iki trend” (abbreviated 

HKIT in fig. 10) within ~1 km of the surface that is also 

connected to the shallower summit magma reservoir but 

extends east from Halema‘uma‘u Crater towards Kīlauea 

Iki. The model does not include a deep rift zone that allows 

magma to bypass the summit magma storage system and 

intrude into, or erupt from, the rift zones. Instead, we favor 

the model of Johnson (1995b), who proposed that the East 

and Southwest Rift Zones host molten cores at 3–5-km depth 

and supply magma vertically both towards the surface and to 

greater depths. Below, we describe these zones in detail.

Summit Magma Storage

Storage of magma beneath Kīlauea’s summit was 

suspected by observers in the 1800s, including the first 

non-Hawaiian visitor to the volcano, William Ellis (Ellis, 

1825), and no doubt by Hawaiians before then, on the basis 

of summit eruptive activity and the presence of the caldera, 

which was thought to be related to removal of subsurface 

magma (see discussion in Peterson and Moore, 1987). Early 

geophysical evidence for magma storage was provided mostly 

by deformation measurements. For example, deflation of the 

summit associated with magma withdrawal in 1924 (fig. 11) 

resulted in subsidence of several meters near Halema‘uma‘u 

Crater (which collapsed as a result of the activity) and 

horizontal displacements of over 1 m toward the center of 

the caldera (Wilson, 1935; Dvorak, 1992). These data were 

used by Mogi (1958) in a now-classic paper that modeled 

the deformation as due to pressure decrease caused by 

withdrawal of magma from point sources at 3.5-km and 25-km 

depth (although Eaton (1962) rejected the deeper source as 

an artifact of scale error in the wooden leveling rods, and 

Dvorak (1992) found that a single source at 4.5-km depth was 

sufficient to fit the data). 

Tilt measurements that started in 1913 with horizontal 

pendulum seismographs (Finch, 1925) and continued with 

the development of water tube tilt measurements in the 1950s 

(Eaton, 1959) revealed that the summit was almost always 

in some state of inflation or deflation, often associated with 

eruptive and earthquake activity (fig. 12). Seismic evidence for 

magma storage was provided by data from the modern seismic 

network installed at Kīlauea in the 1950s (Eaton and Murata, 

1960; Eaton, 1962; Okubo and others, this volume, chap. 2), 

which made high-resolution earthquake locations possible.

The geometry of magma storage is complex. Fiske and 

Kinoshita (1969) tracked inflationary deformation during 

the 22.5 months between Kīlauea’s 1965 ERZ eruption and 

1967–68 summit eruption, noting that the locus of maximum 

uplift changed over time (fig. 13A), with deformation during 

several time periods best approximated by a source at about 

2–3-km depth beneath the southern part of the caldera. 

Their conceptual model of the summit magma system was a 

plexus of dikes and sills, with different zones of the system 

activated at different times—similar to the model of Dieterich 

and Decker (1975) for the same time period, and a proposal 

adopted by numerous authors to explain changes in surface 

deformation over time (for example, Schimozuru, 1981; Yang 

and others, 1992). Ryan and others (1981) and Ryan (1988) 

proposed a series of interconnected magma storage and 

transport zones beneath Kīlauea’s summit and rift zones from 

seismic data and numerical models.

Deformation during the 2003–07 magma supply increase 

(see section entitled “2003–07 Increase in Magma Supply to 

Kīlauea” and Poland and others, 2012) followed a progression 

that highlights several discrete areas of magma accumulation 

(fig. 14). InSAR provides an especially clear view of 

deformation that is indicative of distinct magma storage areas 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 195

KīlaueaCaldera Mauna

Ulu

Nāpau

Crater

Pu‘u ‘Ō‘ō Heiheiahulu

Kapoho Crater

H

SC

K

Makaopuhi

Crater

Mauna Loa

HKIT

Se si mci SWRZ

0123456789

10

Depth below surface, 

in kilometers

Basal décollement

Halema

uma ʻ u- ʻ

Kīlauea Iki trend

Volcanic SWRZ

Seismic SWRZ

East Rift Zone H

K SC

Koa

e Fault System ʻ

Halemaʻumaʻu Crater

00

2 4 KILOMETERS

1 2 MILES

Vo cl an ci SWRZ

East Rift Zone

Deep rift zone

Figure 10. Illustration of proposed structure of Kīlauea’s subsurface magma plumbing system. Schematic cut-away shows a cross section through Kīlauea’s summit and rift zones. Magma pathways and storage areas are exaggerated in size for clarity. H, Halema‘uma‘u reservoir; K, Keanakāko‘i reservoir; SC, south caldera reservoir; SWRZ, Southwest Rift Zone. Plan view gives the relations of magma pathways to surface features and topography in the vicinity of Kīlauea Caldera.
600

-600

-400

-200

0

200

400

1956 1961 1966 1971 1976 1981 1986 1991 1996 2001 2006 2011

North-south

Tilt at Uēkahuna, in microradians

East-west

Year

196 Characteristics of Hawaiian Volcanoes

Figure 11. Map of deformation at Kīlauea’s summit 

during the 1920s. Arrows show horizontal displacements 

from triangulation surveys spanning 1922–26, and 

contours give vertical displacements from leveling 

surveys spanning 1921–27. Maximum measured 

subsidence is about 4 m. Most of this deformation was 

associated with downdrop of the Halema‘uma‘u lava 

column and enlargement of Halema‘uma‘u Crater in 

1924. Adapted from figures 2 and 8 of Wilson (1935).

Figure 12. Plot of north-south (red) and east-west (blue) tilt from 1956 through 2011 at the Uēkahuna water-tube tiltmeter, located about 300 m 

west-northwest of the Hawaiian Volcano Observatory on the northwest rim of Kīlauea Caldera. Positive tilt is to the north or east, and negative tilt to 

the south or west. Northwest tilt indicates inflation and southeast tilt, deflation. Some offsets are associated with large earthquakes (for example, the 

1975 earthquake on Kīlauea’s south flank). A large offset associated with the November 1983 Ka‘ōiki earthquake, beneath Mauna Loa’s southeast 

flank, has been removed from the time series because the vault that houses the tiltmeter was damaged by the shaking.

-0.9

-1.2

-1.5

-1.8

-2.1

Horizontal displacement vector

1.5 meters

Contour interval 0.3 meters

 155°20'

19°20'

19°22'

19°24'

19°26'

 155°18' 155°16' 155°14'

0

0

2 KILOMETERS

1 MILE

1
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 197

Figure 13. Deformation and seismic data indicating magma storage 

beneath Kīlauea Caldera. A, Map of centers of uplift determined 

by leveling during 1966–67. Dashed line indicates trend of section 

in part B. Adapted from figure 5 of Fiske and Kinoshita (1969). B, 

Cross section through Kīlauea Caldera showing seismic velocities 

(in kilometers per second) and shallow earthquakes in the vicinity of 

the caldera during 1980–81 (circles). Zone of low seismic velocity 

surrounded by earthquakes beneath the caldera suggests the 

presence of magma. Cross section extends both northwest and 

southeast of the dashed line in part A, and caldera bounds are 

indicated at the top of the section. From figure 3 of Thurber (1984).

beneath the summit (Baker and Amelung, 2012). Inflation 

associated with subsurface magma accumulation was centered 

just east of Halema‘uma‘u Crater in 2003 (fig. 3, 14A), near 

Keanakāko‘i Crater in 2004–05 (fig. 3, 14B), and in the south 

caldera in 2006 (fig. 14C). Synthesizing these results with past 

studies of deformation, gravity, and seismicity suggests two 

long-term, interconnected magma reservoirs: one at ~3 km 

beneath the south caldera (which we term the “south caldera 

reservoir”) and another at ~1–2 km beneath the caldera center 

(our “Halema‘uma‘u reservoir”). In addition, at least one region 

of intermittent storage exists beneath the Keanakāko‘i Crater 

area near where the ERZ and summit intersect. While satisfying 

geophysical data, the presence of multiple discrete magma 

reservoirs beneath Kīlauea’s summit is also consistent with 

petrologic observations of the preservation of distinct magma 

batches over time (Helz and others, this volume, chap. 6).

South Caldera Reservoir

The southern part of Kīlauea Caldera has long been 

recognized as the main locus of persistent deformation and, 

therefore, the main region of magma storage beneath the 

summit. Starting in the late 1950s and early 1960s, models 

of deformation consistently located a source of volume 

change beneath the south caldera at about 3–4-km depth 

(Eaton, 1959, 1962). The map-view location of the magma 

reservoir is determined geodetically and, thus, is independent 

of models for depth and shape. Dvorak and others (1983) 

modeled numerous episodes of inflation and deflation 

during 1966–70 using point sources, which showed a 

clustering at 2–4 km beneath the south caldera. Davis (1986) 

introduced an ellipsoidal source located in the same place 

to fit deformation of Kīlauea’s summit during the 1970s. An 

ellipsoidal source with a center about 2 km below the south 

caldera was also modeled by Dieterich and Decker (1975) 

from vertical and horizontal deformation data collected 

during inflation in January–February 1967, although the top 

of their reservoir extended to less than 1 km beneath the 

surface. Yang and others (1992) suggested that all deflations 

and inflations were caused by volume changes in a single 

spherical body at 2.6 km beneath the south caldera, and 

that migrating deformation maxima during periods of uplift 

reflected dike intrusions beneath the caldera center and upper 

parts of the rift zones. The geodetically determined location 

of the south caldera magma reservoir correlates with both an 

aseismic zone at 3–6-km depth (Koyanagi and others, 1976; 

Ryan and others, 1981) and low P-wave velocities (Thurber, 

1984; Rowan and Clayton, 1993)—characteristics consistent 

with magma storage (fig. 13B).

Summit deformation during the first 20 years of Kīlauea’s 

1983–present (as of 2014) ERZ eruption was dominated 

by deflation centered on the south caldera, which has been 

modeled by several authors as a small amount of pressure (or 

volume) loss at ~3-km depth (Delaney and others, 1990, 1993; 

Owen and others, 1995, 2000a; Cervelli and Miklius, 2003; 

Baker and Amelung, 2012). Similarly, gravity data collected 

during the eruption found mass changes centered on the south 

caldera with mass loss of only a few percent of the erupted 

volume, which supports interpretations based on deformation 

data that the south caldera reservoir served as a waypoint for 

magma that entered the volcano en route to the eruption site 

(Johnson, 1992; Kauahikaua and Miklius, 2003; Johnson and 

others, 2010). Inflation of the south caldera and the upper 

part of the Southwest Rift Zone (SWRZ) in 2006 was imaged 

with excellent spatial resolution by InSAR (fig. 14C; Myer 

and others, 2008; Baker and Amelung, 2012). Modeling the 

inflation observed by InSAR and GPS—and assuming a region 

1

2

5 4 3 6 7

8 9

10

Halemaʻumaʻu Crater 

(1) January–July 1966

(2) July–October 1966

(3) October 1966–January 1967

(4) January–February 1967

(5) February–February 1967

(6) February–May 1967

(7) May–June 1967

(8) June–July 1967

(9) July–September 1967

(10) September–October 1967

19°24'

19°26'

 155°18' 155°16' 155°14'

Keanakākoʻi Crater 

HVO

0

4

8

0 5 10 15 20 25

Distance along section, in kilometers

Depth below sea level,

in kilometers

Caldera

3 km/s

4

5

6

8 7

A

B

0

0

1 KILOMETER

0.5 MILE
198 Characteristics of Hawaiian Volcanoes

of distributed opening beneath the south caldera and upper 

SWRZ (see appendix for details on modeling procedure)—

suggest a source depth of 3 km (95-percent confidence bounds 

are 2.1–5.1 km) beneath the south caldera and upper SWRZ 

(fig. 15), essentially the same location as that modeled by 

numerous other workers.

Volume estimates of the south caldera magma reservoir 

are diverse but generally cluster between 3 and 20 km3. 

Dawson and others (1999) identified a low P-wave velocity 

anomaly in the south caldera with a volume of 27 km3 (on 

the basis of a 5 percent reduction in velocity contrast from an 

initial model). Decker (1987) pointed out that magma storage 

must have been, at some point, at least the volume of the 

caldera collapse, estimated at 3 km3 and possibly as large as 

22 km3, assuming a spherical shape with a diameter of 3.5 km 

(roughly equivalent to the caldera dimensions). Only part of 

this region may currently be molten, however, which would 

bring the volume closer to the 11 km3 estimate that Wright 

(1984) derived from geodetic data. Johnson (1992) calculated 

an effective volume (that is, the volume that behaves like a 

2007/04/12 – 2007/06/21

32o

2003/09/29 – 2004/07/05

20o

2004/04/19 – 2005/10/31

22o

2007/07/29 – 2009/11/15

40o

Range change

0 28.3 mm

N

2006/01/15 – 2007/05/09

40o

A B C

D E

0

0

2 4 KILOMETERS

1 2 MILES

fluid) of about 13 km3. At the low end of the spectrum is a 

volume of 0.08 km3, based on the magma supply rate and 

average repose period between eruptions (Klein, 1982). 

Residence time analysis of rapid geochemical fluctuations led 

Pietruszka and Garcia (1999b) and Garcia and others (2003) 

to infer a simple source geometry with a volume of 2–3km3, 

which may be smaller than most geophysical estimates 

because it represents the hotter core of the reservoir in which 

magma mixing occurs. As an upper bound, Denlinger (1997) 

calculated a volume of 240 km3 from the ratio of pressure 

change to volume, but this estimate is probably more reflective 

of the entire volume of magma within and beneath Kīlauea 

and not solely that of the south caldera reservoir (Ryan, 1988). 

The wide range of these volume estimates demonstrates 

the difficulty in defining just what constitutes a magma 

reservoir. Zones of magma storage probably grade from host 

rock to molten liquid and will contain regions of crystal mush 

and exsolved volatiles that will respond differently to applied 

stress (for instance, changing pressurization due to episodes of 

magma accumulation and withdrawal), depending also on the 

Figure 14. Interferograms detailing surface deformation of Kīlauea’s summit area during the 2003–07 magma supply increase. Data are from the 

Advanced Synthetic Aperture Radar (ASAR) instrument on the Envisat satellite. For all images, dates spanned are in the lower right, and upper left 

inset gives flight direction (arrow) and look direction (orthogonal line with angle in degrees from vertical). Color scale in the lower right panel applies 

to all five interferograms. One fringe is 28.3 mm of range change along the radar line of sight, where positive (magenta to yellow) indicates increasing 

range, that is, ground motion away from the satellite: subsidence. A, Uplift is focused within the caldera, near Halema‘uma‘u Crater. B, Uplift has 

shifted to near Keanakāko‘i Crater. C, Uplift rate has increased and is centered on the south caldera and upper Southwest Rift Zone. D, Subsidence 

near Halema‘uma‘u Crater and uplift of the East Rift Zone caused by magma withdrawal from the summit and intrusion into the East Rift Zone during 

June 17–19, 2007. E, Subsidence centered in the south caldera and upper Southwest Rift Zone, as well as in the East Rift Zone.
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 199

Figure 15. Source model of summit inflation due to volume increases in the south caldera and upper Southwest Rift Zone. Interoferometric 

Synthetic Aperature Radar (InSAR) and Global Positioning Ssystem (GPS) data span May 13–September 30, 2006. InSAR data are from the 

Advanced Synthetic Aperture Radar (ASAR) instrument on the Envisat satellite, beam mode 3, track 365. Arrows are horizontal displacement vectors 

scaled according to magnitude of deformation; where small, only arrowhead may show. In contrast to figure 14, one fringe represents 28.3 mm of 

displacement, where positive (magenta to yellow) indicates line-of-sight uplift. A, Observed deformation from InSAR and GPS. B, Modeled GPS and 

InSAR displacements assuming distributed opening of a sill-like source at 3-km depth. Boxes indicate sill elements that opened by 1 cm or more. C, 

Residual (observed minus modeled) GPS and InSAR displacements based on the model in part B. D, Map showing distribution of model sill opening.

magnitude and time scale of the stress (Gudmundsson, 2012). 

Such a question is inherent to any study of magma storage. 

We therefore caution against strict interpretations of reservoir 

volumes without a thorough understanding of the underlying 

assumptions that went into volume calculations.

Halema‘uma‘u Reservoir

A second long-term zone of magma storage beneath 

Kīlauea’s summit caldera is suggested by deformation results 

that indicate volume/pressure change beneath, or east of, 

the east margin of Halema‘uma‘u Crater (see, for example, 

location 2 in fig. 13A). Leveling and tilt data from the summit 

over various time periods show deformation centered just 

east-northeast of Halema‘uma‘u Crater, distinct from the 

south caldera reservoir (Fiske and Kinoshita, 1969; Dvorak 

and others, 1983), although the presence of an independent 

magma storage zone in this area (as opposed to occasional 

dike intrusions) was not proposed until later (Cervelli and 

Miklius, 2003). The reservoir probably feeds the eruptive vent 

within Halema‘uma‘u Crater that formed in March 2008, but 

the connection between the vent and storage area is complex 

(Chouet and Dawson, 2011).

The installation of electronic borehole tiltmeters at 

the summit of Kīlauea, starting in 1999, provided further 

evidence of a magma storage reservoir beneath the east 

margin of Halema‘uma‘u Crater. The tiltmeters record 

repeated transient tilt events that last hours to days and 

are characterized by sudden deflation, followed by equally 

sudden inflation (“DI” events; see “Characteristics of 

Magma Transport at Hawaiian Volcanoes” section and 

Anderson and others, in press). Cervelli and Miklius (2003) 

modeled four of these events as due to a point source of 

pressure change at 500–700-m depth about 0.5 km east of 

Halema‘uma‘u Crater. With the start of summit eruptive 

activity in 2008 (ongoing as of 2014), DI events became 

much more common (~5–10 per year during 1999–2007, 

before the eruption, compared to more than 50 per year 

in 2009–13, during the eruption). Modeling the tilt events 

(see appendix for details on modeling procedure) as due to 

a point source of pressure change (a more complex source 

geometry is not resolvable with the limited number of tilt 

stations at the summit) indicates a location about 1–2 km 

beneath the east margin of Halema‘uma‘u Crater (fig. 16), 

although the depth is poorly constrained. Anderson and 

others (in press) inverted tilt data recorded during more than 

450 DI events that occurred between 2000 and 2013 and 

found the same location and depth range, indicating that the 

source position does not vary over time. 

Rapid deflation of the Halema‘uma‘u reservoir 

has been documented repeatedly since the start of the 

1983–present (as of 2014) ERZ eruption and is a result of 

magma drainage to feed ERZ intrusions and eruptions, as 

exemplified by activity in 1997 (Owen and others, 2000b), 

2007 (Montgomery-Brown and others, 2010), and 2011 

(Lundgren and others, 2013). Models of these subsidence 

events suggest a source depth of 1–2 km beneath the east 

margin of Halema‘uma‘u Crater (Poland and others, 2009; 

Montgomery-Brown and others, 2010; Lundgren and others, 

2013). As an example, inversion of InSAR and GPS data 

for summit subsidence associated with the 2011 ERZ fissure 

eruption (see appendix for details on modeling procedure) 

gives a depth of 1.4 km (95 percent confidence range is 

1.0–1.9 km; fig. 17).

0

0.1

0.2

0.3

0.4

0.5

Opening, in meters

A B C D

Displacement (mm)

5 cm 0 28.3

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0

2 4 KILOMETERS

1 2 MILES

19°22'

 155°18'

19°18'

 155°22'
200 Characteristics of Hawaiian Volcanoes

500 1,500 2,500 0

2,000

4,000

6,000

8,000

10,000

Optimal depth = 1,559 m

Number of models

Depth, in meters

−1 -0.5 0 0.5 1 0

2,000

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6,000

8,000

Optimal

volume

change =

0.13 x 106 m3

Volume change,

in cubic meters x 106

UWE

SDH

SMC

IKI

2 microradians

19°24'

19°26'

 155°18' 155°16' 155°14'

0

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Figure 16. Source modeling of transient tilt events at Kīlauea’s summit. Modeling assumes a point-source geometry. Arrows show magnitude and 

direction of observed (blue) and predicted (red) tilts for a sample model at the sites of four tiltmeters (IKI, SDH, SMC, and UWE) positioned around 

Kīlauea Caldera. Best-fitting source location marked by red dot. Distributions of depth and volume change for the sample model are given at right. 

Source locations for models of 151 additional tilt events that occurred during 2000–09 are given by gray circles. Depth ranges for these additional 

models are widespread but generally fall between 0.2 and 2 km beneath the surface.

Seismic and gravity evidence for the Halema‘uma‘u 

reservoir is also convincing. Broadband seismic stations 

in Kīlauea Caldera detected very-long-period (VLP) 

tremor (Ohminato and others, 1998) and VLP seismic 

events (Almendros and others, 2002; Chouet and others, 

2010) originating about 1 km beneath the east margin of 

Halema‘uma‘u Crater, possibly an indication of flowing 

magma. High-resolution (0.5 km) velocity models imaged 

a low-velocity P-wave anomaly in about the same location 

(Dawson and others, 1999), which corresponds to a cluster 

of long-period (LP) seismicity (Battaglia and others, 2003); 

tremor and LP events also occur just above this region 

(Almendros and others, 2001). Finally, gravity data spanning 

1975–2008 detected mass increase beneath the east rim of 

Halema‘uma‘u Crater, although no long-term deformation 

was measured in this location. Models of the gravity data 

suggest a source depth of 1 km, and the lack of deformation 

accompanying the gravity increase implies that magma was 

accumulating in void space that may have been created when 

several tens of millions of cubic meters of magma drained 

from the summit as a result of the 1975 earthquake (Dzurisin 

and others, 1980; Johnson and others, 2010).

The Halema‘uma‘u magma reservoir is an order of 

magnitude smaller than the south caldera reservoir. Johnson 

(1992) suggested that the volume was at least 1.6 km3, based 

on summit deflation associated with draining of the Mauna Ulu 

lava lake in 1973. A model of deformation and gas emission 

data during rapid deflation of the source associated with the 

June 2007 ERZ intrusion and eruption led Poland and others 

(2009) to determine a volume of 0.2–1.2 km3. Anderson and 

others (in press) related lava-level changes within the summit 

eruptive vent to ground tilt recorded during two especially 

large-magnitude DI events in February 2011 and found a 

volume of 0.15 to 2.7 km3. Segall and others (2001) used 

deformation associated with the brief 1997 ERZ eruption to 

model a volume for the Halema‘uma‘u magma reservoir—

shown by Owen and others (2000b) to be the deflation source at 

Kīlauea’s summit during that event—of ~20 km3. They assumed 

a value for the elastic modulus of 20 GPa; reducing this by an 

order of magnitude, which is reasonable, given the fractured 

nature of Kīlauea’s shallow subsurface (for example, Rubin and 

Pollard, 1987), would correspondingly decrease the modeled 

volume to ~2 km3 (Anderson and others, in press).

Keanakāko‘i Reservoir

Multiple authors have noted that modeled locations of 

long-term summit subsidence (in other words, subsidence 

not associated with rapid drainage of magma to feed rift zone 

intrusions and eruptions) cluster beneath the south caldera, 

whereas inflation sources are distributed over a broader area 

(for example, Dvorak and others, 1983; Yang and others, 

1992; fig. 13A). This observation led Yang and others (1992) 

to suggest that inflation of the south caldera reservoir was 

accompanied by dike intrusion, but the intruded dikes did 

not subsequently close during reservoir deflation. As an 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 201

Figure 17. Source model of summit subsidence associated with the 2011 Kamoamoa fissure eruption. Interoferometric Synthetic Aperature Radar 

(InSAR) and Global Positioning System (GPS) data span January 19–March 6, 2011. InSAR data are from the PALSAR instrument on the ALOS 

satellite, orbital path 598. Color scale and arrows as in figure 15. A, Observed deformation from InSAR and GPS. B, Modeled GPS and InSAR 

displacements assuming a point source of volume change (white circle) at 1.4 km beneath the northeast margin of Halema‘uma‘u Crater. C, Residual 

(observed minus modeled) GPS and InSAR displacements, based on the point source model in part B. D–G, Histograms showing the distributions of 

model parameters, including east location (D, with arbitrarily sourced x-axis), north location (E, with arbitrarily sourced x-axis), depth (F), and volume 

change (G). Dashed red lines indicate best-fitting solutions, which do not necessarily align with the peaks of the distributions.

alternative, we suggest that the distribution of inflationary 

centers reflects the combined effects of several persistent 

zones of magma storage. In addition to the Halema‘uma‘u 

and south caldera reservoirs, magma is also stored beneath 

the area near Keanakāko‘i Crater (fig. 3).

Relocated LP earthquakes cluster at about 4-km depth 

beneath Keanakāko‘i Crater (Battaglia and others, 2003), 

which is also a region of low P-wave velocity (Dawson 

and others, 1999). In addition, uplift has frequently been 

localized in that area during periods of summit inflation; for 

example, between the 1959 Kīlauea Iki and 1960 Kapoho 

eruptions (Wright and Klein, 2014), before the 1967–68 

summit eruption (location 1 in fig. 13A; Fiske and Kinoshita, 

1969), and before the 1974 summit eruption (Lockwood and 

others, 1999). Many dikes modeled for episodes of uplift 

cluster near Keanakāko‘i Crater, and seismicity indicates that 

ERZ intrusions are frequently initiated near Keanakāko‘i 

(Yang and others, 1992; Klein and others, 1987), both 

suggesting magma storage in the area. Historical eruptions 

in this region occurred within Keanakāko‘i Crater in 1877 

(Peterson and Moore, 1987) and nearby in 1971 (Duffield 

and others, 1982) and 1974 (Lockwood and others, 1999).

InSAR data that span 2004–05 indicate uplift immediately 

north and east of Keanakāko‘i Crater as part of the sequence 

of deformation associated with the 2003–07 magma supply 

increase (Baker and Amelung, 2012; fig. 14B). A model of 

InSAR and GPS displacements from 2004–05 (see appendix 

for details on modeling procedure) suggests a deformation 

source depth of about 2.6 km (95 percent confidence range is 

2.0–4.9 km) near Keanakāko‘i Crater (fig. 18).

We propose that the localized uplift near Keanakāko‘i 

Crater observed occasionally during periods of caldera 

inflation represents a transient accumulation of magma. The 

accumulation may occur due to a backup of magma that cannot 

be accommodated by the ERZ conduit. Eruptive activity in 

2004–05 included a sudden increase in the effusion rate from the 

ERZ eruptive vent in February 2005 (Poland and others, 2012), 

suggesting that the rift conduit began to carry more magma from 

the summit. Continued summit inflation after this date, including 

uplift near Keanakāko‘i Crater (figs. 14B and 18), implies that 

the rift conduit was full and could not transport larger volumes of 

magma towards the ERZ eruption site. Pressure increase where 

the ERZ and summit magma systems intersect near Keanakāko‘i 

Crater, due to conduit back-ups, such as the one that occurred in 

A B C Displacement (mm)

0 28.3

4 cm

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19°25'

-2 -1 0 1 2 0

1,000

2,000

3,000

4,000

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Number of models

East location, in kilometers

D

7 7.5 8 8.5 0

500

1,000

1,500

2,000

2,500

North location, in kilometers

E

1 1.2 1.4 1.6 1.8 2

0

400

800

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Depth, in kilometers

F

-12 -8 -4 0

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G

Volume change, in

million cubic meters

 155°17' 155°16'

0

0

1 KILOMETER

0.5 MILE
202 Characteristics of Hawaiian Volcanoes

Figure 18. Source modeling of summit inflation recorded by Interoferometric Synthetic Aperature Radar (InSAR) and Global Positioning System 

(GPS) data spanning November 3, 2004–October 19, 2005. InSAR data are from the ASAR instrument on the Envisat satellite, beam mode 2, track 

429. Color scale and arrows as in figure 15. A, Observed deformation from InSAR and GPS. B, Modeled GPS and InSAR displacements assuming a 

point source of volume change (white circle) at 2.6-km depth beneath the south part of the caldera, near Keanakāko‘i Crater. C, Residual (observed 

minus modeled) GPS and InSAR displacements based on the point source model in part B. D–G, Histograms showing the distributions of model 

parameters, including east location (D, with arbitrarily sourced x-axis), north location (E, with arbitrarily sourced x-axis), depth (F), and volume change 

(G). Dashed red lines indicate best-fitting solutions, which do not necessarily align with the peaks of the distributions.

2005, may eventually lead to an eruption from this accumulation 

zone, as exemplified by activity in 1971 and 1974. Each of 

those eruptions was preceded by waning activity at Mauna Ulu, 

presumably due to a blockage between the summit and that vent, 

and the lava erupted at the summit was similar in composition to 

that which had been erupted during the preceding activity at the 

then-active Mauna Ulu (Duffield and others, 1982; Lockwood 

and others, 1999). Indeed, the very presence of Keanakāko‘i 

Crater, and the existence of pit craters in general, argues for 

magma storage, since the crater may have formed during sudden 

drainage of magma from a subsurface storage area (for example, 

Swanson and others, 1976b).

Rift Zones

Kīlauea is generally described as having two rift zones that 

radiate to the east and southwest from the summit. Geologic data 

suggest that each rift zone is comprised of two distinct magma 

pathways with different trends and surface expressions (for 

example, Holcomb, 1987; Fiske and others, 1993)—a model that 

is also supported by geophysical results. The shallower pathways, described below as the volcanic Southwest Rift Zone and 

Halema‘uma‘u-Kīlauea Iki Trend, are within 1 km of the surface 

and are fed from high-level parts of the summit magma system (specifically, the Halema‘uma‘u reservoir). Slightly deeper 

pathways at about 3-km depth—the East Rift Zone and seismic 

Southwest Rift Zone, as described below—are fed by the south 

caldera reservoir (fig. 10). The volcanic Southwest Rift Zone 

and Halema‘uma‘u-Kīlauea Iki Trend are the sites of numerous 

eruptions but are structurally superficial to the volcano (Fiske and 

others, 1993), at least during recent times, while the slightly deeper 

seismic Southwest Rift Zone and East Rift Zone represent the current structural boundaries of the mobile south flank (for example, 

Cayol and others, 2000). The East and Southwest Rift Zones may 

therefore be viewed as parallel structures, each with both shallow 

and deeper magma pathways that are connected to the summit 

magma system at different depths, and with the shallower pathways 

located north of the deeper pathways. Such geometry may be a 

consequence of seaward migration of both rift zones over time, as 

suggested by Swanson and others (1976a). In addition, both rift 

zones are underlain by a zone of deep extension (see the “Deep Rift 

Zones” section below) that has been proposed as a region of magma 

accumulation and transport (for example, Delaney and others, 

1990), but for which evidence for magma storage is contradictory.

Number of models

4 5 6 7 8 0

1,000

2,000

3,000

2 3 4 5 0

400

800

1,200

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East location, in kilometers North location, in kilometers Depth, in kilometers Volume change, in

million cubic meters

D E F G

Displacement (mm)

0 28.3

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2 KILOMETERS

4 cm 1 MILE

A B C

19°24'

 155°16'

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Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 203

Seismic Southwest Rift Zone

Southwest of Kīlauea Caldera is a broad alignment of 

eruptive vents, fractures, and seismicity that extends toward 

the coast. This zone, the Southwest Rift Zone (SWRZ), is 

composed of distinct strands (Holcomb, 1987). We follow 

previous authors in recognizing two primary strands, one 

defined largely by seismicity (the “seismic SWRZ”) and 

the other marked mostly by alignments of fissures and 

eruptive vents (the “volcanic SWRZ,” described in a section 

of the same name below). The seismic SWRZ, following 

the naming convention of Wright and Klein (2006, 2014), 

defines the boundary between the stable north sector of the 

volcano and the mobile south flank, based on modeling of 

rift zone opening (Cayol and others, 2000), and corresponds 

to the “middle rift strand” of Holcomb (1987).

The seismic SWRZ is marked by earthquakes that trend 

south and then southwest from the south caldera magma 

reservoir at a depth of ~3 km, southwest of the alignment of 

eruptive vents and fissures that defines the volcanic SWRZ 

(Klein and others, 1987; Fiske and Swanson, 1992). The seismic 

swarms mark magmatic intrusions, as indicated by deflation of 

the summit during propagation of earthquakes to the southwest, 

such as in 1974–75 (Lockwood and others, 1999). That episode 

resulted in the only post-18th century eruption from the seismic 

SWRZ on December 31, 1974. The eruption lasted a single 

day and was characterized by gas-rich lava that formed shelly 

pāhoehoe and fountain-fed ‘a‘ā flows. In the days following the 

eruption, deformation and seismic evidence indicated intrusion 

of magma to lower parts of the seismic SWRZ (fig. 19A), but no 

lava erupted (Lockwood and others, 1999).

Since 1974, at least three additional magmatic intrusions 

have occurred in the seismic SWRZ, as judged by deformation 

measurements and seismicity (many additional intrusions 

occurred before 1974, as well; Klein and others, 1987). The 

first half of 1981 was characterized by small earthquake 

swarms in the seismic SWRZ that led Klein and others (1987) 

to hypothesize a continuous intrusion that culminated in August 

(fig. 19B). In June 1982, an earthquake swarm in the seismic 

SWRZ (fig. 19C) was accompanied by summit deflation and 

rift extension. Wallace and Delaney (1995) modeled this event 

as a dike intrusion coupled with seaward slip of the volcano’s 

south flank. In April 2006, uplift was detected by InSAR 

(fig. 14C) and GPS (fig. 4A) in the portion of the seismic 

SWRZ near the caldera (Myer and others, 2008; Baker and 

Amelung, 2012; Poland and others, 2012). This deformation 

is best modeled as inflation of a sill-like body beneath the 

south caldera and upper seismic SWRZ at a depth of 3 km 

(fig. 15) and is correlated in time with shallow, high-frequency 

seismicity in the same area (fig. 19D).

Seismicity related to intrusions into the seismic SWRZ 

extends as far as the Kamakai‘a Hills, where the trace of 

epicenters bends abruptly southward and the seismicity 

deepens as if the intrusions activate south-flank faults (fig. 

19; Klein and others, 1987). Deep extension modeled from 

deformation data terminates in about the same place (Cayol 

and others, 2000). Geophysical data therefore indicate that the 

seismic SWRZ does not extend all the way to the coast but, 

instead, ends somewhere in the vicinity of the Kamakai‘a Hills.

When not inflating, the normal deformation mode of 

the seismic SWRZ is opening and subsidence. Leveling 

results during 1996–2002, although dominated by south 

caldera deflation, include a component of subsidence along 

the seismic SWRZ (fig. 5 in Cervelli and Miklius, 2003). 

GPS data from station KOSM, located about 5 km southwest 

of Halema‘uma‘u Crater (fig. 3), also show persistent 

subsidence of several centimeters per year before and after 

the 2003–07 magma supply surge (fig. 4A). Most of this 

subsidence is probably a result of rift-zone opening due to 

south-flank motion, as indicated by models of deformation 

data (Johnson, 1987; Owen and others, 2000a; Cayol and 

others, 2000), although a component of the subsidence may 

be caused by magma withdrawal (Johnson, 1995b) and (or) 

cooling and contraction of stored magma.

East Rift Zone

Post-18th century eruptions from the ERZ were sparse 

before the 1950s, occurring in 1840, 1922, and 1923 (Holcomb, 1987). Starting in the 1950s, the ERZ became the most 

active part of Kīlauea’s magmatic system, with weeks-long 

eruptions from the lower part of the rift zone in 1955 and 

1960. At least eight eruptions occurred in the 1960s from the 

middle ERZ, culminating in the 1969–74 eruption of Mauna 

Ulu—the longest-lasting post-18th century eruption along 

the rift zone to that time. Following the 1975 south flank 

earthquake, numerous ERZ intrusions and a few eruptions 

characterized the latter part of the 1970s. A series of intrusions 

occurred in early 1980, after which the ERZ was quiet until 

the start of eruptive activity in 1983 at Pu‘u ‘Ō‘ō (Holcomb, 

1987), which continues as of 2014. 

Like the SWRZ, the ERZ is connected to the south 

caldera magma reservoir at a depth of ~3 km. Such a depth is 

indicated by the location of seismicity associated with ERZ 

dike intrusions (Klein and others, 1987; Wolfe and others, 

1987). For instance, dikes that intruded into, and ultimately 

erupted from, the middle ERZ in June 2007 and March 2011 

both ascended from depths of about 3 km, according to 

earthquake locations and deformation modeling (Syracuse and 

others, 2010; Montgomery-Brown and others, 2010; Lundgren 

and others, 2013).

The ERZ contains a molten core that connects the summit 

to at least the distal subaerial end of the rift zone, over 50 km 

from the summit, and possibly into the submarine part of the 

rift zone beyond (Fiske and others, 1993). Such a continuous 

magma system was suggested by Dana (1849, description 

starting p. 188), based on second-hand reports of summit lava 

lake drawdown during the 1840 flank eruption. Johnson (1995b) 

proposed that the ERZ molten core, at 3–5-km depth, could 

feed magma vertically downward to deeper levels, towards 

the surface and eruption, and laterally along the rift zone. A 

hydraulically connected magma system along the rift zone is 
204 Characteristics of Hawaiian Volcanoes

Figure 19. Maps of seismicity along Kīlauea’s seismic Southwest Rift Zone (SWRZ). A, December 31, 1974–January 7, 1975. 

B, August 10–12, 1981. C, June 1982. D, March–September 2006. Color and size of circles gives depth and earthquake magnitude, 

respectively. Dashed red lines in part D indicate seismic SWRZ and volcanic SWRZ, and the Kamakai‘a Hills are labeled.

19°10'

19°15'

19°20'

19°25'

19°10'

19°15'

19°20'

19°25'

Earthquake

depth, in

kilometers

M3

M2

M1

Seismic SWRZ

Volcanic SWRZ

Kamakaiʻa

 Hills

A B

C D

Dec. 31, 1975–Jan. 7, 1976 August 10–12, 1981

June 1982 March–September 2006

0

0

2 4 KILOMETERS

1 2 MILES

 155°25' 155°20' 155°15' 155°25' 155°20' 155°15'

0

–5

–10
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 205

required to explain summit subsidence during ERZ eruptions, 

which indicates magma withdrawal from beneath the summit 

(Dvorak and Okamura, 1987). This is especially apparent when 

summit subsidence begins after the start of an ERZ eruption, as 

was the case on the lower ERZ during 1955 and 1960 (Eaton 

and Murata, 1960; Helz and Wright, 1992; Fiske and others, 

1993; Wright and Helz, 1996). Episodes of summit inflation 

also demonstrate hydraulic connectivity between the summit 

and lower ERZ. For example, in late 2003, as the summit 

began to inflate due to a surge in magma supply (see “2003–07 

Increase in Magma Supply to Kīlauea” section and Poland and 

others, 2012), uplift began in both the middle (fig. 5A) and 

possibly lower (fig. 5B) ERZ, indicating magma accumulation 

within the rift zone. During prolonged ERZ eruptions (including 

the Mauna Ulu and Pu‘u ‘Ō‘ō eruptions), magma might enter 

the volcano’s summit magma system and flow through the south 

caldera reservoir directly to the ERZ without spending time in 

storage (or displace a similar amount of stored magma into the 

ERZ), as suggested by gravity and deformation measurements 

from the 1980s through 2000s (Johnson, 1987; Kauahikaua and 

Miklius, 2003; Johnson and others, 2010).

Geophysical, petrologic, seismic, and physical evidence 

demonstrate that the ERZ contains multiple areas of shallow 

magma storage along its length. A reservoir beneath the 

Makaopuhi Crater area was first suggested by Jackson and 

others (1975) on the basis of seismicity and deformation 

associated with ERZ eruptions in 1968, and again by Swanson 

and others (1976b), using similar data from the 1969 ERZ 

eruption. Local gravity highs in the vicinity of Makaopuhi and 

Nāpau Craters suggest the presence of dense bodies—probably 

solidified magma storage zones—in these areas (Kauahikaua 

and others, 2000). High b-values (a seismic quantity associated 

with magma storage in volcanic environments) at multiple 

locations along the ERZ, including near Makaopuhi, also 

suggest magma storage (Wyss and others, 2001). Wolfe and 

others (1987) argued for a reservoir near Makaopuhi, because 

seismicity associated with the first episode of the Pu‘u ‘Ō‘ō 

eruption started in that location (a pattern that was repeated 

in March 2011; Lundgren and others, 2013), and episodes of 

lava fountaining during 1983–86 suggested that a secondary 

reservoir between the summit and Pu‘u ‘Ō‘ō acted as a valve 

that controlled eruptive activity. Owen and others (2000b) and 

Segall and others (2001) similarly found that including magma 

storage near Makaopuhi improved models of deformation 

from data recorded during a 22-hour eruption in Nāpau Crater 

in 1997, and Cervelli and others (2002) argued for a source of 

magma, possibly beneath Makaopuhi Crater, contributing to the 

1999 intrusion.

The somewhat evolved compositions of many lava 

flows erupted from the ERZ require storage at shallow 

levels within the rift zone (Wright and Fiske, 1971; 

Thornber and others, 2003). Likewise, compositions of lava 

from the early stages of the Pu‘u ‘Ō‘ō eruption resulted 

from mixing between isolated stored magmas and a more 

mafic end member (Garcia and others, 1989, 1992, 2000; 

Thornber, 2003). Lava flows from the lower ERZ in 1955 

and 1960 had a common source and show signs of mixing 

between evolved magmas stored within the rift zone (which 

crystallized pyroxenes and plagioclase as well as olivine) and 

more primitive magmas supplied from the summit (which 

crystallized only olivine; Helz and Wright, 1992; Wright 

and Helz, 1996). Drilling at a geothermal power plant in the 

lower ERZ in 2005 intersected a magma body with a dacitic 

composition at a depth of 2.5 km (Teplow and others, 2009). 

That such magma bodies exist is not surprising, given the 

frequent and sometimes long-lived activity along the ERZ 

(Parfitt, 1991); they were inferred from vent distributions 

and petrology prior to being drilled (Moore, 1983). The 

ongoing eruption at Pu‘u ‘Ō‘ō, for instance, has formed a 

small (~1×107 m3) magma storage area beneath the vent, 

as interpreted from geophysical, geochemical, and fluid 

dynamic evidence (Wilson and Head, 1988; Hoffmann and 

others, 1990; Garcia and others, 1992; Owen and others, 

2000b; Segall and others, 2001; Heliker and others, 2003; 

Thornber and others, 2003; Shamberger and Garcia, 2007; 

Mittelstaedt and Garcia, 2007). 

Volcanic Southwest Rift Zone

Holcomb (1987) suggested that the SWRZ included a 

northern strand that extended from Halema‘uma‘u Crater to 

5 km southwest of Maunaiki (fig. 3) and that was connected to 

the deeper magma plumbing system by way of Halema‘uma‘u. 

Fiske and others (1993) termed this the “classic” SWRZ and 

suggested that it was a superficial feature fed by shallow dikes 

intruded laterally from high levels of the summit magma 

system. We term this strand the “volcanic SWRZ” to avoid 

confusion with the “seismic SWRZ” (described in a section 

of the same name above). The volcanic SWRZ follows an 

alignment of fissures and eruptive vents, including the Great 

Crack (fig. 3), as far as the coast and was the source of post18th century eruptions in 1823, 1868, 1919–20, and 1971.

Instead of tapping the summit magma system at the level 

of the south caldera reservoir at ~3-km depth, as do the seismic 

SWRZ and ERZ, the volcanic SWRZ appears to be within 

1 km of the surface and is fed directly from the Halema‘uma‘u 

reservoir and, sometimes, from Halema‘uma‘u Crater itself. 

This connection was demonstrated by the 1919–20 Maunaiki 

(fig. 3) eruption, when the active lava lake at Halema‘uma‘u 

drained into a fissure that propagated southwest from the 

caldera. Jaggar (1919) observed lava flowing in cracks just 

beneath the surface, and occasionally reaching the surface, at 

multiple locations between Halema‘uma‘u Crater and Maunaiki, 

attesting to its shallow nature. When the summit lava lake 

drained in 1922 and 1924, a dike-like structure interpreted 

to be the 1919–20 conduit was exposed in the walls of 

Halema‘uma‘u Crater (fig. 20)—further evidence of its shallow 

connection to the summit magma system. This interpretation is 

at odds with the existence of higher gravity along the volcanic 

SWRZ, compared to the seismic SWRZ (Kauahikaua and 

others, 2000), implying a deeper and more important magma 

pathway. Gravity, however, reflects cumulative magma storage 
206 Characteristics of Hawaiian Volcanoes

throughout Kīlauea’s evolution. The current gravity field may 

indicate that the volcanic SWRZ was a more extensive magma 

pathway earlier in Kīlauea’s history than it currently is, much 

like gravity data that also suggest southward migration of the 

ERZ over time (Swanson and others, 1976a).

Duffield and others (1982) provided an overview of 

historical eruptive activity southwest of the summit caldera, 

noting that the eruptions of 1868, 1919–20, and 1971 were 

fed from lava lakes at Halema‘uma‘u Crater or magma 

stored at shallow depth beneath the crater (on the basis of 

drops in lava level or surface sagging coincident with the 

eruptions). The 1823 Keaīwa eruption issued from the Great 

Crack about 20 km from the summit and included exceptionally thin pāhoehoe (Stearns, 1926; Duffield and others, 

1982; Soule and others, 2004). The eruption was interpreted 

by Ellis (1825) to be contemporaneous with a drop in lava 

level at the summit, forming the “black ledge” that existed 

at the time of his visit to Kīlauea Caldera in 1823, just a few 

months after the Keaīwa outbreak.

Volcanic SWRZ eruptions differ considerably in terms of 

eruption rate, eruption style, vent location, and composition 

from the 1974 eruption, which had a source in the seismic 

SWRZ. These differences led some authors to propose that the 

1971 and 1974 eruptions southwest of the summit originated 

from different structural domains of the volcano (Duffield and 

others, 1982; Lockwood and others, 1999). Evidence from 

deformation suggests that dikes from both the seismic SWRZ 

and volcanic SWRZ are more or less vertical (Pollard and 

others, 1983; Dvorak, 1990; Lockwood and others, 1999); 

thus, the volcanic SWRZ eruptions are probably not a result of 

south-dipping dikes that originate from the seismic SWRZ and 

intersect the surface along the volcanic SWRZ trend.

Halema‘uma‘u–Kīlauea Iki Trend (HKIT)

A shallow magma pathway extends east from 

Halema‘uma‘u Crater towards Kīlauea Iki, defining what 

Hazlett (2002) termed the Halema‘uma‘u-Kīlauea Iki rift zone 

(his figure 31) and what we refer to as the Halema‘uma‘uKīlauea Iki Trend (HKIT). Like the volcanic SWRZ, the 

HKIT is connected to the shallow Halema‘uma‘u magma 

reservoir, transports magma within about 1 km of the surface, 

and is defined by an alignment of eruptive vents and fissures 

(fig. 21), although no accompanying gravity high is apparent 

(Kauahikaua and others, 2000). The western end of the 

HKIT is defined by frequent historical activity adjacent to 

Halema‘uma‘u Crater, with eruptions in 1954, 1971, 1975, and 

1982 (Holcomb, 1987; Neal and Lockwood, 2003). Farther 

east, the HKIT is marked by eruptions in 1832 on Byron 

Ledge (which separates Kīlauea Iki Crater from the deeper 

part of the caldera), in 1868 within Kīlauea Iki Crater, in 1877 

along the east side of the caldera, and in 1959 at Kīlauea Iki. 

The site of the ‘Ailā‘au eruption, which occurred in the 15th 

century and was ongoing for about 50–60 years (Clague and 

others, 1999; Swanson and others, 2012a), also lies along that 

trend immediately east of Kīlauea Iki, and lava flows from 

A

B

Fissure

Fissure

Lava

Figure 20. Photos of the southwest wall of Halema‘uma‘u Crater 

showing the fissure through which the 1919–20 Maunaiki eruption is 

thought to have been fed. A, USGS photograph by Thomas A. Jaggar, 

Jr., taken on May 25, 1922, following a collapse of Halema‘uma‘u Crater. 

A small amount of lava is present at the base of the fissure. B, USGS 

photograph by Howard A. Powers, taken on August 18, 1947, with the 

south flank of Mauna Loa in the background.

‘Ailā‘au would obscure any vents that might have formed 

prior to the 15th century outside the current caldera. The HKIT 

is therefore currently expressed only within the caldera, and it 

is not clear to what extent caldera formation (which occurred 

after the ‘Ailā‘au eruption; Swanson, 2008; Swanson and 

others, 2012a) might have influenced the geometry of existing 

or subsequent magma pathways or whether caldera-bounding 

faults exert control on eruptive vent locations.

Magma transported by way of the HKIT passes through, 

and may be stored within, the shallow Halema‘uma‘u magma 

reservoir before eruption, as demonstrated in 1959. The 

Kīlauea Iki eruption of that year was fed, at least in part, by 

a batch of rapidly ascending, primitive magma, as indicated 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 207

by a deep (~55 km) earthquake swarm three months prior to 

eruption and summit inflation that began two months before 

the eruption (Eaton and Murata, 1960). At least two lines of 

evidence suggest that some of this magma was stored beneath 

the caldera (and probably in the shallow Halema‘uma‘u 

magma reservoir) before erupting. First, deformation and 

seismicity indicate a source near Halema‘uma‘u Crater (Eaton 

and others, 1987; Wright and Klein, 2014), consistent with 

the location of the shallow reservoir. Second, petrographic 

and petrologic evidence from 1959 lava and scoria confirm 

that at least two components—a juvenile magma and a 

shallowly stored magma that is closely related to lava erupted 

in 1954—mixed before and during the eruption (Wright, 1973; 

Helz, 1987; Wright and Klein, 2014; Helz and others, this 

volume, chap. 6). Melt inclusions from Kīlauea Iki scoria also 

indicate mixing between a primitive, parental component and 

a component that had been stored at shallow depths (<3 km); 

the mixing probably occurred during summit inflation in the 

two months before the eruption (Anderson and Brown, 1993). 

The period of storage of the juvenile magma was sufficiently 

brief and deep that little volatile exsolution (except for CO2) 

19°24'

19°25'

 155°17' 155°16' 155°15'

9/71

9/71

9/74

1921

1975

4/82

1921

9/71

1975

1918–19

8/71

1954

8/71

1919

8/71 1892–94

1921 9/71

7/74

7/74

7/74

1892–94

1832

1888–89

1954

1892–94

4/82

8/71

1877

8/71

8/71

1790t 9/71

1975

1921 1954

1954

9/82

1882–85

1959

0

0

1 KILOMETER

0.5 MILES

0.5

0.25

7/74

8/71

1832

9/82

9/71

9/74

1959

1877

1882–85

1921

1919

1918–1919

4/82

1975

1954

1888–1889

1892–1894

1790

eruptive fissure

EXPLANATION

Figure 21. Geologic map showing post-18th century lava flows of the summit region (warm colors represent more recent flows than cool colors) 

and eruptive fissures (hatched red lines). Tephra deposits, like the 1959 Kīlauea Iki deposit, are not shown. From Neal and Lockwood (2003).

occurred before the eruption; gas was therefore available to 

drive high fountaining (Stovall and others, 2012). Lacking 

concurrent continuous deformation data, it is not possible to 

ascertain the relation between the 1959 eruption and the HKIT, 

but we speculate that magma erupted in 1959 utilized this 

pathway to travel from the storage reservoir to the surface.

Deep Rift Zones

Many models of Kīlauea’s magmatic system incorporate 

a deep (~3–9 km) rift zone below the ERZ and seismic SWRZ 

that is thought to contain a mix of magma and solidified 

intrusions. This deep rift zone provides a pathway by which 

magma may enter the ERZ and seismic SWRZ without 

passing through summit reservoirs (for example, Ryan, 1988), 

as well as internal pressure to force the volcano’s south flank 

seaward (for example, Dieterich, 1988; Delaney and others, 

1990; Fiske and others, 1993; Cayol and others, 2000; Wright 

and Klein, 2014). Petrologic evidence has been used to argue 

that the deep rift zone provides a path for primitive magmas 

to circumvent the summit reservoir system prior to eruption 
208 Characteristics of Hawaiian Volcanoes

along the ERZ. For example, Trusdell (1991) suggested that 

picritic (19 percent MgO) magma, which erupted from fissures 

along the lower ERZ in 1840 (a mixed magma was extruded 

as part of the same eruption in the middle ERZ), bypassed the 

summit magma reservoir by way of the deep rift zone. Vinet 

and Higgins (2010) cited the presence of deformed olivine 

in some Mauna Ulu lava flows as evidence that magma had 

moved along the basal décollement to erupt on the ERZ 

without passing through the summit reservoir. Wright and 

Helz (1996) proposed that compositional variations in lavas 

erupted during the 1960 Kapoho eruption reflected magma 

flow along two different paths from the summit region to 

the lower ERZ, with various amounts of mixing between 

magmatic components.

Fundamentally, the petrologic arguments above are 

based on the inference that the eruption of olivine-rich lavas 

on the ERZ would not be possible if the magma had passed 

through the summit reservoir system, or even the shallow 

part of the rift zone, where it would have mixed with resident 

magma before erupting. This inference is not necessarily 

accurate. Most primitive magmas erupted from the ERZ 

have complex evolutions, having fractionated and (or) mixed 

with stored magmas at the base of Kīlauea’s summit magma 

system (for example, Helz and others, this volume, chap. 

6). Olivine cumulates grow in that area, are available to be 

picked up by magma passing through the summit and out 

to the ERZ, and are deformed by flow through the magma 

plumbing system (Clague and others, 1995). Involvement of 

the deep rift zone is therefore not necessary for production 

of an olivine-rich flow with deformed crystals. Indeed, 

Macdonald (1944) proposed such a model for the 1840 

eruption of Kīlauea, with picritic lava extruded from the 

lower ERZ indicating an origin in the deeper part of the 

summit reservoir system, while olivine-poor lava that 

erupted from the middle ERZ came from higher levels of the 

summit magma complex. Even high-MgO Mauna Ulu lava, 

thought by Vinet and Higgins (2010) to indicate a source 

deep beneath Kīlauea’s ERZ, has olivine compositions that 

are consistent with a source no deeper than the base of the 

summit reservoir (Helz and others, this volume, chap. 6). 

Melts with high (>14 percent by weight) MgO have been 

recovered on the Puna Ridge (the submarine extension of 

the ERZ) and provide evidence that high-MgO magma 

can occasionally traverse long distances through Kīlauea’s 

plumbing system, perhaps related to periods of heightened 

magma supply (Clague and others, 1991). Most high-MgO 

lava on the Puna Ridge, however, show evidence of mixing 

between volatile-rich and shallowly degassed end members 

and subsequent fractional crystallization (for example, Dixon 

and others, 1991), indicative of complex evolution; similar 

picritic lava is also found on the submarine southward 

extension of Mauna Loa’s Southwest Rift Zone (Helz and 

others, this volume, chap. 6).

Geophysical evidence for melt in the deep rift zones 

of Kīlauea is ambiguous. The rift zones are largely aseismic 

below 4 km, which supports the presence of melt and (or) 

hot cumulates (Ryan, 1988; Clague and Denlinger, 1994; 

Denlinger and Okubo, 1995). High b-values in the deep 

ERZ suggest melt, although the high values extend several 

kilometers south of the rift zone into areas unlikely to host 

magma (Wyss and others, 2001). In addition, models of 

deformation data require opening of the rift zones below 

about 3 km and above the basal décollement (which defines 

the base of the volcano’s mobile south flank) at ~9-km depth 

(Dieterich, 1988; Delaney and others, 1990; Owen and 

others, 1995, 2000a; Cayol and others, 2000; Denlinger and 

Morgan, this volume, chap. 4). This opening is consistent with 

magma accumulation and deep rift pressurization—a potential 

mechanism for seaward motion of Kīlauea’s south flank (for 

example, Dieterich, 1988; Delaney and others, 1990; Cayol 

and others, 2000; Wright and Klein, 2014).

Gravity and seismic-velocity data contradict evidence 

for magma in the deep rift zones. Kīlauea’s rift zones, 

particularly the ERZ, are characterized by positive residual 

gravity anomalies that imply dense intrusions extending 

from the base of the volcanic pile to high within the edifice 

(Kauahikaua and others, 2000). Similarly, seismic velocities 

within the deep ERZ are elevated, relative to adjacent areas, 

suggesting gabbro-ultramafic cumulates (Okubo and others, 

1997; Park and others, 2007, 2009). Low Vp/Vs ratios are also 

inconsistent with the presence of melt (Hansen and others, 

2004). This dense, seismically fast region has been proposed 

to be a hot cumulate core that is capable of plastic flow and 

might drive south flank motion by gravitational forces (for 

example, Clague and Denlinger, 1994; Park and others, 2007, 

2009; Plattner and others, 2013; Denlinger and Morgan, this 

volume, chap. 4). Magnetic data support such an interpretation, 

as low magnetization at depth below the ERZ could be a result 

of hot, and possibly altered, rock, as opposed to solidified, 

more magnetic intrusions in the shallower part of the rift zone 

(Hildenbrand and others, 1993). 

Given the conflicting petrologic and geophysical interpretations, the pressing questions become, is molten magma 

present in Kīlauea’s deep rift zones, and does the deep rift 

zone provide a pathway for magma to bypass the summit 

reservoir? These questions can be addressed by three possible models for deep rift zone structure: (1) the deep rift 

zones are solid and do not allow for magma transport; (2) the 

deep rift zones are a mix of cumulates and melt and provide 

a means for magma to bypass the summit reservoir; and (3) 

the deep rift zones are a mix of cumulates and melt but are 

fed by downward draining of magma from a molten core that 

originates from Kīlauea’s ~3-km-depth summit magma storage area. Of these models, only the third satisfies geophysical and geological constraints.

While model 1 explains the dense, seismically fast 

structure of the deep rift zone, it is not consistent with the 

model for ERZ development. The asymmetric ERZ positive 

gravity anomaly (with a steeper gradient on the north than 

on the south) suggests southward migration of the ERZ over 

time (Swanson and others, 1976a; Kauahikaua and others, 

2000). The maximum residual gravity is near the current 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 209

ERZ axis, indicating a complex of solidified magma that 

extends well beneath the transport region of the present rift 

zones (Kauahikaua and others, 2000). If the ERZ migrated 

southward over time but no magma were present below 

3-km depth, how can solidified intrusions extend to the 

base of the volcano under the present ERZ? At some point, 

magma must have been present in the deep rift zones, and the 

solidification of that magma provides the dense core imaged 

by gravity and seismic measurements.

Model 2 is compatible with a dense and seismically fast 

deep rift zone (implying solidified magma and cumulates) 

that is also aseismic, opening, and has low magnetization 

(implying the presence of melt, or at least hot and altered 

rock), but is unsupported by observations from recent eruptions and intrusions. If a semimolten deep rift zone provides 

a pathway for magma to bypass Kīlauea’s summit magma 

storage areas, why was there no change in the rate of deep 

rift zone opening during the 2003–07 surge in magma supply 

(fig. 6)? Similarly, why was there no contraction of the deep 

rift zone during the 1997 fissure eruption in Nāpau Crater, 

which was caused by extensional stress from deep rift opening and would presumably have received magma from the 

deep rift zone (Owen and others, 2000b)?

Model 3 is essentially the same as the molten core 

model of Johnson (1995b) and Fiske and others (1993). 

In model 3, the ERZ has a semicontinuous molten core at 

3–4-km depth from which magma moves vertically, either 

upward towards the surface or downward, into the deeper 

part of the rift zone, and all magma that enters the ERZ 

passes through the summit reservoir system first. The ERZ 

is characterized by an interconnected melt zone for at least 

its entire subaerial length (figs. 5, 10). Olivine-rich magma 

that erupts on the ERZ does not bypass the summit reservoir 

but, instead, traverses tens of kilometers, mixing with older 

batches of rift-stored magma along the way (for example, 

Wright and Fiske, 1971; Thornber and others, 2003). The 

Father’s Day dike intrusion of June 2007 provided a glimpse 

of the MgO-rich magma fed into Kīlauea’s summit reservoir 

system. The dike was driven by summit overpressure, and the 

lava that erupted when the dike reached the surface between 

the summit and Pu‘u ‘Ō‘ō (after propagating at a depth no 

greater than 3 km; Montgomery-Brown and others, 2010) 

was much hotter and richer in MgO than lava that had been 

erupting from Pu‘u ‘Ō‘ō a few days earlier (fig. 4D; Poland 

and others, 2012). Similarly, the composition of lava extruded 

three weeks after the start of the 1960 Kapoho eruption was 

relatively olivine-rich and contained up to 30 percent of 

the juvenile component of the 1959 Kīlauea Iki eruption, 

suggesting that 1959 summit magma was able to reach an 

eruption site about 50-km distant in a few weeks (Wright and 

Helz, 1996). Finally, if all magma passes through the summit 

reservoir system, it ascends through cumulates at the base of 

the summit reservoir complex (Clague and Denlinger, 1994; 

Clague and others, 1995; Kauahikaua and others, 2000) and 

may pick up deformed olivine crystals that might later erupt 

at the summit (for example, Helz, 1987) along the subaerial 

ERZ (for example, Wright and Helz, 1996; Vinet and Higgins, 

2010) or even the submarine extension of the ERZ (Clague 

and others, 1995), although submarine olivine compositions 

are more magnesian than those of the summit and subaerial 

ERZ (Helz and others, this volume, chap. 6). Observing 

the seismicity and deformation associated with any future 

submarine eruptions on the Puna Ridge will be invaluable in 

interpreting the relation between the summit magma system 

and submarine part of the rift zone.

Downward transport of magma from the molten core 

explains geophysical measurements that imply the presence 

of both melt and cumulates extending to the base of the 

volcano beneath the rift zones. Magma that passed though 

the summit and intruded along the ERZ but did not erupt 

would gradually crystallize, with dense minerals retained 

at depth to create a massive, high-velocity body. Degassed 

magma, some of which may have drained back into eruptive 

fissures following extrusion (which is commonly observed 

in Hawai‘i—see, for example, Richter and others, 1970), 

might also sink to lower levels. The deep rift zone would 

therefore contain a mix of degassed melt and dense crystals 

that could explain deep opening modeled from deformation 

data, the lack of seismicity, low magnetization, high seismic 

velocities, and positive gravity.

The deep rift zone is ambiguous in character, yet critical 

to overall models of Kīlauea’s magma supply and storage 

system. For example, whether all magma passes through the 

summit storage network or intrudes into the deep rift zone 

and bypasses the summit has important implications for 

magma supply calculations and CO2 degassing (see “Magma 

Supply” section above). Further work is obviously necessary 

to investigate the characteristics and evolution of the deep 

rift zones on Hawaiian volcanoes, but we are hopeful 

that this model will serve as a starting point for studies of 

Kīlauea’s past, present, and future eruptive activity.

Model Summary and Complications

We propose a refined model of Kīlauea’s magma 

plumbing system that includes multiple areas of magma 

storage beneath the summit, with the primary storage 

reservoir at ~3–5 km beneath the south caldera and a smaller 

reservoir ~1–2 km beneath the east margin of Halema‘uma‘u 

Crater. Magma is also at least occasionally stored beneath 

Keanakāko‘i Crater. Kīlauea’s major rift zones radiate east 

and southwest from the deeper south caldera reservoir, while 

shallower pathways extend from the Halema‘uma‘u reservoir 

east toward Kīlauea Iki and southwest toward, and beyond, 

Maunaiki. 

Individual elements of this model are long established, 

having been proposed previously by other authors (for 

example, Holcomb, 1987; Klein and others, 1987; Fiske and 

others 1993), but our overall depiction represents an attempt to 

combine geologic, geochemical, and geophysical observations, 

especially those collected by seismic and geodetic techniques 

since the 1990s, into a comprehensive model for Kīlauea. 
210 Characteristics of Hawaiian Volcanoes

The model nevertheless remains an oversimplification. For 

example, we overlook the geometrical complexity of the ERZ, 

particularly the bend that occurs southeast of the summit that 

may have formed as a result of southward rift zone migration 

over time (Swanson and others, 1976a). We chose, instead, to 

focus on the general configuration of Kīlauea’s magma plumbing system as defined geophysically, leaving these and other 

complications for future study.

We have also neglected important structural elements of 

the volcano, most notably the Koa‘e Fault System (fig. 10), 

which is itself poorly understood (Duffield, 1975; Swanson 

and others, 1976a; Fiske and Swanson 1992; Lockwood and 

others, 1999) but was recognized by Holcomb (1987) as being 

associated with a “southern strand” of the SWRZ. Magma 

was inferred to have intruded the Koa‘e Fault System in 1965 

(Fiske and Koyanagi, 1968) and 1999 (Cervelli and others, 

2002), and intrusions were confirmed in 1969 (Swanson and 

others, 1976b) and 1973 (Zablocki, 1978; Tilling and others, 

1987), with the 1973 dike extending several kilometers into 

the central Koa‘e Fault System from the ERZ, based on 

leveling data (Swanson and others, 2012b). Dikes underlying 

the Koa‘e Fault System have also been inferred from a 

variety of geophysical measurements (for example, Flanigan 

and Long, 1987). Small eruption sites of unknown age, 

but covering ash deposited in 1790, were discovered in the 

central part of the Koa‘e by Swanson and others (2012b), 

demonstrating that these dikes sometimes erupt. The Koa‘e 

Fault System may therefore structurally and magmatically 

link the ERZ and seismic SWRZ in a single “breakaway” 

rift (Fiske and Swanson, 1992) and may one day become 

an important site for the injection and storage of magma 

(Fiske and Swanson, 1992; Lockwood and others, 1999). 

Geophysical monitoring of future intrusions and eruptions 

in the Koa‘e Fault System will be critical in elucidating the 

role and character of the region, as well as incorporating its 

structure into the magmatic model of Kīlauea.

Mauna Loa

Mauna Loa’s magmatic system is less well known than 

Kīlauea’s, owing to a comparative lack of geophysically 

monitored eruptive activity. The modern seismic network 

on Mauna Loa was established in the 1950s (Okubo, 1995; 

Okubo and others, this volume, chap. 2), and deformation 

monitoring began in the 1960s (Decker and others, 1983). 

These measurements, however, missed the period of frequent 

Mauna Loa activity—31 eruptions occurred during 1843–1950 

(Lockwood and Lipman, 1987), but only two eruptions took 

place from 1951 to 2014 (in 1975 and 1984). In addition, while 

geodetic monitoring characterized important changes over 

time (fig. 22), the network was too sparse and the deformation 

insufficient in magnitude to facilitate anything except 

generalized modeling. Despite these shortcomings (compared 

to Kīlauea), enough measurements exist to provide a rough 

picture of Mauna Loa’s magma plumbing system.

Both deformation and seismicity suggest the existence 

of a zone of magma storage at 3–4 km beneath the southeast 

margin of Moku‘āweoweo Caldera. A magma accumulation 

zone is hypothesized to correspond to an aseismic region that 

is capped by earthquakes at about 4-km depth beneath the 

south-southeastern part of the caldera (Decker and others, 

1983; Okubo, 1995). Vertical deformation from leveling 

during 1977–81 was best modeled by a source of inflation 

at 3.1-km depth below the southern part of the caldera. 

Combining vertical, horizontal, and tilt over the same time 

period suggests a depth of almost 4 km (Decker and others, 

1983; Lockwood and others, 1987). The 1984 eruption was 

accompanied by summit deflation as a dike propagated into, 

and erupted from, the Northeast Rift Zone. Tilt and leveling 

data spanning the eruption were modeled by a source of 

volume loss at about 3.5-km depth beneath the southeast 

margin of the caldera, coupled with dike opening from 0–5-km 

depth in the Northeast Rift Zone (Johnson, 1995a). Following 

the eruption, Electronic Distance Mesurement (EDM) and 

leveling data indicated inflation with a model location just east 

of the caldera at 3.7-km depth (Miklius and others, 1995). 

An episode of inflation began in May 2002 and was well 

monitored by both continuous GPS (Miklius and Cervelli, 2003) 

and InSAR (Amelung and others, 2007). InSAR data spanning 

2002–05 are best fit by magma accumulation in two sources: a 

spherical body located 4.7 km beneath the southeast margin of 

the caldera and a dike-like structure extending the length of the 

caldera and into both rift zones, with most opening occurring at 

4–8-km depth (Amelung and others, 2007). GPS displacements 

from 2004 to 2005, the period of most rapid inflation, can be 

modeled by similar sources and demonstrate the importance 

of the dike to the overall inflation pattern and rate (fig. 23). It 

is unknown whether the dike source is a new feature associated with the 2002 onset of inflation or was previously present 

but unresolvable. Given the limited spatial resolution of the 

campaign-style measurements (including tilt, leveling, EDM, 

and GPS) used before the availability of InSAR and continuous 

GPS, we suspect that the dike source has always been present. 

The fact that both sources accumulated magma implies that they 

form a single interconnected magma storage system. During 

the latter half of the inflation episode (which lasted until 2009), 

GPS data indicated that the source beneath the southeast caldera 

was the dominant region of magma storage, suggesting that the 

dike source may be most active during periods of rapid supply 

(which was apparently the case during 2004–05).

In addition to regions of magma storage, deformation 

of Mauna Loa also indicates motion of the volcano’s southeast flank. With the establishment of campaign GPS stations 

around the volcano in the 1990s, surface displacements on 

Mauna Loa could be measured on a broader scale than was 

previously possible. These measurements revealed that the 

southeast flank of Mauna Loa was moving southeast at a 

rate of about 4 cm/yr (Miklius and others, 1995). Similar to 

Kīlauea’s, such motion may reflect slip along a deep subhorizontal fault underlying the volcano’s flank. The role of 

this flank motion in magmatic and tectonic activity at Mauna 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 211

Figure 22. Plot of distance change across Mauna Loa’s summit caldera during 1974–2011. Blue dots give Electronic Distance Measurement (EDM) 

results, black dots are campaign Global Positioning System (GPS) data, and black line is from continuous GPS. Positive change is interpreted as 

inflation and negative, deflation. Inset map shows station locations and names relative to Mauna Loa’s caldera.

Figure 23. Map of displacements from Global Positioning System 

data from Mauna Loa during 2004–05. Observed displacements are 

black, with 95-percent confidence ellipses, and modeled displacements 

are blue. Model includes a point source at 3.5-km depth beneath the 

southeast caldera rim (red dot), with a volume increase of 4×106 m3/yr, 

and a vertical dike along the length of the caldera (red line), with a top at 

3.8-km depth (the bottom is difficult to resolve but is probably no more 

than a few km below the top) and a volume increase of 30×106 m3/yr. 

Loa is unclear, although stress modeling suggests a feedback 

where flank earthquakes encourage dike intrusion, and vice 

versa (Walter and Amelung, 2004, 2006).

Mauna Loa eruptions occur not only from the summit and 

rift zones, but also from submarine and subaerial vents oriented 

radially from the summit on the northwest flank of the volcano 

(Lockwood and Lipman, 1987). The radial vents are located in 

the area of greatest horizontal tension, where the volcano’s stress 

field is controlled by intrusions into the two rift zones, suggesting 

that radial dike eruptions occur during periods of heightened 

magma pressure within the volcano (Rubin, 1990). Petrologic 

studies of radial vent eruptions have found both primitive 

high-Mg lava (Riker and others, 2009), as well as evolved alkalic 

lava (Wanless and others, 2006). If the source magmas for these 

eruptions ascended through the main Mauna Loa magmatic 

system, their compositional signatures would probably have 

been overprinted by the more common tholeiitic magma that 

undoubtedly dominates the magma plumbing system. Radial 

vent eruptions may therefore be evidence of secondary magma 

pathways that bypass the summit plumbing system, allowing 

evolved magmas to reach the surface (Wanless and others 2006), 

and (or) evidence of episodic increases in magma supply that 

allow primitive magmas to erupt (Riker and others, 2009). Eruption

1975 1980 1985 1990 1995 2000 2005 2010

20

 0

40

60

80

100

120

Distance change, in centimeters

EDM

GPS

Eruption

MOKP

MLSP

Year

 155°40' 155°30'

19°20'

19°30'

0

0

2 4 KILOMETERS

1 2 MILES

2 cm/yr
212 Characteristics of Hawaiian Volcanoes

Summary

Frequent eruptive activity and an excellent record of 

geophysical and geological monitoring during the period 1950–

2014 have provided a high-resolution view of Kīlauea’s magma 

plumbing system. Our model of Kīlauea’s magma storage 

areas and transport pathways builds upon previous models (for 

instance, Eaton and Murata, 1960; Decker and others, 1987; 

Holcomb, 1987; Klein and others, 1987; Ryan, 1988; Fiske and 

others, 1993; Tilling and Dvorak, 1993) and incorporates data 

collected by new techniques (for example, continuous GPS and 

InSAR). The improved resolution of the different components 

of Kīlauea’s plumbing system provided by our conceptual 

model offers a new framework for interpreting volcanic and 

tectonic activity at the volcano. Although less active during the 

same time period, the general outline of Mauna Loa’s magmatic 

system should prove valuable for understanding future volcanic 

and seismic activity at that volcano.

Curiously, both Kīlauea and Mauna Loa are characterized 

by primary magma storage areas located about 3–4 km beneath 

their summit calderas. This correspondence suggests that the 

magmatic systems of Hawaiian volcanoes migrate upward as 

the volcanoes grow (Decker and others, 1983; Lockwood and 

others, 1987; Dvorak and Dzurisin, 1997), and that the 3–4-km 

depth represents a favorable level of magma accumulation, 

possibly because it is a level of neutral buoyancy (Ryan, 

1987; Burov and others, 2003). Magma storage zones exist at 

similar depths at other basaltic shields—for example, Piton 

de la Fournaise (Peltier and others, 2009), Etna (Aloisi and 

others, 2011), and Axial seamount (Chadwick and others, 1999; 

Nooner and Chadwick, 2009)—implying that levels of neutral 

buoyancy and magma trapping may be comparable at large 

basaltic volcanoes worldwide. Such volcanoes are also often 

characterized by multiple, vertically stacked magma reservoirs 

and rift zones that radiate from the summit, resembling the 

model of Kīlauea’s magma system.

Magma Transport

At basaltic volcanoes, transport of magma away from a 

source reservoir is most often accomplished by means of dikes and 

sills. Intrusion initiation, propagation, and eruption have been the 

subjects of a large body of literature, a complete (or even partial) 

review of which is impractical. We therefore refer the reader 

to articles cited within this section for additional information, 

although these few articles only scratch the surface of research 

into the process of intrusion. We also focus specifically on dike 

intrusion, which is common at Hawaiian volcanoes, although 

we note that other basaltic systems seem to be dominated by sill 

intrusion over dikes (for example, Bagnardi and others, 2013).

Dike propagation occurs both vertically and laterally 

and is controlled by numerous factors, the most important of 

which include magma pressure, the preexisting stress field, 

magma viscosity, and host rock properties (for example, 

Rubin, 1995; Taisne and Jaupart, 2009; Traversa and others, 

2010). When dike driving pressure (defined as the difference 

between the magma pressure and the least compressive stress 

in the crust) is sufficiently high, the dike will reach the surface 

and erupt; however, geologic evidence suggests that most 

dikes are arrested before they reach the surface, probably 

because mechanical and stress barriers inhibit propagation 

(for example, Gudmundsson and others, 1999; Gudmundsson, 

2002; Rivalta and others, 2005; Taisne and Tait, 2009; Taisne 

and others, 2011a; Geshi and others, 2012). Understanding the 

conditions that favor dike eruption versus arrest is therefore 

critical for volcano monitoring and hazards mitigation (for 

example, Geshi and others, 2010).

Dike emplacement is accompanied by both deformation 

and seismicity. For example, earthquakes indicated lateral 

emplacement of dikes at Krafla, Iceland (Brandsdóttir and 

Einarsson, 1979; Einarsson and Brandsdóttir, 1980), and at 

Miyakejima, Japan (Toda and others, 2002), as well as vertical 

ascent of magma at Piton de la Fournaise (Battaglia and others, 

2005). Seismicity and tilt were used to follow the transition from 

vertical to lateral propagation of dikes at Etna (Aloisi and others, 

2006) and Piton de la Fournaise (Toutain and others, 1992; 

Peltier and others, 2005). Rates of dike propagation in these 

and similar cases are on the order of 0.1–1 km/hr, occasionally 

reaching several kilometers per hour (for example, Dvorak and 

Dzurisin, 1997; Grandin and others, 2011; Wright and others, 

2012) and sometimes displaying a complex propagation history 

in both rate and direction (for example, Taisne and others, 

2011b). Traversa and Grasso (2009) and Traversa and others 

(2010) noted that earthquakes may not necessarily track the 

propagation of the dike tip. In some cases, however, earthquakes 

must indicate the location of a dike’s leading edge—in Iceland 

in 1977, a small eruption occurred through a drill hole when 

seismicity associated with a propagating dike reached the hole 

(Brandsdóttir and Einarsson, 1979).

Exposures within the cores of eroded Hawaiian islands, 

such as O‘ahu, illustrate the importance of dike intrusions to 

volcano construction and evolution (Walker, 1986, 1987), and 

many of the primary characteristics of active dike emplacement 

were deduced from Hawaiian examples (for instance, Pollard 

and others, 1983). Exceptional monitoring data and frequent 

episodes of dike emplacement, especially at Kīlauea, provide 

views into the process of dike formation and intrusion (for 

example, Segall and others, 2001; Rivalta, 2010). Flow of 

magma through conduits can also be studied, thanks to longlived eruptive activity (for example, Cervelli and Miklius, 2003). 

Here, we summarize observations of active magma transport in 

Hawai‘i and explore the mechanisms for the formation of dikes 

within Hawaiian volcanoes.

Characteristics of Magma Transport at Hawaiian 

Volcanoes

Dike emplacement is the most important means of 

transporting magma from source reservoirs to intrusion and 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 213

eruption sites at shield-stage Hawaiian volcanoes. The eroded 

Ko‘olau volcano, on the Island of O‘ahu, displays thousands 

of dikes that, at outcrop or map scale, make up 50–65 percent 

of the host rock. These dikes form highly concentrated 

swarms, similar to oceanic sheeted dike complexes, that 

delineate rift zones in which intrusion is concentrated 

(Walker, 1986, 1987), with Hawaiian rift zones apparently 

formed and maintained by gravitational forces (for example, 

Fiske and Jackson, 1972). 

Compared to dikes exposed by erosion in Iceland— 

another region of frequent basaltic intrusions—those of Hawai‘i 

are thin. This observation led Rubin (1990) to suggest that 

Hawaiian dikes are driven by higher magma pressures than 

their Icelandic counterparts (which are more heavily influenced 

by tectonic stress), and therefore are more likely to reach the 

surface—a model later supported by Segall and others (2001). 

During 1960–75, dikes reached the surface at Kīlauea about 

twice as often as they stalled (Dzurisin and others, 1984)—a 

proportion much greater than that implied by geologic studies 

in Iceland (for example, Gudmundsson and others, 1999). 

Changes in the stress state of the volcano directly influence the 

proportion of magma that is stored versus erupted. For example, 

in the months following the M7.7 earthquake beneath Kīlauea’s 

south flank in 1975 (Nettles and Ekström, 2004; Owen and 

Bürgmann, 2006), the ratio of eruptions to noneruptive 

intrusions changed from 4:1 to 1:4, presumably because dilation 

of the volcano’s magma system caused by the earthquake 

promoted magma storage over eruption (Dzurisin and others, 

1980, 1984; Cayol and others, 2000). A similar relation was 

found at Piton de la Fournaise, which experienced a period 

dominated by noneruptive intrusions following a caldera 

collapse in 2007 (Peltier and others, 2010).

Rates of dike propagation at Kīlauea and Mauna Loa are 

comparable to those determined at other basaltic volcanoes. At 

Mauna Loa, lateral dike propagation associated with the 1984 

Northeast Rift Zone eruption occurred at a rate of 1.2 km/hr, 

based on the appearance of eruptive fissures (Lockwood and 

others, 1987). The relative lack of seismicity during the dike 

propagation was attributed by Lockwood and others (1987) to 

magma flow into a mostly open conduit created by an intrusion 

in 1975 (which immediately followed the summit eruption of 

that year) or perhaps to a stress drop caused by a nearby M6.6 

earthquake in 1983. In contrast, Kīlauea has experienced tens of 

intrusive episodes since the start of dense geophysical monitoring 

in the 1950 and 1960s. Klein and others (1987) provided a 

comprehensive overview of Kīlauea seismicity during 1963–83, 

subdividing earthquake swarms into those that were associated 

with magma intrusion and eruption, inflation of the magma 

system, and constant flow of magma from the summit into the rift 

zones. Seismic swarms due to rapid magma intrusion or eruption 

have migration speeds ranging from 0.1 to 6 km/hr. Propagation 

directions are generally away from the summit, although uprift 

migrations of seismicity have been observed within the ERZ 

between the summit and Mauna Ulu, possibly reflecting a 

complex interplay between varying stresses, pressure gradients, 

and magma conduits (Klein and others, 1987).

Numerous magmatic intrusions have occurred in 

Kīlauea’s ERZ since 1983, despite nearly continuous eruptive 

activity (ongoing as of 2014) and the existence of a magma 

conduit between the summit and ERZ eruption site (fig. 

24; see “Mechanisms of Dike Emplacement at Hawaiian 

Volcanoes” section for a discussion of intrusion driving 

forces). Five intrusions into the upper ERZ took place during 

1990–93, some of which were associated with pauses as long 

as 8 days in the ongoing eruption (Heliker and Mattox, 2003). 

A fissure eruption in Nāpau Crater in 1997 caused collapse of 

Pu‘u ‘Ō‘ō and a pause of 24 days (Owen and others, 2000b; 

Heliker and Mattox, 2003; Thornber and others, 2003), 

and dikes in 1999 and 2000 intruded the ERZ but failed to 

erupt (Cervelli and others, 2002; Heliker and Mattox, 2003), 

although they caused the ongoing eruption to pause for 11 

days and 7 hours, respectively. A dike consisting of two 

discrete segments intruded between Mauna Ulu and Nāpau 

Crater and resulted in a small eruption on the northeast flank 

of Kanenuiohamo in June 2007 (Poland and others, 2008; 

Montgomery-Brown and others, 2010; Fee and others, 2011). 

The intrusion propagated downrift discontinuously in a series 

of pulses, as indicated by bursts of seismicity coincident with 

increases in the rate of summit deflation, with periods of 

propagation characterized by rates of about 3 km/hr (fig. 25). 

During March 2011, the ongoing ERZ eruption was again 

interrupted, this time by a 5-day fissure eruption between 

Nāpau Crater and Pu‘u ‘Ō‘ō (Lundgren and others, 2013). 

These dikes followed the trends established by previous 

intrusions at Kīlauea. The 1997 and 2011 fissure eruptions were 

preceded by only a few hours of intense and localized seismic 

activity, suggesting that a dike propagated from the existing ERZ 

magma conduit at ~3-km depth to the surface at an average rate 

of about 1 km/hr (Owen and others, 2000b; Thornber and others, 

2003). Tilt records over the 1999 intrusion suggest a similar 

velocity in a downrift direction (Cervelli and others, 2002).

In addition to hosting repeated dike intrusions, Kīlauea 

provides an opportunity to study magma flow through established conduits, thanks to the occurrence of long-term eruptive 

activity. Such a conduit has existed since the start of the 1983–

present (as of 2014) ERZ eruption and has evolved over time 

from an episodically active to a continuously active pathway, 

possibly due to the relation between cooling of magma within 

the conduit and pressure in the summit magma reservoir system 

(Parfitt and Wilson, 1994). The primary vent for the eruption, 

Pu‘u ‘Ō‘ō, is located about 20 km from Halema‘uma‘u Crater 

along the curving trend of the ERZ. Given that the average 

propagation rate of a dike at Kīlauea is about 1 km/hr (Klein 

and others, 1987), one can reasonably infer that the transport 

rate of magma along an open conduit is no slower than that, 

and that magma would take no more than 20 hours to travel 

from the summit to Pu‘u ‘Ō‘ō. In fact, the rate of transport is 

possibly much higher, as suggested by ground tilt at the summit 

and at Pu‘u ‘Ō‘ō (Cervelli and Miklius, 2003).

Transient tilt events consisting of deflation-inflationdeflation (DID) or deflation-inflation (DI) cycles (fig. 26) have 

been recorded at Kīlauea since the installation of borehole 
214 Characteristics of Hawaiian Volcanoes

Figure 24. Map of Kīlauea’s East Rift Zone showing intrusive activity during 1990–2011. Seismicity (colored circles) is shown for intrusions 

for which little or no geodetic data are available. Model geometries (colored lines) are given for those dikes that were detected by deformation 

measurements (tilt, Global Positioning System data, and (or) Interferometric Synthetic Aperature Radar), including the 1997 Nāpau fissure eruption 

(Owen and others, 2000b), 1999 noneruptive dike (Cervelli and others, 2002), 2007 Father’s Day intrusion/eruption (Montgomery-Brown and others, 

2010), and 2011 Kamoamoa fissure eruption (Lundgren and others, 2013).

-20

-10

0

10

20

30

0 10 20 30 40 50 60 70

155.26

155.24

155.22

155.20

155.18

155.16

155.14

07:30 08:00 08:30 09:00

155.22

155.21

155.20

155.19

155.18

155.17

155.16

155.15

Tilt, in microraidans Longitude, in degrees

Time, in hours since midnight June 17, 2007, HST

1 km

3.2 km/hr

Local time (HST), June 17, 2007

 155.25° 155.16°

19.33°

19.42°

M3

M2

M1

0

0

2 4 KILOMETERS

1 2 MILES

Figure 25. Plots of tilt and 

seismicity associated with Kīlauea’s 

June 17–19, 2007, East Rift Zone 

intrusion and eruption. Tilt time 

series (top) is from a station located 

about 300 m west-northwest of the 

Hawaiian Volcano Observatory. 

Tilt azimuth is 327°, which is 

approximately radial to the source 

location, implying that positive tilt is 

inflation and negative tilt is deflation 

(assuming a spherical source). 

Seismicity (middle) occurred along 

an approximately east-west trend, 

so longitude is roughly perpendicular 

to the strike of the dike. Map 

shows distribution of earthquakes. 

Earthquake colors on map and 

time series indicate timing (blue, 

00:00–07:30 HST on June 17; red, 

07:30–09:00 on June 17; green, 

09:30 on June 17 to the end of June 

19). The most rapid propagation, 

during 07:30–09:00 HST on June 17 

(zoomed plot at bottom), occurred 

at a rate of about 3.2 km/hr (the 

two earthquakes at the end of the 

zoomed plot do not significantly 

influence the trend line). 

155°20' 155°15' 155°10' 155°05'

19°20'

19°25'

M3

M2

M1

Jan. 1997

Sept. 1999

June 2007

March 2011

Feb. 1993

March 1992

Aug. 1991

March 1991

Dec. 1990

Feb. 2000

Kīlauea

Caldera

Puʻu ʻŌʻō

Nāpau Crater

0

0

2 4 KILOMETERS

1 2 MILES
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 215

tiltmeters in 1999 (Cervelli and Miklius, 2003; Anderson and 

others, in press). The general pattern includes summit deflation, 

followed within minutes to hours by deflation at Pu‘u ‘Ō‘ō and 

a waning of eruptive activity. After a period of hours to a few 

days, inflation begins at the summit and, with a lag of minutes 

to hours, at Pu‘u ‘Ō‘ō (sometimes accompanied by a surge in 

lava effusion). DID and DI events were probably occurring 

before 1999, based on summit tilt cycles that correlated with 

pauses and surges in lava effusion from Pu‘u ‘Ō‘ō, but tilt 

measurements were not available from Pu‘u ‘Ō‘ō o for 

confirmation (Heliker and Mattox, 2003). The coincidence 

between changes in tilt at the summit and at Pu‘u ‘Ō‘ō implies a 

hydraulic connection between the vents (for example, Dvorak 

and Okamura, 1987), and the time lag may represent the flow 

rate of magma through the conduit for at least three reasons: (1) 

summit and Pu‘u ‘Ō‘ō deflation indicate that pressure within the 

conduit was low before the onset of inflation; (2) if the inflation 

were caused by a pressure wave, the pressure pulse would 

approximate the speed of a P-wave through magma, arriving at 

Pu‘u ‘Ō‘ō a few seconds after occurring at the summit (Cervelli 

and Miklius, 2003); and (3) a poroelastic effect (for example, 

Lu and others, 2002) is unlikely, given the presence of an open 

magma conduit between the two locations. Alternatively, the 

delay time may be a function of mechanical factors, like 

dike-wall-rock elasticity and magma viscosity, as modeled by 

Montagna and Gonnermann (2013).

The delay between the onset of inflation at the summit and 

Pu‘u ‘Ō‘ō during DID and DI events generally falls in the range 

of 1–2 hours (fig. 26). If related to magma transport, the delay 

time suggests a magma flow rate of 10–20 km/hr in the ERZ 

conduit—similar to the gravity-driven velocity of lava flowing 

through tubes (for example, Kauahikaua and others, 1998), 

and a few times higher than flow rates modeled for early lava 

fountaining episodes of the Pu‘u ‘Ō‘ō eruption (Hardee, 1987). 

Such a simple model, however, does not account for magma that 

is already in the conduit and that would be pushed at the front 

of a surge from the summit (for example, Parfitt and Wilson, 

1994), implying a shorter transport time than indicated by the 

tilt patterns. The model also ignores the mechanism for varying 

delay times (which can reach up to many hours) and for DI 

events at the summit that were not reflected at Pu‘u ‘Ō‘ō during 

2005–07 despite continued ERZ eruptive activity. The time lag 

in tilt between the summit and Pu‘u ‘Ō‘ō therefore represents a 

fast end member for the transport time along the ERZ.

3

2

1

0

-1

-2

-3

-4

2

1

0

-1

-2

-3

1.0

0.5

0

-0.5

-1.0

-1.5

Tilt, in microradians

A

3

B

C

0 1 2 3

Days (start date varies)

0 1 2 3 4 5 6 7

0 1 2 3 4 5

Figure 26. Plots of some example transient tilt events at Kīlauea’s summit and Pu‘u 

‘Ō‘ō. Summit tiltmeter (black line) is located about 300 m west-northwest of the Hawaiian 

Volcano Observatory, and tilt azimuth is 327° (northwest). Pu‘u ‘Ō‘ō tiltmeter (red line) is 

located on the north flank of the eruptive cone, and tilt azimuth is 308°. Both tilt azimuths 

are radial to deformation sources (assumed spherical), so that positive tilt changes are 

interpreted to be inflation and negative changes are deflation. The time axes for all three 

plots are at the same scale for ease of comparison. A, Tilt record spanning December 

8–12, 2001, showing a deflation-inflation-deflation (DID) event. The time between the 

onset of inflation at Pu‘u ‘Ō‘ō (dashed red line), relative to the summit (dashed black 

line), is approximately 36 minutes. B, Tilt record spanning March 21–27, 2010, showing a 

deflation-inflation (DI) event with a U-like shape. The time between the onset of inflation 

at Pu‘u ‘Ō‘ō relative to the summit is approximately 271 minutes. C, Tilt record spanning 

September 27–29, 2010, showing a DI event with a V-like shape. The time between the 

onset of inflation at Pu‘u ‘Ō‘ō, relative to the summit, is approximately 177 minutes.
216 Characteristics of Hawaiian Volcanoes

Mechanisms of Dike Emplacement at Hawaiian 

Volcanoes

The driving pressure of a dike can be enhanced by either 

increasing the magma pressure in the dike (active intrusion) or 

by decreasing the least compressive stress (σ3) that opposes dike 

opening (passive intrusion). Both conditions facilitate intrusion, 

but eruption is more likely when magma pressure is high 

(Rubin, 1990). At Hawai‘i’s active volcanoes, dikes have been 

observed to have characteristics suggesting emplacement due to 

both increased magma pressure and decreased σ3.

Dikes that are initiated due to high magma pressure 

commonly occur after periods of repose at Kīlauea and Mauna 

Loa and are usually preceded by summit inflation and (or) 

earthquake activity. For example, both the 1975 and 1984 

eruptions of Mauna Loa were heralded by about a year of 

elevated seismicity, although cross-caldera extension 

consistent with inflation was observed only before the 1975 

eruption (Koyanagi and others, 1975; Lockwood and others, 

1987; fig. 22). At Kīlauea, summit inflation is a common 

precursor to dike intrusions, as demonstrated by summit 

tiltmeter records (fig. 12), and is usually accompanied by 

seismicity (Klein and others, 1987). The 1959 Kīlauea Iki 

eruption provides an excellent example. Following a swarm of 

deep earthquakes, Kīlauea’s summit began to inflate in August 

1959. The inflation accelerated in October and November, and 

a swarm of small, shallow earthquakes began and increased in 

number as the inflation rate accelerated. This activity 

culminated in an eruption on November 14 at Kīlauea Iki 

(Eaton and Murata, 1960).

Passive intrusions were documented in 1997 and 1999 

during the 1983–present (as of 2014) ERZ eruption of Kīlauea. 

The 1997 intrusion resulted in a 22-hour-long eruption in Nāpau 

Crater (Owen and others, 2000b; Thornber and others, 2003), 

but the 1999 dike did not breach the surface (Cervelli and 

others, 2002). The intrusions are thought to be a consequence of 

shallow rift zone opening in response to a buildup of extensional 

stress above the opening deep rift zone. The extension is 

ultimately caused by long-term seaward motion of Kīlauea’s 

south flank (Owen and others, 2000b; Denlinger and Morgan, 

this volume, chap. 4), possibly enhanced by the presence of 

preexisting melt pockets within a few kilometers of the surface 

that create zones of weakness (Thornber and others, 2003). 

Such a model is supported by the fact that neither the 1997 nor 

the 1999 intrusion was preceded by summit inflation; instead, 

the summit was deflating prior to both events, suggesting that 

increasing magma pressure did not trigger dike emplacement 

(Owen and others, 2000b; Cervelli and others, 2002). Seismicity 

during these two intrusions was localized along the ERZ and did 

not propagate downrift from the summit—additional evidence 

that the intrusion process was initiated from the ERZ conduit.

The presence of an active magma conduit between the 

summit and Pu‘u ‘Ō‘ō is probably a prerequisite for these 

passive intrusions, since much of the magma supplied to the 

1997 and 1999 dikes came from the conduit system itself 

(Owen and others, 2000b; Cervelli and others, 2002). In 

addition, stress modeling of the 1999 intrusion based on a 

model of long-term GPS velocities from Owen and others 

(2000a) indicates extensional stress favoring dike opening 

in the area of the ERZ where the dike actually intruded 

(Cervelli and others, 2002). Passive dike initiation and ascent 

is not unique to the time period of the Pu‘u ‘Ō‘ō eruption but 

probably also occurred at other times in Kīlauea’s history. For 

example, negligible summit deformation was recorded before 

the 1955 (Helz and Wright, 1992; Wright and Klein, 2014) and 

September 1977 (Dzurisin and others, 1980, Wright and Klein, 

2014) ERZ eruptions, both of which involved fractionated 

magmas thought to have been stored within the rift zone for 

years to decades (Wright and Klein, 2014).

The June 2007 Father’s Day intrusion and eruption 

provides an example of a hybrid type of dike intrusion in which 

both magma pressure and extensional stress contributed to dike 

growth. Summarized here is the chronology and interpretation 

of activity (Poland and others, 2008; Montgomery-Brown and 

others 2010, 2011). At 02:16 HST on June 17, 2007, summit 

tiltmeters began recording rapid deflation as an earthquake 

swarm began on the ERZ, indicating magma withdrawal from 

the summit and intrusion into the ERZ near Pauahi Crater. After 

~6 hours, seismicity jumped downrift, and the rate of summit 

deflation increased as the intrusion migrated to the east (fig. 25). 

During the intrusion, GPS stations on the south flank of Kīlauea 

suggested the occurrence of an aseismic slip event along the 

basal décollement (Brooks and others, 2008; MontgomeryBrown and others, 2010, 2011). Just after midnight HST 

on June 19, a small-volume (~1,500 m3) fissure eruption of 

high-MgO basalt (compared to what had been erupting from 

Pu‘u ‘Ō‘ō) began on the northeast flank of Kanenuiohamo 

(Poland and others, 2008, 2012; Fee and others, 2011). Summit 

deformation reversed to inflation at about 10:30 HST on June 

19, indicating an end to the Father’s Day intrusive and eruptive 

activity. Models of deformation data suggest that the intrusion 

comprised two en echelon dike segments (Montgomery-Brown 

and others, 2010).

The Father’s Day event was preceded by an increase in 

magma pressure at the summit of Kīlauea, as suggested by 

summit inflation that had been ongoing since 2003, due to 

an increase in magma supply to the volcano (see “2003–07 

Increase in Magma Supply to Kīlauea” section above and 

Poland and others, 2012). The magma pressure finally ruptured 

the already-full ERZ conduit early on June 17, 2007, and 

magma was pushed from the summit into the ERZ, forming the 

western dike segment (using the terminology of MontgomeryBrown and others, 2010). The fact that summit deflation 

occurred contemporaneously with the onset of the intrusion, 

instead of lagging by minutes or hours, also argues for summit 

magma pressure (and not extension due to south flank motion) 

as a trigger for dike initiation.

The decreasing rate of summit deflation after the first 

6 hours indicates that growth of the intrusion was waning 

when the aseismic décollement slip event began at about 

08:00 HST on June 17. Stress models suggest that the 
Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 217

décollement slip may have been triggered by the intrusion 

(Brooks and others, 2008). Once begun, the décollement 

slip would have promoted additional dike opening, possibly 

prompting the jump of dike emplacement to the eastern 

segment (Montgomery-Brown and others, 2010), which 

ultimately became larger in volume and had a longer period 

of growth than the western segment. Feedback between 

dike opening and décollement slip therefore combined to 

drive much of the Father’s Day dike emplacement activity, 

demonstrating that the event was a combination of active and 

passive processes.

Summary

Dikes are the dominant means of initial magma transport 

within shield-stage Hawaiian volcanoes away from the summit 

reservoir system. Most dikes in Hawai‘i appear to be driven 

by high magma pressure, as opposed to low levels of stress in 

the crust that oppose dike opening, which means that many 

Hawaiian dikes erupt. This condition can be reversed (intrusion 

favored over eruption) during times of enhanced extensional 

stress, as occurred following the 1975 earthquake beneath 

Kīlauea’s south flank, but such periods apparently last only 

months to years. Rates of dike propagation in Hawai‘i are 

similar to those measured at other basaltic volcanoes worldwide 

(generally about 1 km/hr). During long-term eruptions, the 

initial dikes evolve into open pipe-like conduits, and magma 

flow rates may reach 10–20 km/hr based on tilt measurements 

during transient deformation events.

Dike emplacement in Hawai‘i is a consequence of both 

active and passive processes. Active intrusions, common following periods of intrusive/eruptive repose, are driven by magma 

pressure in summit reservoirs at Kīlauea and Mauna Loa. 

Passive intrusions, in contrast, are driven by extensional stress 

(especially prevalent at Kīlauea due to seaward motion of the 

volcano’s south flank), the best examples of which include the 

1997 and 1999 ERZ dikes at Kīlauea. Feedback between dike 

emplacement and flank slip creates intrusions that are a consequence of both magma pressure and extensional stress, like the 

2007 Father’s Day intrusion and eruption along Kīlauea’s ERZ.

Data collected on dike emplacement in Hawai‘i provide 

important insights into magma transport at basaltic volcanoes. 

Perhaps most critical from a hazards perspective is that active 

intrusions are preceded by summit inflation, earthquake swarms, 

and other activity, making them straightforward to forecast 

(although predicting their precise timing remains elusive). Previous work has demonstrated as much—for example, Koyanagi 

and others (1975) correctly interpreted the buildup in seismicity at Mauna Loa during 1974–75 as a precursor to eruption 

(the eruption occurred shortly after their manuscript had been 

submitted for publication). Passive intrusions, on the other hand, 

are a consequence of extensional stress buildups and have no 

obvious precursors. The likely location of such activity can be 

forecast from stress models (Cervelli and others, 2002), but their 

general timing cannot, as yet, be anticipated.

Conclusions

In the more than 100 years since the Hawaiian Volcano 

Observatory was founded by Thomas A. Jaggar, Jr., in 1912, a 

wide array of volcanic and tectonic activity has been tracked in 

Hawai‘i. Monitoring of Kīlauea has resulted in one of the most 

comprehensive sets of geological, geophysical, and geochemical data available for any volcano in the world. Mauna Loa is 

nearly as well observed, especially since the establishment of 

seismic monitoring in the 1950s and continuous deformation 

networks in the 1990s. These data provide a context for research 

into magma supply to, and storage and transport within, shield 

stage basaltic volcanoes.

Examinations of the topography, bathymetry, and gravity of the Hawaiian-Emperor chain of islands and seamounts 

have allowed for the reconstruction of the magma supply from 

the hot spot over time. The rate of supply has fluctuated over 

millions of years, with the present representing a peak in supply 

rate. Measurements of deformation and erupted volumes suggest a relatively constant supply rate to Kīlauea over decades, 

although supply appears to fluctuate between Mauna Loa and 

Kīlauea. For example, between 1952 and 2014, Kīlauea has 

been almost continuously active while Mauna Loa has erupted 

only twice (individual eruptions may be large in volume, but 

Mauna Loa’s time-averaged effusion rate has dropped; Lockwood and Lipman, 1987). The supply rate to Kīlauea from 

the mantle more than doubled during 2003–07, indicating that 

short-term changes in supply do occur. The 2003–07 increase in 

supply caused important changes in volcanic and tectonic activity at the surface and was preceded by seismic activity, inflation, 

and increased gas emissions. It should be possible to forecast 

future changes in supply by tracking these indicators, especially 

CO2 emissions. During periods of heightened magma supply, 

both Mauna Loa and Kīlauea may receive magma from the hot 

spot, as indicated by inflation at Mauna Loa during 2002–09, 

suggesting a deep connection between the hot spot source of 

magma, Kīlauea, and Mauna Loa. The lack of significant CO2

emissions from Mauna Loa raises the intriguing possibility that 

most of the CO2 degassed by magma supplied from the hot spot 

may ascend through, and be emitted from, Kīlauea.

Magma storage within Hawaiian volcanoes is 

accommodated by a complex series of interconnected 

reservoirs and pathways. We introduce a refined model for 

Kīlauea’s magma plumbing system based largely on past 

studies, combined with more recent seismic imaging and 

deformation data collected by GPS and InSAR. The model is 

consistent with geochemical data and previous geophysical 

results. There are at least two long-term (that is, decades 

or longer) magma storage areas beneath Kīlauea’s summit, 

each of which is connected to a rift zone system. The deeper 

summit magma reservoir, located beneath the south caldera, is 

the primary magma storage area for Kīlauea and feeds magma 

into the East and seismic Southwest Rift Zones at about 3-km 

depth; both rift zones themselves contain isolated pockets 

of stored, crystallizing magma emplaced at earlier times. A 

shallower reservoir, beneath the east margin of Halema‘uma‘u 
218 Characteristics of Hawaiian Volcanoes

Crater, transports magma to summit eruptive vents, as well 

as to the east (toward Kīlauea Iki) and southwest (toward 

Maunaiki) at depths of 1 km or less. Magma also accumulates 

episodically beneath Keanakāko‘i Crater. The question of 

magma storage in Kīlauea’s deep rift zones is still a matter of 

uncertainty and is an important topic for future research. 

Mauna Loa’s magma system appears less complicated, 

but a lack of eruptive activity there since the advent of high 

spatial- and temporal-resolution deformation measurements 

limits our ability to detect all elements of the volcano’s 

plumbing system. The main magma reservoirs beneath Mauna 

Loa are located beneath the southeast margin of the caldera 

and along the length of the caldera, probably reflecting an 

interconnected storage area with complex geometry. These 

conceptual models of the current magma plumbing systems 

of Kīlauea and Mauna Loa will no doubt be refined as new 

geophysical and geochemical techniques are developed and 

applied, and as future eruptive activity offers additional 

opportunities to study magma storage. For now, they provide a 

framework for interpreting magmatic and volcanic activity in 

Hawai‘i and, by extension, other basaltic volcanoes.

Given their frequent eruptions, Hawaiian volcanoes 

afford an unprecedented opportunity to observe magma 

transport. Studies of magma intrusion in Hawai‘i have 

provided much of the foundation for understanding dike 

emplacement in general, from the mechanics of propagation 

to the seismic and geodetic changes that result. During longlasting eruptions, dikes evolve into open conduits, and magma 

flows freely between storage areas and the eruption site. Dike 

intrusion is triggered by both an increase in magma pressure 

within a storage reservoir (which generally occurs following 

periods of repose or during surges in supply) and buildups in 

extensional stress (generally due to flank motion), even during 

ongoing eruptions, when an open magma conduit already 

exists. These two processes reinforce each other, resulting in 

complex triggering relations between magmatic and tectonic 

activity, as demonstrated by the June 2007 Father’s Day 

intrusion and eruption at Kīlauea.

During the first century of the Hawaiian Volcano 

Observatory’s operation, understanding of magma supply, 

storage, and transport at Hawaiian volcanoes has evolved 

from a rudimentary set of observations and generalizations 

to quantitative models and detailed characterizations. These 

insights have been applied to numerous volcanoes on the Earth 

and on other planets and have aided in the development of 

volcano monitoring techniques and hazard mitigation efforts. 

We expect that Hawaiian volcanoes will continue to serve in 

this role for the next 100 years.

Acknowledgments

We are grateful to the staff and volunteers of the 

Hawaiian Volcano Observatory, past and present, who 

collected the data and made many of the interpretations 

that we drew upon. Several figures were prepared with the 

use of the Generic Mapping Tools software (Wessel and 

Smith, 1998). We are indebted to Lisa Faust (Cascades 

Volcano Observatory) for her careful and artistic rendering 

of figure 10. Numerous colleagues offered feedback on 

this manuscript. We are particularly indebted to editorial 

assistance from Jane Takahashi, comments from Tom Wright, 

Dan Dzurisin, Kyle Anderson, and Don Swanson, and formal 

reviews from Dave Sherrod, Christina Plattner, and Jim 

Kauahikaua. 

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Magma Supply, Storage, and Transport at Shield-Stage Hawaiian Volcanoes 233

Appendix. Inverse Modeling Methods

All inverse models in this study (shown in figs. 15–18) 

are kinematic and assume a homogeneous, isotropic, linearly 

elastic half-space, with data weighted according to their 

respective variances. We assume a Poisson’s ratio of 0.25 

and a shear modulus of 3×1010 Pa. Modeled displacements 

are computed from expressions that relate deformation 

of point (Mogi, 1958) or planar (Okada, 1985) sources 

to surface displacements. Parameters estimated for point 

sources include east position, north position, depth, and 

volume change. The parameters estimated for the planar 

sources are length, width, depth, dip, strike, east position, 

north position, and opening.

Model parameters (east position, north position, depth, 

and volume change) for point sources in the Halema‘uma‘u 

(figs. 16 and 17) and Keanakāko‘i (fig. 18) models were 

determined using a Markov Chain Monte Carlo (MCMC) 

optimization algorithm (Metropolis and others, 1953). We 

assumed that the a priori distributions of the parameters 

are uniform between broadly chosen bounds. The posterior 

distributions show the range of parameters that fit the data 

acceptably. The preferred model is the one with the smallest 

misfit.

Sill opening for the south caldera plus upper 

Southwest Rift Zone model (fig. 15) was determined by first 

constraining the parameters describing a single, uniformly 

opening planar dislocation (length, width, depth, dip, strike, 

east position, north position and opening) with an MCMC 

optimization (fig. 27). The preferred dislocation was then 

expanded horizontally and subdivided into 375-m by 375-m 

squares for the distributed opening model. We used a nonnegative least squares algorithm that minimizes the L2‐norm 

of the weighted residuals. Spatial smoothing was applied 

using a Laplacian operator with the optimal weight chosen 

by the L-curve criterion (fig. 28; Hansen, 1992).

Figure 27. Histogram showing the distribution of sill depth that fits 

deformation of the south caldera and upper Southwest Rift Zone (fig. 15A). 

The best-fitting sill, which is used to constrain the distributed opening 

model (fig. 15B), is at 2.9-km depth (red dashed line), with 95-percent 

confidence values spanning 2.1–5.1 km. The best-fitting depth does not 

align with the peak of the distribution.

0 2 4 6 8

0

1,000

2,000

3,000

4,000

5,000

Depth, in kilometers

Number of models
234 Characteristics of Hawaiian Volcanoes

Figure 28. Plots of model misfit and smoothing parameters for distributed-opening sill model shown in fig. 15. A, 

Root-mean square (RMS) misfits for sill models of varying depths. The misfit is minimized at ~3-km depth. B, L-curve 

(Hansen, 1992) for models of varying values of the smoothing weight. The curve compares the weighted residual norm 

(a measure of data misfit) versus model norm (a measure of model roughness). The optimal smoothing weight is chosen 

as the corner value that represents the smoothest model with minimal increase in misfit.

2,000 3,000 4,000 5,000 6,000

0.0165

0.017

0.0175

0.018

0.0185

0.019

0.0195

0.02

Depth, in meters

RMS, misfit

0 1x10-4 2x10-4

0.1

0.15

0.2

0.25

0.3

0.35

8.86e3 4.28e3 2.07e3 1e3 1.83e4

3.79e4

7.85e4

1.62e5 (preferred value)

3.36e5

6.95e5

1.44e6

2.98e6

6.16e6

1.27e7

2.64e7

5.46e7

1.13e8

 1e9

Residual norm

Model norm

A B
Hawaiian Volcano Observatory scientists collect leveling data along Kīlauea’s East Rift Zone to measure vertical deformation of the 

surface. Pu‘u ‘Ō‘ō is in the background. HVO volunteer Connie Parks (foreground) and Dan Johnson (background) are holding the leveling 

rods, while Maurice Sako reads the gun and Christina Heliker records. USGS photograph by J.D. Griggs, November 12, 1985.
Shielding his face from the heat, Thomas A. Jaggar, Jr., approaches the Halema‘uma‘u lava lake in Kīlauea’s summit caldera. USGS 

photograph, January 11, 1917 (photographer unknown).
Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

1

U.S. Geological Survey.

2

Monterey Bay Aquarium Research Insitute.

Abstract 

Study of the petrology of Hawaiian volcanoes, in particular 

the historically active volcanoes on the Island of Hawai‘i, has 

long been of worldwide scientific interest. When Dr. Thomas 

A. Jaggar, Jr., established the Hawaiian Volcano Observatory 

(HVO) in 1912, detailed observations on basaltic activity at 

Kīlauea and Mauna Loa volcanoes increased dramatically. 

The period from 1912 to 1958 saw a gradual increase in the 

collection and analysis of samples from the historical eruptions 

of Kīlauea and Mauna Loa and development of the concepts 

needed to evaluate them. In a classic 1955 paper, Howard 

Powers introduced the concepts of magnesia variation diagrams, 

to display basaltic compositions, and olivine-control lines, to 

distinguish between possibly comagmatic and clearly distinct 

basaltic lineages. In particular, he and others recognized that 

Kīlauea and Mauna Loa basalts must have different sources.

Subsequent years saw a great increase in petrologic 

data, as the development of the electron microprobe made it 

possible to routinely monitor glass and mineral compositions, 

in addition to bulk rock compositions. We now have 100 years’ 

worth of glass compositions for Kīlauea summit eruptions, 

which, together with expanding databases on prehistoric 

tephras, provide important constraints on the nature of Kīlauea’s 

summit reservoir. A series of chemically distinctive eruptions 

in the 1950s and 1960s facilitated evaluation of magma 

mixing and transport processes at Kīlauea. At Mauna Loa, 

lava compositions are distinctive only at the trace element 

level, suggesting that its summit reservoir is better mixed than 

Kīlauea’s. Most summit lavas at both volcanoes, however, lie 

on olivine control lines having the same olivine composition 

(Fo86–87). Study of the ongoing East Rift Zone eruption at 

Kīlauea has further illuminated the complexity of magma 

storage, resupply, and mixing along this very active rift zone. 

Studies of active and closed-system lava lakes have been 

part of HVO’s efforts since Jaggar’s unique descriptions of the 

Halema‘uma‘u lava lake that existed before 1924. Detailed 

study of closed-system bodies, including the 1959 Kīlauea Iki, 

1963 ‘Alae, and 1965 and prehistoric Makaopuhi lava lakes and 

the Uēkahuna laccolith, have allowed recognition and quantification of processes of basalt differentiation. Specific topics 

reviewed herein include the occurrence of segregation veins and 

related structures, overall cooling history, and patterns of crystallization and reequilibration of olivine in various lava lakes.

In recent decades, study of the submarine slopes of the 

Island of Hawai‘i and of Lō‘ihi Seamount has revolutionized 

our understanding of the early history of Hawaiian volcanoes. 

Observations of Lō‘ihi lavas first established the existence of 

an early alkalic stage in the evolution of Hawaiian volcanoes. 

Stages of volcanic development from inception to tholeiitic 

shield building can be observed in Kīlauea’s submarine and subaerial sections. One distinctive feature of submarine volcanics at 

Kīlauea, Mauna Loa, and Hualālai is that picritic lavas are more 

abundant than in subaerial eruptions. Also, olivine compositions 

of submarine lavas are more magnesian, ranging from Fo88 at 

Kīlauea to Fo89 at Hualālai. The most magnesian glasses known 

from Kīlauea (MgO=14.7–15.0 weight percent) were found 

along the submarine part of Kīlauea’s East Rift Zone.

Contributions to our knowledge of the nature of the mantle 

source(s) of Hawaiian basalts are reviewed briefly, although this 

is a topic where debate is ongoing. Finally, our accumulated 

petrologic observations impose constraints on the nature of 

the summit reservoirs at Kīlauea and Mauna Loa, specifically 

whether the summit chamber has been continuous or segmented 

during past decades. 

Introduction

The youngest volcanoes of Hawai‘i have attracted volcanologists to the Island of Hawai‘i for almost two centuries 

because, first, they erupt frequently, being among the most 

active volcanoes in the world; second, Hawaiian eruptions over 

the past 200 years have usually been effusions of basaltic lava, 

Chapter 6

Petrologic Insights into Basaltic Volcanism at Historically 

Active Hawaiian Volcanoes

By Rosalind T. Helz1, David A. Clague2, Thomas W. Sisson1, and Carl R. Thornber1
238 Characteristics of Hawaiian Volcanoes

rather than violent explosions, so they can be approached and 

observed in reasonable safety; and finally, Hawai‘i’s geographic 

isolation from all continents makes these volcanoes exemplars 

for basaltic volcanism, uncomplicated by interactions with 

continental crust. This chapter reviews our petrologic knowledge of four volcanoes: Kīlauea, Mauna Loa, Hualālai (which 

last erupted in 1801), and Lō‘ihi, a submarine shield volcano off 

the southeast coast of the Island of Hawai‘i (fig. 1), all of which 

have erupted since 1790. Among them, they exhibit three of the 

main stages of Hawaiian volcanism: the preshield alkalic stage 

(Lō‘ihi), the tholeiitic shield-building stage (Kīlauea and Mauna 

Loa), and the postshield alkalic stage (Hualālai), as described 

in Macdonald and Katsura (1964), further developed by Clague 

and Dixon (2000) and summarized in Clague and Sherrod (this 

volume, chap. 3). Of these, Hualālai, Mauna Loa, and Lō‘ihi lie 

on the geochemically defined Loa line (as shown by Jackson 

and others, 1972, expanding on an observation of Dana, 1849), 

whereas Kīlauea lies on the Kea line (fig. 1). Volcanoes on these 

two lines differ in isotopic compositions (Tatsumoto, 1978), as 

well as in the character of their postshield alkalic caps (Wright 

and Helz, 1987), so the four active volcanoes exhibit most of the 

range of behavior of older Hawaiian volcanoes.

Scientists from the Hawaiian Volcano Observatory 

(HVO) and their collaborators have pursued a wide range of 

petrologic and chemical studies at these volcanoes. These 

Loa line

Kea line

156° 155°

19°

20°

Kohala

Hualālai

Mauna Kea

Mauna

Loa

Kīlauea

Lō‘ihi

Waimea

Hilo

Kailua

1678

2515

4206

4169

1247

-950

1000

2000

1000

3000

3000 2000

0

0

20 40 KILOMETERS

10 20 MILES

studies have elucidated processes of basalt crystallization and 

fractionation, magma mixing, and transport in the volcanic 

plumbing system, both subaerial and submarine. Lastly, 

Hawai‘i’s erupted materials, especially from the currently active 

volcanoes, continue to serve as a test-bed for studies of various 

trace-element and isotopic systems, allowing scientists to 

evaluate their behavior during generation, transport, degassing, 

fractionation, and mixing of basalt, free from contamination by 

crustal and most sedimentary rocks. 

The topics reviewed herein include (1) petrologic work 

at HVO from 1912 to 1958, which focused on eruptions at 

Kīlauea and Mauna Loa, (2) petrologic studies of subaerial 

activity at Kīlauea and Mauna Loa, especially summit 

activity from 1959 to the present, (3) Kīlauea’s ongoing East 

Rift Zone eruption, in its 30th year as of the time of this 

writing, (4) selected topics on Kīlauea lava lakes, and (5) the 

petrology of the submarine sections of these four historically 

active volcanoes. The final section will consider what this 

work, in particular that of HVO scientists, has contributed to 

our understanding of primary tholeiitic magma compositions 

and their origin, the origin of preshield lavas, and the nature 

of the summit magma chambers of Kīlauea and Mauna Loa. 

Early History of Petrology at HVO 

(1912–58)

Many of the earliest scientific studies of Kīlauea and 

Mauna Loa were petrologic in nature, including analyses of 

lava flows and descriptions of their mineralogy and texture, 

carried out by scientific visitors to Hawai‘i (see reviews 

in Peterson and Moore, 1987; Wright and Helz, 1987). 

The 1912 founding of HVO, located on the rim of Kīlauea 

Caldera, provided greater continuity of observation, and the 

work of Thomas Jaggar and others associated with HVO 

greatly increased the extent and quality of observations of 

eruptive behavior and degassing activity. These early reports 

have been compiled in Bevens and others (1988) and Fiske 

and others (1987) to make them more accessible to modern 

researchers. The collection and analysis of lava samples 

was not the main focus of the early efforts, but the scientific 

community continued to collect, analyze, and slowly build 

up a body of petrographic and chemical data on Kīlauea and 

Mauna Loa basalts. 

Jaggar’s observations on the active lava lake in 

Halema‘uma‘u Crater remain unique, because the behavior 

observed from before 1912 until 1924 was only partly 

replicated in subsequent Halema‘uma‘u eruptions of 1952 

(Macdonald, 1952, 1955) and 1967–68 (Fiske and Kinoshita, 

1969; Kinoshita and others, 1969). Jaggar (1917a, 1917b) 

described the filling, circulation patterns, and draining of the 

lava lake in Halema‘uma‘u and established its depth at that 

time as 14 m. He also obtained a temperature profile through 

the active lava lake (Jaggar, 1917a). Lastly, he offered a 

thorough description and evaluation of the 1924 phreatic 

Figure 1.  Index map of the Island of Hawai‘i, showing the five

volcanoes that make up the island (delineated by dotted lines; triangles 

mark summits) and the active submarine volcano Lō‘ihi to the south. After

Peterson and Moore (1987). Gray bands show trajectories of the Loa and 

Kea geochemical lines. Elevations in meters; contour interval 500 m.
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 239

eruption and noted the occurrence of incandescent intrusions 

in the walls of Halema‘uma‘u, which had been truncated 

during its widening (Jaggar and Finch, 1924). 

The gradual refilling of Halema‘uma‘u after its 1924 

collapse was documented by Jaggar in detail in The Volcano 

Letter (Fiske and others, 1987). Petrologic sampling of 

this activity was restricted to the collection of small spatter 

samples; only one sample, taken in 1931 (as reported in 

The Volcano Letter, no. 288; Jaggar, 1932), was big enough 

to analyze by the methods then available. During this 

period, eruptions at Mauna Loa were more spectacular, 

with significant activity in 1940 and 1942 (Macdonald and 

Abbott, 1970) and especially in 1949 and 1950 (Macdonald 

and Orr, 1950; Finch and Macdonald, 1953), but sampling 

for analysis was minimal, as can be seen in the database used 

by Wright (1971). 

Petrologic research during this period expanded the base 

of samples analyzed petrographically and chemically. Lava 

classification at the time depended on optical characterization 

of groundmass minerals, especially pyroxenes, so samples 

that were holocrystalline and as coarse-grained as possible 

were desirable. Macdonald (1949a), using these techniques, 

correctly identified the shield-building lavas of Hawaiian 

volcanoes as tholeiitic in nature. His work on the 1840 

eruption of Kīlauea (Macdonald, 1944), in which upper 

rift olivine-poor flows and lower rift picritic lavas erupted 

almost simultaneously, was the first to suggest that olivine 

settling in the conduits of Hawaiian volcanoes might explain 

such differences in composition; subsequent analysis of the 

olivine-poor material (Macdonald, 1949b), together with 

analyses of augite separates from the 1921 Kīlauea lava, 

suggested that the uprift lava might have lost some pyroxene 

and plagioclase, as well (Muir and Tilley, 1957). Macdonald 

and coworkers also provided descriptions of the 1952 

Halema‘uma‘u activity (Macdonald, 1952, 1955) and the 

brief 1954 summit eruption (Macdonald, 1954; Macdonald 

and Eaton, 1957). The final eruption at Kīlauea during this 

period occurred in 1955, the first along the lower East Rift 

Zone since the 1840 eruption. This eruption was extensively 

investigated, with 10 samples analyzed for major elements 

and 8 for selected trace elements, the latter data the first such 

for Kīlauea (Macdonald and Eaton, 1964). The 1955 lavas 

were unlike the 1840 lavas, being much more differentiated 

than those of any previously observed eruption at Kīlauea. 

Powers’s (1955) paper on the composition and origin 

of Hawaiian basalts offers the most advanced discussion of 

Hawaiian petrology of the period. That report introduced the 

use of magnesia variation diagrams (with MgO replacing 

SiO2 on the abscissa) as a superior means of visualizing 

relationships among lavas where addition or removal 

of olivine was the dominant process. Their use enabled 

Powers to recognize that Hawaiian olivine basalts differed 

from one volcano to the next and, specifically, that the 

historical lavas of Kīlauea and Mauna Loa were distinct. He 

introduced the concept of “magma batches” to describe these 

differences, and the term “olivine-control lines” to describe 

variations within a magma batch produced by variation in 

olivine only. Additional analytical and petrographic data on 

Kīlauea and Mauna Loa lavas presented by Tilley (1960) and 

Tilley and Scoon (1961) provided further evidence that the 

historical lavas at these two volcanoes are chemically and 

petrographically distinct. 

Subaerial Activity at Kīlauea and Mauna 

Loa

Petrologic investigations expanded rapidly as the 1950s 

drew to a close, stimulated by the spectacular 1959 summit 

(Kīlauea Iki) and 1960 Puna (Kapoho) eruptions. Developments that have greatly enhanced the subsequent evolution of 

petrology include the following: 

(1) Collection of more samples, especially in real time, 

through the course of an eruption. This was begun during 

the 1955 eruption, but such collecting was greatly expanded 

during the 1959 and 1960 eruptions and has been maintained 

in all subsequent activity.

(2) Submission of more samples for whole-rock analysis, 

beginning with the 1959 and 1960 eruptions (Murata and 

Richter, 1966; Richter and Moore, 1966). Gaps in earlier 

sampling and analysis were filled in wherever possible, 

including sampling of young prehistoric flows (Wright, 1971). 

(3) Development of microbeam analytical techniques, 

beginning with the electron microprobe, which have made it 

possible to obtain chemical analyses of glass and individual 

mineral phases in volcanic samples. The first significant 

contributions involving microprobe data were Moore and 

Evans (1967) and Evans and Moore (1968), on mineral 

compositions from a suite of samples from the prehistoric 

Makaopuhi lava lake. Other early microprobe studies looked 

at silicate and Fe-Ti oxide phenocrysts (Anderson and Wright, 

1972) and sulfides (Desborough and others, 1968) in various 

Kīlauea lavas. 

(4) Use of experimental petrology to determine the phase 

relations of natural basalt samples (for example, Thompson 

and Tilley, 1969), which supports field observations on the 

crystallization sequence of lavas. Melting experiments, 

coupled with microprobe analysis of the resulting phases, led 

to the use of glass composition (the MgO content of olivinesaturated melts) as an indicator of quenching temperature 

at both Kīlauea and Mauna Loa. Originally developed to 

allow temperature estimates for glassy core from Kīlauea 

Iki lava lake (Helz and Thornber, 1987), this technique has 

subsequently been extended to other lavas at Kīlauea (Helz 

and others, 1995) and Mauna Loa (Montierth and others, 

1995). 

These expanded observations and quantitative analyses 

of rocks, glasses, and mineral phases have produced advances 

and further refinement of models for magma fractionation, 

mixing, transport to eruption sites, and storage. 
240 Characteristics of Hawaiian Volcanoes

Melt and Rock Compositions at Kīlauea and 

Mauna Loa

Before reviewing the range of melt and rock compositions observed for Kīlauea and Mauna Loa, let us look at the 

overall geography and structure of these two volcanoes

(fig. 2). Each volcano has a summit caldera and two prominent rift zones; these features are the focus of almost all 

eruptive activity. Magma rises from the mantle beneath the 

summit, where it may enter the summit magma reservoir or 

move out into either of the two rift zones. In discussing the 

compositional data, we review Kīlauea summit compositions 

first; discussion of rift lavas through the 1970s follows, in 

chronological sequence, followed by a summary of Mauna 

Loa compositions. The ongoing Pu‘u ‘Ō‘ō-Kupaianaha eruption on Kīlauea’s East Rift Zone (January 3, 1983–present) is 

discussed last, in a separate section below. 

Summit Melts at Kīlauea

Figure 3 presents data for selected major elements in 

melts erupted at the summit of Kīlauea from 1912 to 2010, a 

period of almost 100 years. The data for the period 1912–34 

are newly presented microprobe analyses (appendix table 1) 

of small spatter samples collected by Jaggar. The results show 

that melts erupted effusively within Halema‘uma‘u Crater and 

on the floor of Kīlauea Caldera had consistently low MgO 

contents of 6.4–7.6 weight percent over the entire period. 

MgO content may have decreased slightly with time from 

1912 to 1924, but then increased during and after the refilling of Halema‘uma‘u, following the 1924 phreatic eruption, 

by approximately 0.9 weight percent MgO over the 86-year 

period, as shown by the least-squares-fit lines in figure 3. 

For most other major oxides (for example, CaO and TiO2, 

lines shown in fig. 3), the concentrations decrease as MgO 

Figure 2. Map showing the summit calderas and rift zones of Kīlauea and Mauna Loa Volcanoes and the sites of the Mauna Ulu and Pu‘u ‘Ō‘ō eruptions

on Kīlauea’s East Rift Zone. The location of the submarine volcano Lō‘ihi is also shown. Towns shown by black dots. After Babb and others (2011).

Lōʻihi

Hilo

Kapoho

Hawaiian Volcano

Observatory

Mauna Loa

summit caldera

(Mokuʻāweoweo) Kīlauea Iki

Mauna Ulu 

Puʻu ʻŌʻō

East Rift Zone

Southwest 

Rift 

Zone

Northeast Rift Zone

Southwest

Rift

Zone

Hualālai

Mauna Kea

Kohala

HAWAI‘I

Area

of map

Halemaʻumaʻu

Kīlauea

summit 

caldera 

Kalapana

Kīlauea

Kīlauea

Mauna Loa

155°45' 155°30' 155°15' 155°

19°

19°15'

19°30'

19°45' Mauna Kea

Mauna Loa

0

0

10 20 KILOMETERS

5 10 MILES

Kupaianaha
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 241

48

49

50

51

52

53

SiO2 content, in weight percent

9

10

11

12

13

14

15

16

Oxide content, in weight percent

Al2O3

FeO

0

2

4

6

8

10

12

14

1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 2010

Oxide content, in weight percent

Year of eruption

CaO

MgO

TiO2

1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 2010

1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 2010

Figure 3.  Time plots showing major oxide composition (color-coded by oxide) of intracaldera glasses erupted at the summit of Kīlauea from

1912 to 2010 (rectangles for scoria and spatter, diamonds for 1924 Pele’s tears), together with extracaldera glasses from the 1959, 1971, and

1974 eruptions (crosses) and selected whole-rock analyses from 1911 to 1982 (circles). Glass analyses from 1912–34 are in appendix tables 1 

(M. Beeson, analyst) and 2 (T. Rose, analyst); most 1952–82 intracaldera glass data are from Helz and others (1995) and Wright and Helz (1996). 

Analyses showing the range of MgO in the 1959 glasses are from Helz (2009), and those for 1971 and 1974 are in appendix table 3 (R.T. Helz, 

analyst). Data showing the range in the current (since 2008) summit activity are from Wooten and others (2009) and C.R. Thornber (unpublished 

data). Whole-rock data for lavas erupted from Halema‘uma‘u or onto the floor of Kīlauea Caldera are from Wright (1971), Basaltic Volcanism

Study Project (1981), Duffield and others (1982), Baker (1987), and Lockwood and others (1999). Arrows mark the time of the 1924 phreatic

eruption and draining of Halema‘uma‘u. Trend lines are discussed in the text.
242 Characteristics of Hawaiian Volcanoes

<6 6.0 6.5 7.0 7.5 8.0 8.5 9.0 9.5 10.0 10.5 >11

0

10

20

30

40

50

60

80

Keanakāko‘i glasses

n = 298

MgO content, in weight percent

<6 6.0 7.0 8.0 9.0 10.0 11.0 12.0

0

20

40

60

80

100

120

140

160

180

Pāhala glasses

n = 707

0

5

10

15

20

25

30

1959 glasses

n = 94

Number of analyses

0

50

100

150

200

250

300

Kulanaokuaiki glasses

n = 1,163

70

>13

increases. Trends for the alkalies and P2O5 parallel the variation seen in TiO2

. 

By contrast, some extracaldera summit glasses are more 

magnesian (fig. 3). Glasses from the 1959 eruption, the vents 

of which lay just outside the caldera near Kīlauea Iki (fig. 2), 

range from 6.4 to 10.0 weight percent MgO (Helz, 1987a, 

2009) and vary widely in other oxides, as well. New data 

on glasses from the 1971 and 1974 summit eruptions (fig. 3, 

appendix table 3) show that some extracaldera summit glasses 

are more magnesian than the intracaldera glasses, even along 

the same eruptive fissure.

Glasses with MgO contents greater than the intracaldera 

summit glasses (MgO=6.4–7.6 weight percent) also occur 

in tephra layers at Kīlauea. The 1924 tephra includes glassy 

Pele’s tears with as much as 8.7 weight percent MgO (appendix table 2). Glasses in older tephra range to even higher MgO 

contents. The range and relative frequency of glass MgO 

contents in the 1500–1800 C.E. Keanakāko‘i Tephra Member 

of the Puna Basalt (Mastin and others, 2004; Swanson and 

others, 2012; Helz and others, 2014) and of five sections of the 

400–1000 C.E. Kulanaokuaiki Tephra (Fiske and others, 2009; 

Rose and others, 2011; Helz and others, 2014) are shown in 

figure 4, with the 1959 glasses for comparison. Figure 4 also 

includes data from the uppermost thick tephra in the Hilina 

Pali section, referred to by Easton (1987) as the Pāhala Ash 

(Clague and others, 1995a; Helz and others, 2014), with an 

estimated age of 25–10 ka. The range and distribution of glass 

compositions for the Keanakāko‘i section shown are similar 

to the results of Garcia and others (2011) on another section of 

the deposit and are broadly similar to the data from the 1959 

eruption. The 1959 results are also consistent with those of 

Harris and Anderson (1983) and Anderson and Brown (1993), 

who found MgO contents as high as 10.5 weight percent for 

glass inclusions in olivines from the 1959 eruption. Glasses in 

the two younger units lie mostly between 7 and 11 weight percent MgO. The Kulanaokuaiki and Pāhala histograms include 

a greater number of relatively differentiated (6.0–7.0 weight 

percent MgO) glasses than do the 1959 and Keanakāko‘i plots, 

Figure 4.  Histograms showing frequency distributions of MgO content in glasses from the 1959 Kīlauea Iki scoria (Helz, 2009; and unpublished

data) and various other tephra units. The Keanakāko‘i data are from the Sand Wash section (Mastin and others, 2004; Helz and others, 2014),

which does not include the upper units shown in McPhie and others (1990). The Kulanaokuaiki data are from multiple sections (Rose and others, 

2011; Helz and others, 2014), and the Pāhala data are from Helz and others (2014). The Kulanaokuaiki and Pāhala tephra contain glasses with

MgO>11.0 weight percent, unlike the younger tephras.
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 243

but in these older tephra, the distribution has a tail that extends 

beyond 12 weight percent MgO, higher than observed in the 

Keanakāko‘i and 1959 distributions. 

Summit Lava Compositions Versus Melt

Compositions

Figure 3 also includes whole-rock analyses for intracaldera 

lavas from 1911 through 1982, including data for one eruption (1975) for which no glass data are available. The wholerock data run about a percent higher in MgO (7.0–8.8 weight 

percent) than the intracaldera glasses, and the data from 1932 

through 1982 show the same slight upward trend in MgO with 

time as the glasses. Other oxide components vary less, suggesting that the bulk samples differ from the glasses mainly in containing varying amounts of olivine crystals. Lavas erupted at the 

summit but immediately outside the caldera have much higher 

bulk MgO contents, correlating with significantly higher olivine 

phenocryst contents. This pattern can be seen in the 1959 

Kīlauea Iki eruption (Murata and Richter, 1966) and also in the 

1971 and 1974 summit eruptions, where lavas immediately outside the caldera have MgO contents of 10.2–11.5 weight percent 

(Duffield and others, 1982; Lockwood and others, 1999). In 

the 1959 eruption, higher bulk MgO contents correlate strongly 

with more Mg-rich glasses; new data on samples of 1971 and 

1974 scoria (appendix table 3) confirm that this is true for those 

eruptions, as well.

The data in figure 3 demonstrate that melts differentiated to 

the point of multiple saturation (where olivine would coprecipitate with augite and perhaps plagioclase) are present in the 

upper part of Kīlauea’s summit magma reservoir over long time 

intervals. Similar stability in the summit reservoir’s melt composition was inferred for the ‘Ailā‘au lavas (Clague and others, 

1999) on the basis of glasses found as rinds on the ‘Ailā‘au 

flows, which were erupted over a period of at least 50 years, 

ending in about 1470 C.E. However, in the absence of tephra 

or scoria samples for ‘Ailā‘au, this inference is not entirely 

conclusive, for two reasons. First, glasses from near-vent scoria 

or tephra can be more magnesian (and hence quenched from 

somewhat higher temperatures) than selvages from lava flows, 

as was seen in samples from the 1984 Mauna Loa eruption 

(Helz and others, 1995) or during the high-fountaining episodes 

at Pu‘u ‘Ō‘ō, where flow selvages never preserved the more 

magnesian melts seen in tephra from fountains (as discussed 

in Helz and others, 2003). Second, some extracaldera summit 

glasses, such as those from the 1959, 1971, and 1974 eruptions, 

have relatively high MgO contents, indicating the availability 

of more magnesian liquids in the summit region. It is therefore 

possible that the MgO content of erupted melts for the ‘Ailā‘au 

flows was higher; this seems particularly likely for the picritic3

‘Ailā‘au lavas. 

The occurrence of olivine-rich lavas near Kīlauea’s 

summit is similar to that of the highly magnesian glasses. 

The 1959 lavas, erupted just outside the caldera, were picritic 

tholeiites (Murata and Richter, 1966), and the picritic flows 

within the (precaldera) ‘Ailā‘au sequence are near the ‘Ailā‘au 

vents, just to the east of Kīlauea Iki Crater (fig 2). Other historical picritic tholeiites have been erupted lower on the East 

Rift Zone, including those of the 1840 eruption (Macdonald, 

1944) and the 1968 Hi‘iaka eruption (Nicholls and Stout, 

1988).

Rift Zone Lavas 

Kīlauea has also erupted differentiated lavas, that is, those 

with bulk compositions produced by multiphase fractionation, 

along its rift zones. The best studied are the 1955 lavas (Macdonald and Eaton, 1964; Wright and Fiske, 1971), which have 

bulk MgO contents of 5.04–5.68 weight percent in the early 

part of the eruption and 6.19–6.69 weight percent in the later 

lavas. The early 1955 lavas have very complex phenocryst 

assemblages, including three Fe-Ti oxide phases in addition to 

olivine, augite, plagioclase, and orthopyroxene (Anderson and 

Wright, 1972). The 1977 lavas, compositionally similar to the 

early 1955 lavas (MgO=5.28–5.84 weight percent) are nearly 

aphyric, containing minor plagioclase plus microphenocrysts 

of olivine+augite+plagioclase (Moore and others, 1980). The 

recent discovery of a subsurface body of dacitic magma in the 

Puna geothermal field (Teplow and others, 2009), similar to 

the Mauna Kuwale rhyodacite at Wai‘anae on O‘ahu (Bauer 

and others, 1973), raises the possibility that more differentiated lavas may be buried at Kīlauea. 

Mauna Loa Melts and Lavas

The distributions of glass and whole-rock compositions 

at Mauna Loa are simpler than at Kīlauea. Summit and upperflank lavas and glasses are fairly differentiated, with the 1975 

and 1984 eruptions being representative (whole-rock data in 

Lockwood and others, 1987, and Rhodes, 1988; glasses from 

the 1984 eruption in appendix to Helz and others, 1995). More 

magnesian (and olivine-rich) flows and their more magnesian 

host glasses are found lower on the slopes, as reported by Riker 

and others (2009) for the 1859 eruption and by Rhodes (1995) 

for the picritic phases of the 1852 and 1868 lavas; these have 

glass or groundmass MgO contents ranging from 7.9 to 9.9 

weight percent in whole-rock compositions, with as much as 

22.2 weight percent MgO. C.R. Thornber and F.A. Trusdell 

(unpublished data) observe a similar but more modest trend 

in bulk compositions of the 1950 lavas, with MgO contents 

increasing from 6.8 weight percent at the highest-elevation 

vents to 10 weight percent at the lowermost vents. Lastly, 

Sparks’s (1990) study of the precaldera lavas and intrusives 

exposed in the walls of Moku‘āweoweo Caldera at Mauna 

Loa’s summit identified a period of eruption/intrusion of 

MgO-rich magma (suite B, average Mg=10.9 weight percent) 

3

Olivine-rich tholeiites with MgO≥15 weight percent are commonly 

referred to in the Hawaiian literature as picritic. These are, in all cases, tholeiitic, and the various authors consider at least some of the olivine phenocrysts 

to be cumulate.
244 Characteristics of Hawaiian Volcanoes

in prehistoric time. Lavas with bulk compositions significantly more differentiated than 6.8 weight percent MgO, rare 

at Kīlauea, are unknown among Mauna Loa subaerial lavas, 

although one example (MgO content=5.50 weight percent; 

Wright, 1971) has been found in the shallow submarine section.

Olivine Control at Kīlauea and Mauna Loa 

Olivine is the only phenocryst phase in subaerial lavas 

at Kīlauea, except for rare differentiated lavas, such as those 

erupted in 1955 (Macdonald and Eaton, 1964; Wright and Fiske, 

1971; Anderson and Wright, 1972), and it is also the dominant 

phenocryst in Mauna Loa lavas (Wright, 1971). Studies of relatively rare olivine-rich lavas have provided much information 

on variations in the petrographic character and composition of 

phenocrystic olivine (Helz, 1987a; Wilkinson and Hensel, 1988; 

Schwindinger and Anderson, 1989; Rhodes, 1995). The most 

important petrologic consequence of this restricted phenocryst 

population, however, is that closely related lavas may differ in 

composition solely by addition or subtraction of olivine. This 

led to the use of magnesia variation diagrams to display chemical variation in Hawaiian lavas (Powers, 1955), a technique 

extended by Wright (1971, 1973), who found that addition 

or removal of olivine (Fo86.0–87.5) was the dominant process of 

compositional fractionation at both Kīlauea and Mauna Loa for 

all but the most differentiated (bulk MgO<6.8 weight percent) 

lavas. 

Olivine-controlled lava compositions from Mauna Loa are 

distinct from those at Kīlauea in major elements (Powers, 1955; 

Tilley, 1960; Wright, 1971; Rhodes, 1988, 1995; Rhodes and 

Hart, 1995) and in trace elements (Tilling and others, 1987a, 

1987b; Rhodes and Hart, 1995); however, magma batches definable by major-element chemistry are lacking (Wright, 1971). 

By contrast, cross-trend variations within olivine-controlled 

Kīlauea lavas exist and are coherent in time and space (Wright, 

1971), and the distinctive lava groups are referred to as “magma 

batches.” Recognition and tracking of these magma batches 

at Kīlauea have provided critical insights into the processes of 

magma mixing and preeruptive shallow transport in that system 

(Wright and Fiske, 1971). 

Magma Batches, Magma Mixing, and Lateral 

Magma Transport at Kīlauea 

Beginning with the 1952 eruption, scientists routinely 

sought to obtain summit lava samples large enough for bulk 

analysis. This allowed recognition of the diverse nature of magmas fed into Kīlauea from 1952 to the present. Information on 

the succession of magma batches erupted at the summit during 

this time is summarized in table 1. 

New magma batches appeared at Kīlauea’s summit every 

1–7 years from 1952 through 1982. The spacing between these 

batches is quite short, probably less than the time they existed 

in the combined summit-rift system, which for several batches 

is about 10 years (table 1). This relatively rapid succession has 

implications for the size of the batches and the size and continuity of Kīlauea’s summit reservoir. Estimation of storage times is 

difficult because of the paucity of phenocrysts in most summit 

lavas. Mangan (1990), however, using crystal-size distribution 

results for olivines in the 1959 lavas, estimated storage times of 

~10 years, somewhat longer than the spacing between batches. 

This suggests that either storage times are longer than the interval between batches or some olivine is retained in the summit 

reservoir while successive melts pass through. 

Examining the relationship between summit batches and 

rift eruptions, three different periods can be inferred from table 

1, as follows: 

(1) From 1952 through 1968, summit eruptions dominated, 

and the magma batches were defined by a series of distinct 

summit lava compositions (Wright and Fiske, 1971). The 1952 

and 1959E magma batches appeared as components of rift lavas 

only after they had erupted at the summit, while the 1961 and 

1967–68 batches first appeared in the 1960 Puna eruption on the 

East Rift Zone (near Kapoho; fig. 2) before they were erupted at 

the summit.

(2) From 1971 through 1975, activity at Kīlauea was dominated by eruptions at Mauna Ulu (fig. 2), though moderately 

extensive summit eruptions (1971, 1974) also occurred. The 

10 magma batches defined for Mauna Ulu (Wright and others, 

1975; Wright and Tilling, 1980) differed more subtly from each 

other than had the magma batches of the earlier period. Summit lavas for the 1971, 1974, and 1975 eruptions were defined 

in terms of Mauna Ulu variants, as shown in table 1. This is the 

reverse of the earlier process, in which summit batches were 

used as the starting point for unraveling magma mixing in rift 

zone lavas; however, the timing of their eruption (they appeared 

at the summit after they had appeared at Mauna Ulu) was the 

same as for all summit batches since 1961. The November 

1975 Kalapana earthquake (M7.7) appears to have brought this 

behavior to an end (Holcomb, 1987).

After a hiatus of two years, eruptive activity at Kīlauea 

resumed with the 1977 eruption (Moore and others, 1980), followed by minor activity in 1979 and 1980, all on the East Rift 

Zone (Holcomb, 1987). 

(3) From 1982 to the present, activity at Kīlauea has been 

dominated by the ongoing Pu‘u ‘Ō‘ō-Kupaianaha East Rift 

Zone eruption (fig. 2), reviewed in the next major section. 

Summit activity has been minimal, consisting, until 2008, of 

two 1-day eruptions in 1982 (Banks and others, 1982; Helz and 

others, 1995), so summit magma batches have been essentially 

undefined. The progressive change in magma partitioning 

between the summit and East Rift Zone seen in table 1 suggests 

that that Kīlauea’s East Rift Zone has become steadily more 

accessible to new magma since 1952.

Period 1: Magma Mixing and Transport at Kīlauea,

1952–68

Magma mixing was proposed early on to be a major 

process at Kīlauea (Wright and Fiske, 1971; Wright, 1973), a 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 245

Time of 

appearance of lava 

at the summit

Time after 

previous 

summit batch

Eruption(s) with summit batch 

as mixing component

Time between 

appearances

Comments, references

1952 ? Late 1955 Puna

1960 Puna

3 years after summit

8 years after summit

Wright and Fiske (1971), Helz and Wright 

(1992) Wright and Fiske (1971), Wright 

and Helz (1996)

1954 2 years 1959 (also a summit eruption) 5 years after summit Equivalent to 1959W component as in 

Wright (1973), Helz (1987a)

1959E 5 years 1960 Puna (latest samples) 3 months after summit Wright and Fiske (1971), Wright and Helz 

(1996)

1961 1.5–2 years 1960 Puna 1 year before summit Wright and Fiske (1971), Wright and Helz 

(1996)

1967–68 6 years 1960 Puna 

August 1968, October 1968, 

February 1969 middle east rift

7–8 years before summit 

1 year after summit

Wright and Fiske (1971), Wright and Helz 

(1996) Wright and others (1975)

August 1971 3–4 years 1969–71 Mauna Ulu 1–2 years before summit Made up of Mauna Ulu 3 & 5 (results in 

Wright and Tilling, 1980); analyses in 

Duffield and others (1982)

July 1974

September 1974

3 years

3 years

1972–74 Mauna Ulu

1972–74 Mauna Ulu

2 years before summit

1–2 years before summit

Equivalent to Mauna Ulu 6 (Wright and 

Tilling, 1980)

Made up of Mauna Ulu 6 & 10 (Wright 

and Tilling, 1980); analyses in Lockwood 

and others (1999)

November 1975 1 year 1972–74 Mauna Ulu 2 years before summit Analysis from Basaltic Volcanism Study 

Project (1981); Mauna Ulu variant 10 

(Wright and Tilling, 1980)

April 1982

September 1982

7 years Pu‘u ‘Ō‘ō episodes 17–20 2 years after summit Analyses in Baker (1987); comparison 

with Pu‘u ‘Ō‘ō data shown in Garcia and 

Wolfe (1988)

2007–present 25 years Equivalent to current Pu‘u 

‘Ō‘ō lava

Concurrent Rowe and others (2009)

Table 1.  Timing of appearance of magma batches at the summit of Kīlauea Volcano from 1952 to the present (2014).

recognition facilitated by the diverse compositions of the summit batches. These early studies were based largely on wholerock major-element chemistry and mixing calculations and 

so were contested as not proven for the late 1955 lavas (Ho 

and Garcia, 1988) or for any of the 1954–60 lavas (Russell 

and Stanley, 1990). However, further work—which included 

petrographic and microprobe analysis, as well as modeling of 

trace element compositions—has demonstrated conclusively 

that the late 1955 and middle to late 1960 lavas were mixed 

magmas (Helz and Wright, 1992; Wright and Helz, 1996). 

One outcome of reexamining the late 1955 lavas was recognition that mixing of thermally disparate magmas produces 

resorption of olivine, even though both magmas are saturated 

with olivine (Helz and Wright, 1992). Thermal equilibration 

is achieved much faster than chemical equilibrium, so crystals from the cooler mixing component are resorbed during 

initial mixing. The widespread occurrence of resorbed olivine 

phenocrysts in Hawaiian lava flows, noted earlier by Macdonald (1949a) and Powers (1955), and examples of which can 

be found even in the picritic 1959 eruption (Helz, 1987b), is 

rarely due to reaction with melt to form pyroxene, but rather 

is evidence that magma mixing is widespread at Kīlauea and 

other Hawaiian volcanoes. 

Reexamination of the lavas of the 1955 and 1960 eruptions also provides insight into the length of time that petrographic evidence for mixing may be expected to survive. 

The late 1955 lava flows had mixed phenocryst assemblages, 

with resorbed and (or) reversely zoned crystals, consistent 

with very recent magma mixing (Helz and Wright, 1992). 

The early 1960 lava flows, although identical in composition 

(Wright and Fiske, 1971), show no such evidence for mixing, containing only a modest amount of microphenocrystic 

olivine+augite+plagioclase (Wright and Helz, 1996). The 5 

years that elapsed between the two eruptions appears to have 

been sufficient for chemical and thermal equilibrium to be 

achieved in magmas reinjected by hotter material. 

By contrast, the early 1955 lavas (with their highly 

complex phenocryst assemblages) may have been isolated 

for a long period. A study by Cooper and others (2001) using 

230Th-226Ra dating suggests that the plagioclase phenocrysts in 
246 Characteristics of Hawaiian Volcanoes

those lavas had been in storage for at least 550 years before 

eruption. This is much longer than has been inferred for other 

East Rift Zone magmas, which appear to have crystallized at 

rates of 1–2 weight percent per year (Wright and Tilling, 1980; 

Wright and Helz, 1996) and so may not reflect the age of the 

rest of the lava. 

Samples from late 1960 lavas show renewed petrographic 

and chemical evidence for magma mixing, interpreted as 

coeval with the eruption (Wright and Helz, 1996), with both 

major and trace-element data requiring the addition of a 

succession of four summit batches. The pattern of appearance 

of the summit batches in the 1955 and 1960 lavas elucidates 

transport processes within Kīlauea during this time. The 1952, 

1961, and 1967–68 summit components followed one after the 

other in time (see fig. 16 in Wright and Helz, 1996), as if these 

three batches were transported to the 1960 mixing chamber 

along the same conduit. The 1959E component, which makes 

up as much as 27 percent of the latest 1960 lavas, is out of 

sequence with the 1961 and 1967–68 batches. This very hot, 

highly phyric component may have followed its own path to 

Puna (fig. 16 in Wright and Helz, 1996), just as it followed an 

aberrant path to the surface at the summit (Eaton and Murata, 

1960; Eaton and others, 1987; Helz, 1987b). The presence 

of clusters of mildly deformed olivine crystals (Wright and 

Helz, 1996) in those 1960 samples that required the 1959E 

as a mixing component supports the identification of this 

component, and its relatively deep path out to Puna, because 

such olivines are unusual in subaerial Kīlauea lavas. 

Some late 1960 scoria contain phenocrystic phases 

(orthopyroxene±ilmenite) like those seen in the early 1955 

lavas (Anderson and Wright, 1972), suggesting that there were 

small pockets of magma similar to the early 1955 compositions that had survived until at least 1960. Such isolated 

pockets of magma are plausible parents (Helz, 2008) for the 

dacitic magma discovered in a geothermal drilling project in 

Puna (Teplow and others, 2009).

Period 2: Magma Batches, Mixing, and Transport at 

Kīlauea, 1969–75

During the period from 1969 through 1975, new magma 

batches, defined in terms of Mauna Ulu variants, surfaced along 

the East Rift Zone before erupting at the summit and the upper 

Southwest Rift Zone (table 1). The only earlier stored batch 

recognized was the 1967–68 magma, which contributed to a 

series of eruptions in the middle East Rift Zone one year after 

the Halema‘uma‘u activity ceased (Wright and others, 1975). 

The tiny November 1975 eruption, which occurred during, and 

perhaps as a consequence of, the Kalapana earthquake (Tilling 

and others, 1976), most closely resembles Mauna Ulu variant 10 

(Wright and Tilling, 1980). 

In general, the chemical and petrographic contrasts 

between Mauna Ulu variants were subtler than between the 

summit batches of 1952 through 1967–68; for example, 

variations in trace element contents showed almost continuous, 

monotonic decreases (Hofmann and others, 1984). The Mauna 

Ulu lavas contain as much as 19 weight percent MgO (Wright 

and others, 1975; Wright and Tilling, 1980), consistent with high 

olivine phenocryst content. In a recent paper, Vinet and Higgins 

(2010) document the occurrence of highly magnesian and (or) 

deformed olivine phenocrysts in the Mauna Ulu lavas and 

suggest that they were entrained from the décollement at the base 

of Kīlauea’s volcanic edifice (shown in their figure 20 as below 

11 km depth). However, olivine control for the Mauna Ulu lavas 

shows that the average olivine composition remains Fo86–87, the 

same as is found in Kīlauea’s summit lavas, so it is not clear that 

a source region as deep as the base of the décollement under the 

East Rift Zone is required. 

The time span 1969–75 encompassed the two periods 

of Mauna Ulu activity (Swanson and others, 1979; Tilling 

and others, 1987c) and the 1971 and 1974 summit eruptions 

(Duffield and others, 1982; Lockwood and others, 1999). During 

this period, magma first moved up into the base of the summit 

reservoir, then out the East Rift Zone to Mauna Ulu, without 

erupting at the summit. Later, when the upper East Rift Zone 

became blocked, the summit inflated and the complex 1971 and 

1974 summit eruptions occurred (Duffield and others, 1982; 

Lockwood and others, 1999). In both instances, the summit 

activity was followed by eruptive activity along the upper 

Southwest Rift Zone, which otherwise had seen no activity since 

the 1919–20 eruption (Holcomb, 1987). Most notably, this period 

saw a major change in volcano monitoring: magma movement, 

formerly inferred by using chemically defined magma batches (as 

described above), has since been tracked using newly enhanced 

geophysical monitoring techniques (see Okubo and others, this 

volume, chap. 2, and Poland and others, this volume, chap. 5).

Subsequent work at Kīlauea has been largely focused on 

the ongoing Pu‘u ‘Ō‘ō eruption (January 3, 1983, to the present), which is reviewed in the “Petrologic Overview of Kīlauea’s 

Ongoing Eruption: 1983–2011” section below.

Magma Batches, Magma Mixing, and Magma 

Transport at Mauna Loa

Magma batches defined at the major-element level are lacking 

at Mauna Loa (Wright, 1971), and the lavas are unusually uniform 

within eruptions (Rhodes, 1988). Lavas with phenocrysts of 

olivine and plagioclase plus various pyroxenes are fairly common 

(Macdonald, 1949a,b; Rhodes and Hart, 1995), but lavas with 

complex, obviously mixed, phenocryst assemblages have not 

been reported. The inference from these observations is that the 

magma chamber under Mauna Loa is large enough to buffer melt 

compositions, and that a large part of that volume is crystallizing 

the three-phase assemblage olivine+pyroxene+plagioclase 

(Rhodes, 1988; Rhodes and Hart, 1995). 

Compositional shifts have been observed in trace elements 

and isotopic ratios, however, with concentrations of incompatible 

trace elements decreasing with time (Tilling and others, 1987a,b; 

Rhodes and Hart, 1995). Early results raised the possibility that 

the shift to lower concentrations was caused by disruption of the 

volcano’s plumbing by the 1868 earthquake (Tilling and others, 

1987a); subsequent work (Rhodes and Hart, 1995) shows the 
Island of Hawai‘i

156o 155o

20o

19o

MAP AREA

Mauna Loa

Mauna Kea

Kohala

Hualālai

Kīlauea

Kīlauea

Caldera

Hawai‘i Volcanoes

National Park

Mauna Ulu

Kanenuiohamo Nāpau

Crater

Pu‘u ‘Ō‘ō

Kupaianaha East Rift Zone

Pulama pali

Pulama pali

Flows emplaced 1982

Flows emplaced Jan 1983–Jun 1986

Flows emplaced 1790–1979

Flows emplaced Jul 1986–Feb 1992

Kalapana

Kapaahu

Kamoamoa

Makaopuhi

Crater

Flows emplaced Feb 1992–Jun 2007

Flows emplaced Jul 2007–Mar 2011

EXPLANATION

Halemaʻumaʻu vent

Pacific Ocean

0

0

2.5 5 KILOMETERS

1 2 3 MILES

Major road Minor road

National Park boundary

Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 247

decrease to have begun before 1868, and to have bottomed out 

in 1880. Levels rebounded somewhat by 1900 and have shown 

little variation through the 20th century. 

In contrast to Kīlauea, post-18th century Mauna Loa 

eruptions have consistently begun at the summit. The eruptions 

have then either continued immediately as rift zone eruptions 

(along the Northeast or Southwest Rift Zone, or both; see fig. 

2) or less commonly from vents on the north flank or resumed 

as a rift zone eruption following a hiatus (in which case a time 

delay of 1–2 years is typical), as summarized in Lockwood 

and Lipman (1987). This contrasts with post-1960 behavior at 

Kīlauea, where magma batches consistently appear first along 

the East Rift Zone and then at the summit, and only rarely and 

subsequently appear along the volcano’s Southwest Rift Zone.

Petrologic Overview of Kīlauea’s 

Ongoing Eruption, 1983–2011

The Kīlauea East Rift Zone has been erupting since 1983 

and continues to do so at the time of this writing. This eruptive 

era was ushered in by summit eruptions in April and September 

1982. Since March 2008, eruptive activity has expanded to 

include simultaneous summit and rift zone eruptions.

In January 1983, magma erupted through a downrift 

sequence of fissure vents extending 7.5 km from the west 

side of Nāpau Crater, past the eventual Pu‘u ‘Ō‘ō vent, and 

nearly to the eventual Kupaianaha vent (fig. 5). The details 

of eruption onset and early activity, including the highfountaining episodes, are presented in Wolfe and others (1987, 

1988). More recently, Heliker and others (2003) provided 

summaries of HVO’s monitoring-related research during the 

first 20 years of the Pu‘u ‘Ō‘ō-Kupaianaha eruption. Included 

in that volume, and briefly reviewed here, is a chapter that 

interprets the petrology and geochemistry of an extensive suite 

of well-quenched lava samples collected from East Rift Zone 

vents between January 1983 and October 2001 (Thornber, 

2003; data in Thornber and others, 2003a). 

Yet another 10 years of petrologic data gathering, during 

2001–11, have further enlightened our perspectives on the 

fundamental traits of near-continuous magma recharge and 

eruption associated with the shield-building stages of Hawaiian 

volcanism. The petrologic details of simultaneously erupted 

summit and rift zone lava provide opposing perspectives on 

magma complexities prior to rift eruption. Combined with 

geophysical and behavioral observations, petrologic monitoring 

at both early and later stages of the rift zone eruption has led to 

better understanding of how and where magma is transported, 

stored, erupted, and recharged within Kīlauea’s edifice. 

Figure 5.  Map showing the extent of lava flows of the ongoing Kīlauea East Rift Zone eruption (from January 1983 through March 2011).
248 Characteristics of Hawaiian Volcanoes

An overview of changes in eruption petrology is best 

facilitated with reference to the temporal variation of an 

MgO index. The weight percent magnesia (MgO) in bulk 

lava and matrix glasses tracks fluctuations in the eruption 

temperature and the character of magma tapped throughout 

that eruption (fig. 6). The lowest MgO values were recorded 

during the early stages of the eruption and again during (and 

shortly after) the brief Nāpau Crater eruption in January 1997. 

In both cases, petrologic studies confirm magma mixing 

shortly before eruption between cooler rift-stored magma 

pockets and hotter olivine-phryic magma at shallow depths 

(Garcia and others, 1989, 1992; Thornber and others, 2003b). 

Otherwise, as detailed by Thornber (2003), sparsely olivinephyric lava was erupted from well-established vents at Pu‘u 

‘Ō‘ō and Kupaianaha between 1985 and 2001. During those 

periods of prolonged and steady eruption, the consistent 

MgO thermometry difference of 30 oC between bulk lava and 

matrix glasses is explained by equilibrium olivine growth with 

cooling during summit-to-rift transit (Thornber, 2001). 

Thornber (2003) suggested that the persistent overall 

limits to the range of MgO cycles of steady-state eruption 

products during the period 1985–2001 reflect a geostatic 

balance between two distinct magmatic components that 

persist within the shallow volcanic edifice during prolonged 

intervals of recharge and eruption. One of them is a shallow 

recharge component with an upper MgO limit of 10 weight 

1,115

1,135

1,155

1,175

1,195

1,215

5

6

7

8

9

10

1982 1986 1990 1994 1998 2002 2006 2010

MgO content, in weight percent

MgO thermometry, in degrees Celsius

MgO (bulk rock)

Corrected glass temperature

percent. The other end member, at 6.8 to 7.0 weight percent 

MgO, is the most common of Kīlauea eruption compositions. 

This is a low-pressure multitectic composition for olivine 

tholeiite magma at which clinopyroxene and plagioclase 

crystallize, together with olivine (Helz and Thornber, 1987), 

and is therefore a thermodynamically favored open-system 

reservoir component that serves as a magmatic buffer during 

recharge (Thornber, 2003). The observed limits of cyclic MgO 

variation of this eruption (fig. 6) are within the slightly larger 

range of historical intracaldera summit lavas and define a 

persistent temperature range of olivine-saturated end-member 

magmas, which is presumably regulated by recharge of the 

shallow magmatic plumbing system during prolonged shieldbuilding eruptions. 

Chemical Signature of Recently Erupted Kīlauea 

Magma 

As with all eruptive sequences of Hawaiian olivine 

tholeiites, bulk-lava MgO variation diagrams for major and trace 

elements provide a first-order means of assessing magmatic 

conditions, such as low-pressure fractionation, mixing, or 

assimilation. Short-term cycles of temperature, MgO, and 

incompatible element concentrations are broadly consistent with 

repeated cycles of olivine fractionation from melts of ~10 to ~7 

Figure 6. Time plot showing temporal variation from 1983 to 2010 of MgO weight percent (left axis) and corresponding MgO glass geothermometer 

values (Helz and Thornber, 1987) (right axis) for bulk lava samples (black triangles) and average matrix glasses (red triangles). The latter are ventcorrected if sampled away from the vent by adding 0.9 oC or 0.05 weight percent MgO glass per kilometer of tube distance (Thornber, 2001). 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 249

weight percent MgO. As exemplified by CaO variations (fig. 7), 

olivine incompatible elements increase with progressive olivine 

fractionation, delineating an olivine-control trend or an olivinesaturated liquid-line-of-descent (LLD). The LLD for steady-state 

eruption products has shifted progressively, from 1985 to 1995, 

toward lower concentrations of incompatible elements. 

The trend of decreasing incompatible elements over time 

is apparent in ratios of highly to moderately incompatible 

elements, such as La/Yb, in olivine-normalized trace element 

patterns for successive intervals of steady-state eruption 

(Thornber, 2003) and is consistent with subtle variations of Sr, 

Nd, and Pb isotopes reported by Thornber and others (2003c), 

Garcia and others (1996, 2000), and Pietruszka and Garcia 

(1999a). This long-term decrease of incompatible-element 

ratios has been proposed by Garcia and others (1996, 2000) 

to represent changes in mantle-source conditions. In contrast, 

Thornber (2003) makes a case for a top-down rather than a 

bottom-up mechanism, attributing the change in chemical 

signature to progressive flushing of diminishing proportions of 

pre-1983 summit magma (maintained at multitectic conditions) 

by magma derived from a chemically uniform mantle source. 

Long-term summit deflation since 1983 is coincident with the 

long-term geochemical trends seen in subsequent products 

of continuous eruption, suggesting that progressive summitreservoir depletion may be responsible. In this scenario, magma 

derived from a uniform mantlesource has apparently flushed 

older resident magma from theshallow edifice and could now 

completely occupy the shallow magmatic plumbing system. 

This idea was further supported when new magma began 

erupting at the summit in March 2008.

9.5

10.5

11.5

CaO content, in weight percent

9.5

10.5

11.5

10 9 8 7 6 5

MgO content, in weight precent

E55 LLD

OLV+SP CPX+PLG

10 9 8 7 6 5

Episode 55 Episode 56 Episode 57

Episode 58 Summit 2008–2011

1983–2011

1983–2011 1998–1999 1999–2001

2005–2007 2001–2005

CaO content, in weight percent

MgO content, in weight precent

A B

E55 LLD

Post-2001 Eruption Petrology: Unexpected 

Changes Associated with an Increase in Magma 

Supply 

Between October 2000 and April 2001, the petrology 

of steady-state Pu‘u ‘Ō‘ō-Kupaianaha lava underwent a 

fundamental and long-term change following the steady 

decline in bulk MgO from a 1998 high of 9.5 down to around 

7.5 weight percent (fig. 6) and amidst the onset of the first 

prolonged summit inflation in more than 20 years (Poland 

and others, 2012). For the first time in more than 20 years, 

the steady-state rift zone eruption was no longer tapping an 

olivine-only shallow magma source. The preeruptive condition 

changed to a more complex comagmatic mixture of hotter 

olivine-only magma and a cooler magma with clinopyroxene 

(±olivine, ±plagioclase; Thornber and others, 2007). All post2001 low-temperature lava is characterized by a disequilibrium 

low-pressure phenocryst assemblage. Olivine (Fo80.5–81.5) 

coexists with phenocrysts of lower temperature clinopyroxene 

(±plagioclase, ±Fe-rich olivine). Mixing between hotter and 

cooler magmas is texturally documented by complex pyroxene 

zoning and resorption, along with olivine overgrowths of 

resorbed pyroxene and plagioclase grains. 

The prolonged decline in the MgO index bottomed out 

in 2004 at 6.8 weight percent. This value corresponds to the 

low-pressure multitectic composition for Kīlauea magma 

and is the lowest sustainable MgO value during periods of 

continuous recharge and eruption (Thornber, 2003). The bulk 

MgO of olivine-clinopyroxene-phyric lava from vents at Pu‘u 

Figure 7.  Graphs of CaO versus MgO contents in Kīlauea lava erupted from 1983 through 2011. A, Analyses from selected time periods from 

1998 to 2005, highlighted in color to emphasize the episode 55 trend of decreasing MgO with time. B, Analyses from selected time periods 1997 to 

2007, highlighted to contrast episode 56, episode 57, and episode 58 lava compositions, along with 2008–11 summit juvenile tephra compositions. 

E55LLD is episode 55 liquid-line-of-descent, as shown in figure 2 of Thornber (2003), which shows episode 48–55 changes and provides additional

explanatory text. The arrows indicate the change in magma compositions resulting from low-pressure fractional crystallization of (1) olivine 

and spinel (OLV+SP) at high temperatures (from recharge magma with up to 10 weight percent MgO) and (2) clinopyroxene and plagioclase

(CPX+PLG) at lower temperatures (from isolated magma with ~6.8 weight percent MgO). The fulcrum in this plot is the 5-phase multitectic point for

shallow Kīlauean magma that occupies cooler zones within the actively recharged reservoir. During prolonged recharge and eruption, variations

along the olivine-controlled liquid-line-of-descent (LLD) with time are affected by proportional mixing of these perpetually present components.
250 Characteristics of Hawaiian Volcanoes

‘Ō‘ō continued to hover around 7.0 to 7.3 weight percent until 

the June 19, 2007, eruption at Kanenuiohamo (episode 56, the 

Father’s Day activity), a few kilometers uprift of Pu‘u ‘Ō‘ō, 

of the hottest and most primitive olivine-only lava seen since 

1998. This occurred at the culmination of the 4-year summit 

inflationary cycle and provided a glimpse of the relatively 

primitive component that is perpetually present during 

magmatic recharge of the shallow edifice. The relatively 

primitive magma erupted at Kanenuiohamo lies along the 

same LLD as all post-1995 lava (see episode 56 in fig. 7B) 

and is a petrologic match for the hotter, olivine-phyric end 

member of simple binary mixing between recharge magma 

and stagnant crystal-saturated magma. 

When the East Rift Zone eruption revived downrift 

of Pu‘u ‘Ō‘ō one month after the Kanenuiohamo eruption, 

the July 2007 fissures once again issued low-MgO, olivineclinopyroxene-phyric lava. MgO of rift zone lava increased 

steadily to 8.0 weight percent 5 months after the onset 

of Kīlauea’s summit eruption in March 2008. Increasing 

eruption temperature with time accompanied compositional 

migration back up the olivine-controlled LLD (fig. 7B) and, 

by August 2008, the preeruptive condition retransitioned to 

olivine-only for another 10.5 months. Meanwhile, bulk MgO 

began to decline steadily in September 2008 and breached 

7.5 weight percent in mid-June 2009 as clinopyroxene 

phenocrysts began to reappear in rift lava. The MgO decline 

bottomed out in October 2009 near 7.1 weight percent, where 

it remained through December 2010. 

The hybrid lava erupted on the East Rift Zone from 2001 

to 2010 is a mixture of two end-member components within 

the shallow magmatic regime that feeds the eruption. The 

high-MgO and high-temperature magma is equivalent to that 

which was sparsely erupted at Kanenuiohamo in June 2007 

(and which is currently being erupted at the summit vent). 

The cooler multitectic component of this magma mixture 

persists in pockets or perhaps throughout the active shallow 

plumbing system; it has the same time-depleted incompatibleelement signature (for example, La/Yb, Zr/Y; C.R. Thornber, 

unpublished data) as the high temperature component, which 

suggests that this component is maintained in a reservoir (or 

dike network) open to near-continuous magmatic recharge. 

Furthermore, the phenocryst phase relations, textures, and 

compositions, along with sulfur concentrations of melt 

inclusions, define a dynamic shallow mixing environment that 

is driven by magmatic recharge into a zone of denser, cooler, 

degassed, and partially crystalline shallow magma (Thornber 

and others, 2010). 

The long-term petrologic monitoring effort during the 

East Rift Zone eruption from 1983 to present suggests that 

such quasi-stagnant magma zones are perpetually flushed 

during continuous eruption of lower-MgO hybrid lava. 

Geophysical and gas monitoring data are consistent with 

an increase in magma supply during the post-2001 interval 

(Poland and others, 2012). Counterintuitively, the apparent 

increase of magma supply to the edifice coincides with a 

pulse of low-MgO hybrid lava through the East Rift Zone 

that continues to the end of 2010 (fig. 6). This inverse 

relation suggests that cooler crystal mush zones within the 

shallow edifice were being flushed more efficiently during 

the period of increased magma supply.

Summit-to-Rift Magmatic Continuity in the 

Ongoing Eruption

Throughout the duration of this epic East Rift Zone 

eruption, physical connectivity between a summit magma 

reservoir and the rift zone vents has been inferred by 

correlations of geophysical flexure at the summit to 

surges and lulls in activity and changes in the MgO index 

(Thornber, 2001, 2003; Thornber and others, 2003b). Since 

the onset of summit eruption in 2008, petrologic analysis 

of lava simultaneously erupted at both ends of the eruptive 

plumbing system has provided a unique and unequivocal 

demonstration of magmatic continuity between the shallow 

summit and throughout the 18-km-long rift zone conduit. 

Trace-element signatures for glasses, glass inclusions, 

and bulk lava in 2008 rift zone and summit samples are 

indistinguishable (Rowe and others, 2009; Thornber and 

others, 2010). Olivine-phyric lava from the 2008-to-present 

Halema‘uma‘u vent lies along the same incompatibleelement-depleted LLD as all post-1995 Kīlauea lava 

(fig. 7B). This has significant implications for the overall 

magmatic condition of the Kīlauea edifice. 

Although earthquake and summit deformation patterns 

have long implied a shallow reservoir of magma beneath 

Halema‘uma‘u Crater (see, for example, Poland and others, 

this volume, chap. 5), its size, age, longevity, and physical 

characteristics have been topics of considerable speculation 

(Pietruszka and Garcia, 1999b; Thornber, 2003). Many 

assumed it would likely comprise the stagnant magma left 

from past summit eruptions in 1982, 1974, or perhaps even 

decades earlier. 

Our recent petrologic evidence shows that the magma 

of the long-term East Rift Zone eruption has flushed out 

older resident summit magma and occupies the entire 

shallow magmatic plumbing system (Thornber, 2003). 

Furthermore, the persistent mixed-crystal assemblage in rift 

zone lava attests to the perpetual presence of magma mush 

along cooler margins of the active magma pathways and in 

cavities, cracks, or sills within the shallow volcanic edifice. 

A zone of mixing between hotter magma that recharges the 

edifice and cooler magma mush that resides within it may 

exist in the summit reservoir, adjacent to the East Rift Zone 

conduit, but it is also possible that preeruptive magma mixing 

occurs all along the transport path to the East Rift Zone vent. 

The persistently low-MgO lava must reflect a steady-state 

condition in which a volumetrically significant open reservoir 

within the shallow edifice is being perpetually refreshed by 

recharge magma. This dynamic condition of variable recharge 

with two simultaneous eruptions at different sites underlines 

the open-system nature of Kīlauea’s magmatism. 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 251

Investigations of Hawaiian Lava Lakes 

Investigations of lava lake activity and behavior have been 

an important component of HVO’s scientific output for the entire 

century of the observatory’s existence. Two distinct features have 

been called lava lakes: the first are active lava lakes underlain by 

a magma column (as at Halema‘uma‘u before 1924 and during 

1952, 1967–68 and 2008–present; Mauna Ulu during 1969–74; 

and Pu‘u ‘Ō‘ō, especially during 1983–86), and the second 

are closed-system lava lakes, produced when lava ponds in 

preexisting pit craters and is no longer in contact with its source. 

Active Lava Lakes

The most famous Hawaiian lava lake occupied 

Halema‘uma‘u Crater for most of the 19th and early 20th centuries, dominating the attention of early observers, as well as that of 

Dr. Jaggar in HVO’s early years (Jaggar, 1917a, 1917b; Bevens 

and others, 1988). In spite of the hazardous environment, Jaggar 

collected a small suite of glassy spatter samples discussed earlier 

but was not able to collect the more viscous material he referred 

to as “bench lava,” which surrounded the more fluid parts of the 

lava lake; however, he described seeing incandescent slabs of 

bench lava spall off the sides, buckling and then disintegrating to 

glowing talus (Jaggar, 1917b) during draining of the lava lake. 

His vivid description suggests that this material was similar to 

the melt-rich mush encountered in various historical lava lakes. 

The mush (fig. 8) consists of a very tenuous crystal network in a 

matrix of as much as 45 volume percent melt. Although this mush 

has been drilled and recovered, bore holes in it collapse within 

days (Helz and Wright, 1983), much as the bench lava collapsed 

gradually after withdrawal of the active lava lake. 

More recently, significant lava lake activity in 

Halema‘uma‘u occurred during the 136-day 1952 eruption 

(Macdonald, 1955) and for 18 months in 1967–68 (Kinoshita 

and others, 1969). The current summit activity (2008 to present) 

differs from these earlier eruptions in that the top of the magma 

column is visible within a small pit crater 160 m in diameter (T. 

Orr, written commun., 2012) against the wall of Halema‘uma‘u 

Crater (location shown in fig. 5) rather than occupying much 

of the crater’s floor. Webcam images of the magma pool show 

similar magma surface patterns to those observed on the 

magma column at Mauna Ulu (Duffield, 1972). The relatively 

small area is similar to the active lava lake that existed at the 

beginning of HVO’s life, which occupied only a small part of 

Halema‘uma‘u (Jaggar, 1917a, 1917b).

The style of activity of this lava lake generates mostly 

spatter, but microbeam techniques now allow extensive 

analysis of even tiny amounts of material in support of 

petrologic study. Samples from the 2008–present summit 

eruption are similar to Halema‘uma‘u scoria and spatter 

sampled over the past 100 years—they contain few crystals, 

and both glass and bulk compositions have a limited range 

of MgO (figs. 3 and 9), which lies at the end of Kīlauea’s 

olivine control line (as discussed in the previous section). 

500 μm

6

6.5

7

7.5

8

1,135

1,140

1,145

1,150

1,155

1,160

1,165

1,170

1,175

2008 2009 2010 2011

MgO glass temperature, in degrees Celsius

MgO glass composition, in weight percent

Maximum value

Average value

Figure 8. Photomicrograph showing groundmass in partly molten 

core from near the limit of drillability of the crust of the Kīlauea Iki lava

lake. Brown glass makes up 40–45 percent by volume of the core. 

Crystals include phenocrystic, microphenocrystic, and groundmass 

olivine, plus groundmass augite (green) and plagioclase. The crystals 

tend to nucleate on each other, leaving more melt-rich areas between 

the clusters and chains of crystals. 

Figure 9. Graph showing variation of MgO content and glass 

quenching temperature of tephra from the ongoing (as of 2014) 

Kīlauea summit eruption versus time. Maximum values, shown as

red dots, provide the best indication of changing magmatic conditions 

in the summit reservoir. Average values with standard deviation are 

indicated with a black dot and vertical bar. Data from Wooten and 

others (2009) and C.R. Thornber (unpublished data).
252 Characteristics of Hawaiian Volcanoes

Closed-System Lava Lakes and Sills

The second class of lava lakes, not fed from below but 

formed by lava flowing into preexisting pit craters (as described in 

Richter and others, 1970; Peck and Kinoshita, 1976; Wright and 

others, 1968), has received major attention at HVO, and study of 

these bodies has contributed greatly to our understanding of the 

crystallization and fractionation behavior of tholeiitic magma. 

Many lava lakes have formed in the 20th century (fig. 10); three 

of these, the 1959 Kīlauea Iki, 1963 ‘Alae, and 1965 Makaopuhi 

lava lakes, have been studied extensively. Older analogues include 

a prehistoric Makaopuhi lava lake (Moore and Evans, 1967; Evans 

and Moore, 1968), the Uēkahuna laccolith (Murata and Richter, 

1961), a prehistoric lava lake at Mauna Loa that has been recovered in fragments (McCarter and others, 2006), and similar bodies 

observed in the walls of Moku‘āweoweo (Sparks, 1990). Table 2 

summarizes some of the properties and data sets available for the 

lava lakes and sill at Kīlauea. 

Kīlauea Iki, ‘Alae, and the 1965 Makaopuhi lava lakes have 

been the focus of elaborate field studies, including (1) observations 

of their filling and surficial subsidence, (2) repeated core drilling, 

with sample collection and petrologic and chemical analysis, and 

(3) field determinations of temperature, melt viscosity, oxygen 

fugacity, seismic, electrical and other geophysical properties. 

Wright and others (1976) summarized the range of results for all 

these lakes, and further results for Kīlauea Iki were summarized 

in Helz (1987b) and Barth and others (1994). The sections below 

compare selected results from the historical lava lakes with those 

for the prehistoric Makaopuhi lava lake and the Uēkahuna laccolith. 

‘Alae

(filled)

Aloi

(filled)

Hiiaka

Pauahi

Makaopuhi

Kīlauea Iki

Keanakāko‘i

Kīlauea

Caldera

Halema‘uma‘u

Crater

Mauna Ulu

155° 18'

19° 24'

19° 26'

19° 22'

155° 16' 155° 14' 155° 12' 155° 10'

HVO

0

0

2 KILOMETERS

1 MILE

1

0.5

Controls on Whole-Rock Compositions in Lava Lakes 

and Sills

Petrologic studies of lava lakes at Kīlauea have 

documented how tholeiitic basalt differentiates as it cools 

and crystallizes by tracking changes in bulk composition and 

interstitial melt composition. The crystallization behavior and 

liquid lines of descent for the relatively olivine-poor 1965 

Makaopuhi and 1963 ‘Alae lava lakes (MgO=7.5–8.2 weight 

percent; table 2) were described by Wright and Okamura 

(1977), Peck and others (1966), and Wright and Peck (1978), 

while the liquid lines of descent and extreme differentiation 

products of Kīlauea Iki were described in Helz (1987b). 

Unfortunately, the ‘Alae and 1965 Makaopuhi bodies were 

covered by lavas of the Mauna Ulu eruption (as was the cliff 

face through the prehistoric Makaopuhi lava lake), so they 

have not been studied further. 

Figure 11 shows magnesia variation diagrams for selected 

oxides (SiO2, Al2O3, and CaO) for Kīlauea Iki (MgO=15.5 

weight percent; Wright, 1973) and the Uēkahuna laccolith 

(MgO=16.0 weight percent; Murata and Richter, 1961), both 

of which are still accessible. The figure also includes analyses 

from the prehistoric Makaopuhi lava lake (MgO=9.67 weight 

percent; Moore and Evans, 1967; Gunn, 1971), because the 

analyzed section includes its olivine-enriched zone, allowing 

direct comparison with the olivine-rich zones in the picritic 

bodies. The three diagrams chosen are sufficient to show the 

effects of crystallization of olivine, augite, and plagioclase on 

the range of whole-rock compositions observed. 

Figure 10. Index 

map of Kīlauea’s

summit and upper 

East Rift Zone

showing locations 

of recent lava 

lakes (red), plus 

the prehistoric 

Makaopuhi lava 

lake (gray). Of the 

lava lakes shown, 

only ‘Alae (1963), 

Makaopuhi (1965), 

Kīlauea Iki (1959),

and the prehistoric 

Makaopuhi lava 

lakes have been 

extensively studied. 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 253

Table 2.  Properties and data available for lava lakes and sills at Kīlauea Volcano.

[m, meters; est., estimated; max., maximum; XRF, X-ray fluorescence; INAA, instrumental neutron activation analysis; EDXRF, energy-dispersive X-ray fluorescence; PGE, platinum group elements; px, pyroxenes; Re, rhenium]

 1963 ‘Alae 1965 Makaopuhi 1959 Kīlauea Iki Prehistoric 

Makaopuhi

Uēkahuna 

laccolith

Thickness of body (m) 13.7 83 (est.) 130–135 (est. 

from 1988 holes)

69 (cliff) to 

120 (est.)

8.5 (sampled) 

to 27 (max.)

Average bulk MgO 

content (weight percent) 7.6 8.2 15.5 9.7 16

Present 

accessibility

Covered by Mauna 

Ulu shield

Covered by 

Mauna Ulu lavas Still exposed Cliff covered by Mauna Ulu lavas Still exposed

Temperature profiles 1963–67, through 

body

1965–1969, to 

11.6 m depth

1960–1988, to 76 m 

(center), through body 

at edge (95.4 m)

None None

Thermal modeling Yes No Yes Yes No

Samples Drill core* 

through body

Drill core* 

to 20 m

Drill core* to 

114.6 m, through body’s 

(95.4 m) north edge

Outcrop samples 

from cliff (69 m)

Outcrop samples 

from cliff (8.5 m)

Major element chemistry 33 classical 72 classical 194 classical

64 XRF

13 classical 

8 XRF 8 classical

Reference(s) 

(majors) 

Wright and Peck 

(1978)

Wright and 

Okamura (1977)

Richter and Moore (1966); 

Helz and others (1994); 

Helz and Taggart (2010)

Moore and Evans 

(1967); Gunn (1971)

Murata and 

Richter (1961)

Trace element 

 chemistry 14 spectrographic None 72 INAA + 57 EDXRF; 

19 PGE + Re 8 XRF None

References (traces)

Wright and Peck 

(1978) -- Helz (2012); Pitcher and 

others, (2009)

Gunn (1971)

--

Electron microprobe 

phase chemistry One residual glass 3 px, 2 residual 

glasses, oxides

Olivine (this report), 

glasses (Helz, 1987b), 

olivine + chromite 

(Scowen and others, 1991)

All crystalline phases 

(Moore and Evans, 

1967; Evans and 

Moore, 1968)

--

*

 Drill core available through the National Museum of Natural History, Smithsonian Institution, Washington, D.C.
254 Characteristics of Hawaiian Volcanoes 468101214 CaO content, in weight percent 46810121416 Al2O3 content, in weight percent 424446485052545630

252015105

0

SiO2 content, in weight percent

MgOcontent,inweightpercent

Figure 11. Graphs 

showing whole-rock 

contents of SiO

2

, Al

2

O3,

and CaO, plotted against 

MgO content (all quantities 

in weight percent), for the 

prehistoric Makaopuhi 

lava lake (Moore and 

Evans, 1967; Gunn, 1971; 

red squares), and for the 

Uēkahuna laccolith (Murata

and Richter, 1961; blue 

squares). Also shown for 

comparison are data from 

1967–88 core from Kīlauea

Iki lava lake, shown as 

black dots (Helz and others, 

1994; Helz and Taggart, 

2010). Rock compositions 

that contain less than 6 

weight percent MgO are 

segregation veins and 

dikelets and are present 

in all three bodies. Arrows 

indicate vertical olivine-rich 

bodies (vorbs) in Kīlauea

Iki (inverted black triangles) 

and in the prehistoric 

Makaopuhi lava lake. The 

colored field shows the

range of compositions of 

the 1959 eruption samples 

(Murata and Richter, 1966) 

and 1960–61 drill core 

from Kīlauea Iki lava lake

(Richter and Moore, 1966). 

The sides of this field are

olivine control lines for 

Fo86–87. See text for further 

discussion.
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 255

For most sets of samples, olivine control is the dominant source of bulk compositional variation; in each case, 

this has been attributed to redistribution of phenocrystic 

olivine within the body (Murata and Richter, 1961; Moore 

and Evans, 1967; Helz and others, 1989). Data for 1967–88 

Kīlauea Iki core samples lie mostly within the colored field 

that is defined by the compositional ranges of the 1959 

eruption samples and early (1960–61) core samples in SiO2

versus MgO. Those for the two prehistoric bodies lie parallel 

to the field at slightly higher SiO2 contents. This distribution 

shows that, in all cases, the olivine composition was Fo86–87

at the time of redistribution. The array of points in Al2O3 

versus MgO is also parallel to the Fo86–87 control lines for all 

three bodies, showing that plagioclase was not materially 

involved as a fractionating phase, except for samples with 

MgO<7 weight percent, as has been widely observed for 

other Kīlauea lavas (Wright, 1971). 

The patterns for CaO versus MgO are more complex: 

highly magnesian Kīlauea Iki samples are enriched in CaO, 

relative to the 1959 eruption samples, while the Uēkahuna 

samples run lower, but parallel to, the Fo86–87 control lines. 

However, samples with 7–11 weight percent MgO, both 

from Kīlauea Iki and from the prehistoric Makaopuhi lava 

lake, lie on a slope flatter than olivine control, which gives 

the appearance of olivine+augite fractionation. Augite does 

not crystallize until melt MgO=7.2 weight percent (Helz 

and Thornber, 1987), so the flattening is not produced by 

fractionation of olivine+augite from liquids having the bulk 

composition of the samples, but by migration of interstitial 

liquid from within the olivine-enriched zone to upper parts 

of the lake (Moore and Evans, 1967; Helz and others, 1989). 

In Kīlauea Iki, this process clearly occurs after settling of the 

olivine phenocrysts. The Uēkahuna laccolith does not show 

signs of this process, probably because it is much thinner 

(table 2). 

The lava lakes and laccolith discussed here contain 

internal differentiates in addition to the dominant, olivinephyric matrix rock. Figure 11 includes analyses of segregation veins from Kīlauea Iki and prehistoric Makaopuhi 

lava lakes, and one analysis of aphanitic dikelets from the 

Uēkahuna laccolith. All have similar compositions, with bulk 

MgO contents of 4–6 weight percent and low CaO and Al2O3 

relative to the olivine-controlled samples. They correspond 

to melts with liquidus temperatures of 1,100–1,135 °C 

(Wright and Okamura, 1977; Helz, 1987b; Helz and others, 1989) and have been interpreted as filter-pressed liquids 

from within a mush crystallizing olivine+augite+plagioclase 

(Moore and Evans, 1967; Helz, 1980, 1987b; Helz and others, 1989). Similar coarse differentiates were reported by 

McCarter and others (2006) for a disrupted lava lake erupted 

as blocks on Mauna Loa. 

Kīlauea Iki lava lake also contains vertical, cross-cutting pipes of varying character, enriched in segregation-vein 

liquid, which have been interpreted as diapir tracks related 

to the formation of segregation veins. Most abundant are the 

vertical olivine-rich bodies (“vorbs”) observed in all cores 

from Kīlauea Iki in the depth range 18–58 m from 1975 

onward (Helz, 1980; Helz and others, 1989). Other vertical 

structures found at and below 78-m depth resemble the melt 

chimneys of Tait and others (1992) and are interpreted as 

points of escape of plumes of vesicles+melt from within the 

lower crust of the lava lake (Helz, 1993; Helz and Taggart, 

2010). Both sets of vertical structures, when quenched in a 

partly molten state, are variably enriched in Fe-rich olivine (Fo78–80; Helz and others, 1989) and highly enriched in 

segregation-vein melt and vesicles relative to the adjacent 

matrix (fig. 12A,B,D). 

Moore and Evans (1967) noted a population of “pipe-like 

masses, rich in olivine” in the prehistoric Makaopuhi lava 

lake, one of which is shown in figure 12C. The one analysis of such a body is similar to the vorbs in Kīlauea Iki: it is 

enriched in segregation-vein liquid and in Fe-rich olivine, so 

(like the vorbs) is low in SiO2, Al2O3, and CaO relative to the 

matrix rock, although high in MgO, as shown in figure 11. The 

position of these pipes low in the cliff face and their relatively 

short (0.15 m) height (Moore and Evans, 1967), however, 

makes them more like the deeper vertical structures found 

in Kīlauea Iki (for example, fig. 12D). McCarter and others 

(2006) describe an “open-textured dunite” from the Mauna 

Loa body as a possible vorb, but its relationship to the rest of 

the dismembered lava lake is unknown.

Appearance, Distribution, and Abundance of 

Segregation Veins

All Kīlauea lava lakes contain segregation veins. 

Though variously described as diabase (Peck and others, 

1966), ferrodiabase (Helz, 1980, 1987b), or mafic micropegmatite (Moore and Evans, 1967), they are very consistent 

in chemical and petrographic character. Figure 13A shows a 

partly molten segregation vein from Kīlauea Iki; it is olivinefree and consists of pyroxenes, plagioclase, Fe-Ti oxides, and 

brown glass, with all crystalline phases 1–2 orders of magnitude coarser grained than their equivalents in the underlying 

olivine-phyric matrix. The subhorizontal, relatively sharp 

lower contact is typical. The 13.4-m thick 1963 ‘Alae lava 

lake contains one such vein (0.4 m thick, or 3 percent of the 

lake’s thickness), and several were encountered in drilling 

the upper 20 m of the 1965 Makaopuhi lava lake. The disaggregated lava lake described by McCarter and others (2006) 

also contains abundant coarse-grained internal differentiates 

like the segregation veins described here. 

The cliff face through the prehistoric Makaopuhi lava 

lake and the extensive drilling program in Kīlauea Iki give 

the most complete picture of the extent of these internal 

differentiates. Figure 14 shows the array of segregation veins 

exposed in a section of the cliff face that cut the prehistoric 

Makopuhi lava lake. The veins are coarse-grained, with 

relatively sharp boundaries, and typically show lateral extents 

20–100 times greater than their thickness. For comparison, 

figure 15 shows the depth and thickness of all segregation 
256 Characteristics of Hawaiian Volcanoes

Figure 12. Photomicrographs of drill-core samples 

from Kīlauea lava lakes. A, Contact between 

vertical olivine-rich body, or “vorb” (left), and 

normal olivine-phyric matrix (sample KI75-1-141.5, 

T=1,095 °C) in plane light, Kīlauea Iki. Horizontal

cut through vertical contact; field of view is 3.2

cm across. B, Irregular vertical contact between 

melt-rich chimney (right) and normal olivine-phyric 

matrix (KI81-5-258±1, T=1,130 °C) with nicols partly 

crossed, Kīlauea Iki. Melt chimney contains coarser

plagioclase and more melt and vesicles than matrix. 

Vertical cut; field of view is 2.2 cm across. C, Contact 

between normal matrix (right) and “pipe-like mass 

rich in olivine” from the prehistoric Makaopuhi lava 

lake (MP-200) with crossed nicols. Field of view is 3 

cm across. D, Small plume enriched in olivine, melt, 

and vesicles (gray) in normal olivine-phyric matrix 

(KI88-1-275.1, T=1,100 °C) in plane light, Kīlauea

Iki. This is a variant melt chimney. Vertical cut, up is

to the right. Field of view is 2.5 cm across.

Figure 13. Photomicrographs 

showing contacts in drill-core 

samples from Kīlauea Iki lava

lake. A, Subhorizontal contact of 

segregation vein with underlying 

olivine-phyric host in sample 

KI79-3-164.5 (T=1,068 °C) in 

plane light. Vertical cut; field of

view is 1.7 cm across. B, Contact 

between black, quenched ooze 

(upper part of image) and olivinephyric wall rock in borehole KI76-

2, recovered at a depth of 141.7 

ft (43.2 m), with nicols partly 

crossed. Vertical cut through core;

contact is vertical, up to the right. 

Field of view is 1.7 cm across.

A

B

C

D
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 257

Distance, in feet

Distance below top of cliff, in feet

Distance below top of lava lake, in feet

90

100

110

120

40

50

60

70

0 10 20 30

12" 1" 3"

6"

18"

Sketch by J.G. Moore veins encountered in the 1967–81 Kīlauea Iki drill 

cores. The overall pattern is similar to that observed 

in the prehistoric Makaopuhi cliff face. The individual 

segregation vein with the greatest inferred lateral 

extent (highlighted in red in fig. 15) is found at a depth 

of 45.72 m (150 ft) below the lake surface; it occurs in 

10 cores and has the same thickness (~0.3–0.4 m) over 

a lateral distance of more than 150 m.

These stacks of internally produced sills make up 

an appreciable fraction of the upper crust of the lava 

lakes. That fraction is 6.2 to 11.5 percent by volume 

in Kīlauea Iki above ~53 m (fig. 15); the frequency 

of such sills in the prehistoric Makaopuhi lava lake 

(fig.14) is similar. This corresponds to 2.5–4.5 percent 

of the thickness of the entire body, which compares 

closely with the 3 percent observed in the much 

thinner ‘Alae lava lake. Equivalent liquids would 

make up 45–50 percent by weight of Kīlauea Iki’s 

bulk composition (Wright, 1973), so the amount of 

liquid segregated into veins is an order of magnitude 

less abundant than in the bulk composition.

The crosscutting aphanitic dikelets in the north 

body of the Uēkahuna laccolith have the composition 

of segregation veins (fig. 11) but differ in grain size 

and in their relation to the host body (Murata and 

Figure 14. Sketch of segregation veins exposed in the cliff through the 

prehistoric Makaopuhi lava lake. This sketch was made by J.G. Moore while 

examining the cliff in a bosun’s chair during the summer of 1962. The exposure

was covered in 1969, when lava from Mauna Ulu filled the west pit to the level

of the top of the cliff. One foot equals 0.3 m; 1 inch (1") equals 0.4 cm.

A B
258 Characteristics of Hawaiian Volcanoes

7.3% 6.2% 8.7% 7.3% 9.4% 7.2% 11.5% 8.3% 6.6% 10.1% 4.8%

Depth below lava lake surface, in meters

0

102030405060

KI81-6

KI79-3

KI79-4

KI75-1, 75-1R

KI67-3

KI81-1

KI81-5

KI81-7

KI81-2

KI81-4

KI79-2

KI76-1

KI76-2

KI79-1

KI81-3

KI79-6

KI67-1

KI75-2

KI75-3

KI79-5

KI67-2

73.2 m 77.7 m 77.7 m 47.2 m 24 m 16.4 m 5.5 m 13.1 m

5.5 m

18.3 m

Figure 15.  Correlation chart showing the depths and thicknesses of all segregation veins encountered in 1967–81 drill core in Kīlauea Iki lava lake; core identification numbers given at top.The solid black intervals show all veins thicker than 1 cm. Vertical thicknesses are to scale; horizontal distances between drill holes are not to scale, but distances between some cores indicatedin meters (m). Dashed lines show correlations of segregation veins between drill holes based on depth. One particularly widespread vein at a depth of 45.7 m (150 ft) is highlighted in orange. Numbers below the columns give the fraction of segregation veins in the core between the surface of the lake and the deepest segregation vein. Other details as in Helz (1980, fig. 7).
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 259 Depth below surface of Kīlauea Iki lava lake, in meters

0

20

40

60

80

100

120

140

Temperature, in degrees Celsius

0 200 400 600 800 1,000 1,200

1961

1962

1967

1975

1979

1981

1988

Estimated position of lower contact

Thermocouple

Magnetite-ilmenite

Glass

Thermal maximum

Contact

0

1

2

3

4

5

6

7

8

9

10 11

Temperature, in degrees Celsius

0 300 600 900 1,200

Depth below surface of ‘Alae lava lake, in meters

0

5

10

A B

Figure 16. Graphs showing temporal evolution of vertical temperature profiles through Kīlauea lava lakes. A, Computed (continuous curves) 

and observed (symbols) temperature profiles in the 1963 ‘Alae lava lake during the 4 years following its formation, after figure 4 in Peck and

others (1977). Data for individual temperature profiles in Peck (1978). B, Variation of temperature with depth below the surface of Kīlauea Iki

lava lake as a function of time for drill holes near the center of the lake. Temperatures estimated by glass geothermometry (Helz and Thornber, 

1987), from coexisting magnetite-ilmenite compositions (Buddington and Lindsley, 1964), or measured by thermocouples lowered down the hole, 

as indicated. Shapes of curves below ~100 m are constrained by (1) an assumed depth of 135 m (based on results for drill hole KI88-2; see

discussion in Helz, 1993) and (2) a lower contact temperature of ~600 °C, based on an initial lake temperature of ~1,190 °C (Helz, 2009).

Richter, 1961). They are not sill-like, and their aphanitic grain 

size resembles that of oozes recovered in boreholes in Kīlauea 

Iki (and other lava lakes), an example of which is shown in 

figure 13B. The oozes are crystal-poor melts, usually derived 

from the mush below the base of the drillable crust; they 

quench rapidly as they move into open boreholes. A possible 

origin for the aphanitic dikelets in Uēkahuna is that they 

represent interstitial melt released from partly crystalline mush 

during shearing of the host body. Exposure of the sheared 

face to surface conditions during caldera formation would 

have quenched the resulting melt stringers to an aphanitic 

texture. If so, the Uēkahuna body would have resembled the 

incandescent, truncated intrusives observed by Jaggar and 

Finch (1924) in the wall of Halema‘uma‘u after the 1924 

eruption. This model would constrain the age of the laccolith 

to be only slightly older than the age of formation of Kīlauea’s 

current caldera (1470–90 C.E.; Swanson and others, 2012), 

making it approximately coeval with the ‘Ailā‘au activity 

(Clague and others, 1999).

Thermal History of Lava Lakes 

‘Alae and Kīlauea Iki lava lakes were monitored 

thermally for most of their crystallization history. Cooling 

of the prehistoric Makaopuhi lava lake was modeled (Moore 

and Evans, 1967) using a simple conductive cooling model 

(Jaeger, 1961). Peck and others (1977) included the effects of 

rainfall in their modeling of the cooling of ‘Alae lava lake; 

with this modification, they were able to match observed field 

measurements of temperature very closely (fig. 16A). Ryan 

(1979) reported results of two-dimensional thermal modeling 

of Kīlauea Iki, but the grid used was coarse and the results not 

directly comparable to those reviewed here. 

A composite of temperature profiles for the central part 

of Kīlauea Iki, obtained for 1961 through 1988, is shown in 

figure 16B. Above 100-m depth, the thermal evolution of the 

lava lake is well documented. Constraints on the curves below 

100 m are provided by (1) an estimate of the depth of the 

lava lake, which is generally about 130 m (Barth and others, 

1994), but probably slightly deeper at the center of the lake 

(Helz, 1993), and (2) the assumption that the lower contact 

temperature is half of the initial temperature of 1,190 °C, as 

expected for conductive cooling (Jaeger, 1961). The observed 

curves compare closely with both observed and model 

temperature profiles through the much thinner ‘Alae lava lake, 

which cooled to subsolidus temperatures in about 9 months 

(Peck and others, 1977; Peck, 1978). Based on extrapolation 

of these data, the interior of Kīlauea Iki reached subsolidus 

temperatures in the mid-1990s, giving a cooling time of 

about 35 years. This solidification time is similar to the 30 
260 Characteristics of Hawaiian Volcanoes

Time interval (years) Temperature interval (°C) Cooling rate

1959 to 1967 (84.5)* 1,190 to 1,090* 12.5 °C/year

1959 to 1975 (145.1)* 1,190 to 1,120* 4.4 °C/year

1959 to 1976 (150)* 1,190 to 1,115* 4.4 °C/year

1959 to 1979 (203.6)* 1,190 to 1,141* 2.4 °C/year

1959 to 1981 (219.8)* 1,190 to 1,140* 2.3 °C/year

1959 to 1988 (266.6)* 1,190 to 1,104* 3.0 °C/year

1979 (203.6) to 1981 (205.4) 1,141 to 1,135 3.0 °C/year

1981 (269.9) to 1988 (266.6) 1,128 to 1,106 3.1 °C/year

1981 (299.9) to 1988 (300) 1,099 to 1,081 2.6 °C/year

Table 3. Cooling rates in the partly molten interior region of Kīlauea Iki lava lake.

[Numbers in parentheses are depths (in feet) of the individual samples in figure 16B used in calculations. Temperatures are based on glass geothermometry (Helz 

and Thornber, 1987)]

*

 Hottest sample recovered in the year the core was obtained.

years Moore and Evans (1967) estimated for the prehistoric 

Makaopuhi body to reach the solidus throughout.

The effect of high rainfall at Kīlauea Iki is clearly visible 

in the upper parts of the temperature profiles. The subsolidus 

interval of the curve for 1975, a year of unusually high 

rainfall, is flatter than most, while that of the 1988 curve, 

monitored during a drought, is steeper. Rainfall did not affect 

the shapes of curves in ‘Alae, but the modeling showed that it 

did hasten cooling of that body. Similar modeling for Kīlauea 

Iki predicted a rate of growth of the upper crust that was 

much faster than observed (Wright and others, 1976; Peck 

and others, 1977), with the drillable crust being only 43 m 

thick in 1975 instead of the predicted 58m. Possible factors 

include: (1) condensation of steam and recirculation of water 

in the thicker upper crust of Kīlauea Iki (Wright and others, 

1976; Peck and others, 1977) and (2) extensive transfer of 

melt and heat upward in Kīlauea Iki (Helz and others, 1989). 

Whatever the balance of these effects, the rate of growth of 

the upper crust in Kīlauea Iki was nearly constant between 

1967 and 1979, at 2.3–2.4 m per year (Helz, 1980).

Thermal data for Kīlauea Iki show that cores collected 

at different times and depths have very different cooling 

histories. For example, partly molten core recovered from 

60–90 m deep cooled at 2–3 °C per year, over the period 

from 1959 to 1979–88, for all time intervals where data are 

available (table 3). By contrast, the hottest samples recovered 

in 1967 cooled from 1,190 °C to 1,090 °C over 8 years, at an 

average rate of 12.5 °C per year. The hottest parts of cores 

from 1975 and 1976 cooled at rates between these values. 

Profiles in the upper crust flatten at less than ~1,100 °C (fig. 

16B); below that temperature, the solidus is reached in one 

year’s time, for a cooling rate of 100+ °C per year. These 

variations in cooling rate have significantly affected the 

crystallization and reequilibration of minerals in Kīlauea Iki, 

especially olivine.

Crystallization and Reequilibration of Olivine in 

Kīlauea Lava Lakes

The classic study of Moore and Evans (1967) on the 

prehistoric Makaopuhi lava lake focused on the observed 

compositions and inferred reequilibration of its olivine 

phenocrysts. For Kīlauea Iki, Helz (1987b) described the 

effects of reequilibration of olivine on interstitial melt 

compositions and groundmass assemblages, while Scowen 

and others (1991) documented how chromite inclusions in 

olivine phenocrysts reequilibrated as olivine became more 

Fe-rich. This section presents olivine data for comparison 

with the results of Moore and Evans (1967). Because Kīlauea 

Iki was sampled as it crystallized, and glass composition 

varies with temperature (Helz and Thornber, 1987), it is 

possible to determine how olivine has crystallized and 

reequilibrated as a function of cooling rate and local bulk 

composition.

Kīlauea Iki contains abundant olivine crystals of all 

sizes. The phenocrysts were present in the eruption scoria 

(Helz, 1987a), as were many smaller olivines, but some 

smaller crystals (<1 mm in length), plus overgrowths on 

inherited phenocrysts, have grown in situ in the lava lake. 

Table 4 summarizes data on the compositions of smaller 

olivines and rims on phenocrysts (hereafter refered to as 

“in situ” olivines) for 36 samples chosen to show the range 

of crystallization and reequilibration behavior in Kīlauea 

Iki. The average composition of in situ olivine decreases 

linearly as temperature decreases from 1,150 to 1,080 °C 

(fig. 17) and is alwaysmore Fe-rich than the initial average 

composition (Fo86–87). At temperatures below the incoming of 

the second pyroxene (pigeonite in the less magnesian samples, 

orthopyroxene in the most olivine-rich core; see table 4), the 

average forsterite (Fo) content of in situ olivine changes more 

slowly. Also, the individual curves fan out, moving toward 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 261

Sample no. Average

(mol % Fo)

 Range 

(mol % Fo)

Temperature

(°C)

Depth 

(m)

Bulk MgO

 (wt %)

Any Low-Ca pyroxene? Comments

A. 1967 samples and subsolidus equivalents

KI67-3-75.0 61.0 69.5 to 52.2 995 22.86 9.47 Pigeonite

KI67-3-77.3 61.8 72.1 to 54.5 1,026 23.56 7.90 Pigeonite

KI67-3-80.7 64.4 72.5 to 56.0 1,057 24.60 7.73 Pigeonite

KI67-3-83.8 64.1 69.9 to 61.1 1,076 25.54 7.54 Pigeonite

KI67-3-83.8R 65.8 71.6 to 61.2 1,076 25.54 7.54 Pigeonite

KI67-3-84.5 69.0 74.5 to 63.4 1,084 25.76 7.69 Pigeonite

KI75-1-85.5 60.0 64.4 to 52.2 Subsolidus 26.06 8.33 Pigeonite

KI81-1-86.0 56.8 59.7 to 53.4 Subsolidus 26.21 8.74 Pigeonite

B. 1975 samples and subsolidus equivalents

KI75-1-130± 63.9 71.2 to 59.0 985 39.62 9.73 Pigeonite

KI75-1-133.3 65.4 71.7 to 58.8 1,035 40.63 10.75 Pigeonite

KI75-1-134.4 65.4 73.4 to 61.6 1,050 40.96 10.90 Pigeonite

KI75-1-138± 67.4 75.0 to 63.3 1,066 42.1 --- Pigeonite

KI75-1-139.3 70.1 77.1 to 67.1 1,084 42.46 11.64 Pigeonite

KI75-1-143.8 75.2 77.8 to 74.5 1,115 43.83 12.15 None

KI75-1-143.8R 76.1 79.1 to 74.5 1,115 43.83 12.15 None

KI75-1-144.9 77.2 77.6 to 76.5 1,117 44.16 --- None

KI75-1-144.9R 77.2 77.3 to 77.0 1,117 44.16 --- none

KI81-1-145.1 64.8 69.1 to 63.0 Subsolidus 44.23 --- Pigeonite

KI88-2-138.9 58.6 63.8 to 55.2 Subsolidus 42.34 --- Pigeonite

C. 1979 samples and subsolidus equivalents

KI79-3-157.7 67.1 69.6 to 61.4 985 48.07 --- Pigeonite

KI79-3-157.7R 66.5 72.8 to 61.2 985 48.07 --- Pigeonite

KI79-3-163.7 67.2 71.1 to 64.4 1,057 49.90 --- Pigeonite

KI79-3-163.7R 69.1 73.4 to 66.1 1,057 49.90 --- Pigeonite

KI79-3-166.1 74.1 75.1 to 72.7 1,071 50.63 27.41 Pigeonite? Vorb

KI79-3-166.1R 75.0 76.5 to 74.1 1,071 50.63 27.41 Pigeonite? Vorb

KI79-3-172.9 75.7 79.7 to 74.1 1,107 52.70 18.71 None

KI79-3-172.9R 76.9 79.5 to 74.4 1,107 52.70 18.71 None

KI79-1-187.4 79.8 80.9 to 79.1 1,130 57.12 19.3 None

KI79-1-203.6 80.9 81.3 to 80.5 1,141 62.06 22.68 None Phenocrysts incl.

KI88-2-147.4 65.8 69.2 to 63.1 Subsolidus 44.93 14.8 Pigeonite

KI88-2-160.5 68.6 74.2 to 65.0 Subsolidus 48.92 17.7 Pigeonite

Table 4.  Average composition and range for olivine in samples from Kīlauea Iki lava lake, with glass quenching temperatures (Helz and Thornber,

1987), depth and bulk MgO content of the sample, and identity of coexisting low-Ca pyroxene, if present.

[Sample numbers give year and drill core number (KI67-3), followed by the sample depth in feet (75.0). All samples contain augite and plagioclase; R, replicate 

analysis; phenocryst core compositions excluded from data except as noted (incl.); mol % Fo, molar percent forsterite; wt %, weight percent; ---, no data; vorb, 

vertical olivine-rich body; Opx, orthopyroxene]
262 Characteristics of Hawaiian Volcanoes

Sample no. Average

(mol % Fo)

 Range 

(mol % Fo)

Temperature

(°C)

Depth 

(m)

Bulk MgO

 (wt %)

Any Low-Ca pyroxene? Comments

D. 1981 and 1988 samples and subsolidus equivalents

KI81-1-181.5 74.7 78.2 to 71.4 1,036 55.32 17.84 Opx

KI81-1-181.5R 73.8 77.2 to 73.2 1,036 55.32 17.84 Opx

KI81-1-181.5 78.5 80.2 to 74.7 1,036 55.32 --- None Dunite inclusion

KI81-1-181.5R 79.2 81.2 to 74.3 1,036 55.32 --- None Dunite inclusion

KI81-1-186.7 75.8 76.8 to 74.6 1,085 56.91 --- Opx

KI81-1-205.4 79.9 80.2 to 79.8 1,134 62.61 --- None Phenocrysts incl.

KI81-1-205.4R 80.1 81.0 to 79.3 1,134 62.61 --- None Phenocrysts incl.

KI81-1-219.8 80.0 80.0 to 79.8 1,140 67.00 23.58 None Phenocrysts incl.

KI81-1-219.8R 80.4 80.5 to 80.2 1,140 67.00 23.58 None Phenocrysts incl.

KI81-1-269.9 77.1 78.0 to 76.8 1,128 82.26 18.79 None

KI81-1-306.7 75.2 76.5 to 74.1 1,087 93.48 17.47 Opx

KI81-1-306.7R 74.9 76.5 to 74.2 1,087 93.48 17.47 Opx

KI88-2-205.2 74.2 76.9 to 72.3 Subsolidus 62.54 19.8 Opx

KI88-2-266.6 77.4 78.5 to 76.8 1,105 81.26 20.9 Opx

KI88-2-300 72.4 74.7 to 71.8 1,081 91.4 ---- Opx

KI88-2-310 73.8 75.3 to 73.2 1,066 94.5 22.5 Opx Foundered crust

KI88-2-336.3 61.4 62.8 to 60.7 ~970 102.5 9.49 Pigeonite Below foundered crust

Table 4.  Average composition and range for olivine in samples from Kīlauea Iki lava lake, with glass quenching temperatures (Helz and Thornber,

1987), depth and bulk MgO content of the sample, and identity of coexisting low-Ca pyroxene, if present.—Continued

[Sample numbers give year and drill core number (KI67-3), followed by the sample depth in feet (75.0). All samples contain augite and plagioclase; R, replicate 

analysis; phenocryst core compositions excluded from data except as noted (incl.); mol % Fo, molar percent forsterite; wt %, weight percent; ---, no data; vorb, 

vertical olivine-rich body; Opx, orthopyroxene]

Figure 17. Graph showing average forsterite content (Fo) 

of in situ olivine versus temperature for core samples from 

Kīlauea Iki lava lake (data in table 4). The red star shows

the average initial composition of olivine in the 1959 eruption 

samples. Individual points show the average in situ olivine 

composition for partly molten samples (blue=1967 core, 

green=1975 core, orange=1979 core, red=1981 and 1988 

cores). Subsolidus samples (from 1975, 1981, and 1988 

cores, recovered at T~100 °C) are shown in gray and plotted

at arbitrary temperatures. Samples plotted at 900 °C are 

6–9 years postsolidus; those at 800 °C are 14–21 years 

postsolidus. The color of the dashed lines shows which molten 

and subsolidus samples most resemble each other in depth 

and bulk MgO content. Further discussion in text. 50

55

60

65

70

75

80

85

90

750 800 850 900 950 1,000 1,050 1,100 1,150 1,200 1,250

Fo content, in mol percent

Temperature, in degrees Celsius

Average initial olivine composition

Low-Ca pyroxene in

below 1,090 °C
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 263

Figure 18. Graphs showing olivine compositions versus depth in Kīlauea Iki lava lake drill cores. A, Compositional ranges for all olivine in 

partly molten 1967 core, plus near- and subsolidus samples with similar bulk MgO content from 1975, 1981, and 1988, plotted against depth. 

The average and extreme in situ olivine compositions are shown by the red dots, highlighted by orange field. Gray diamonds are compositions

of olivine phenocryst cores; green crosses show composition of megacrysts (green lines show ranges). The vertical pink band shows the original 

average olivine composition (Fo86–87) for all Kīlauea Iki drill core. B, As in part A, but for subsolidus samples. The range of compositions in the 

dunite in sample KI81-1-181.5 is included here (same symbol as megacrysts), as it contains no glass within its boundaries. 

120

100

80

60

40

20

0

50 60 70 80 90

Depth below surface, in meters

Olivine composition, in mol percent Fo

Average initial olivine composition

Field of in situ olivine

Olivine core in

foundered crust

Subsolidus samples

120

100

80

60

40

20

0

50 60 70 80 90

Depth below surface, in meters

Olivine composition, in mol percent Fo

Field of in situ olivine

Average initial olivine composition

A B

Dunite

different subsolidus olivine compositions, depending on local 

bulk composition. That local bulk compositional control 

is important is shown by the location of the isolated red 

point in figure 17, which corresponds to the deepest sample 

analyzed (KI88-2-336.3, with MgO=9.49 weight percent), 

with an average in situ olivine composition like that in the 

low-MgO 1967 core. Data for subsolidus samples (plotted 

arbitrarily at 900 °C or 800 °C), are linked in figure 17 with 

the curves best matching their depth and bulk MgO content 

(table 4). The results suggest that the composition of in situ 

olivine continued to change slowly below the solidus (T=970–

1,000°C; Helz, 1987a). 

Figure 18 shows how the compositions of olivine 

phenocryst cores, in situ olivine, and olivine megacrysts 

vary with depth in Kīlauea Iki lava lake; the partially molten 

samples from table 4 are shown in figure 18A and the smaller 

group of subsolidus samples in figure 18B. This projection 

allows comparison with the cooling history of the lava lake 

(fig. 16B), showing clearly that the range of compositions 

for in situ olivine (buff field) within partly molten samples 

reflects the average cooling rate. The relatively shallow, 

more rapidly cooled 1967 and 1975 samples show a wide 

range of in situ olivine composition. By contrast, the 1981–

88 partly molten samples, which have cooled at 2–3 °C per 

year over 22–29 years (see table 3; see also clustered cooling 

curves in fig. 16B), have narrow ranges of in situ forsterite 

content. For these hottest, most slowly cooled samples, even 

phenocryst core compositions fall within the range observed 

for microphenocrystic and groundmass olivine (table 4), so 

they are not shown separately in figure 18A. 

Figure 18 also shows that few olivine crystals retain 

the Fo86–87 composition observed in the eruption samples. In 

particular, olivine phenocrysts in the olivine-rich zone present between 60 and 90 m in the lake have reequilibrated from 

Fo86–87 to Fo74–82, with compositions confined to 79.8–81.3 

mol percent Fo in the very hottest samples. Because of this 

uniformity, no zoning is visible in back-scattered electron 

(BSE) imagery of these samples (fig. 19A). Data from 

Scowen and others (1991) and Vinet and Higgins (2011) 

show similar, restricted ranges of olivine compositions for 

samples from these depths. Much of this reequilibration took 

place before 1975, as the olivine in the vorbs, interpreted as 

derived from the cumulate zone, was already Fo78–80 (Helz 

and others, 1989) in the 1975 core. 
264 Characteristics of Hawaiian Volcanoes

Average forsterite content of in situ olivine drops and 

its range tends to expand as core recovery temperature 

decreases below ~1,100 °C. Phenocryst core compositions 

begin to lag the compositional shift of the in situ olivine 

below 1,130 °C, as can be seen most clearly in the deepest 

samples (below 90m) in figure 18A. This may reflect 

decreasing diffusion rates for Mg and Fe in olivine at lower 

temperatures (Chakraborty, 2010), as well as somewhat 

faster cooling rates below1,100 °C. 

Comparison of figures 18A and B show that in situ 

olivine compositions continue to shift to more Fe-rich 

compositions, and also that the range of compositions (buff 

field) becomes narrower. Phenocryst cores tend to lag, so 

that some samples have developed a second generation of 

olivine compositional zoning after passing through a stage of 

near-uniformity (compare KI79-1-203.6 and KI81-1-205.4 to 

KI88-2-205.2 in table 4). 

This second-generation zoning can be seen clearly in 

the dunite inclusion in sample KI81-1-181.5 (shown in figs. 

25.8a and b in Helz, 1987a). Olivine in the dunite is more 

magnesian, on average, than the in situ olivine (table 4) but 

less magnesian, than phenocryst cores in KI81-1-181.5. The 

dunite is now melt-free, so is plotted in figure 18B, but the 

presence of interstitial augite+plagioclase suggests that grain 

boundaries were formerly coated with melt; the melt films 

and the fine grain size of the olivine presumably facilitated 

reequilibration of the dunite, so that it reequilibrated more 

extensively than the phenocrysts. BSE imagery (fig. 19B) 

shows that the interior of the dunite is very uniform in 

composition, with the present zoning peripheral to the entire 

inclusion. 

It is clear from the data presented here that, in order 

for original (1959) Fo86–87 compositions to be preserved, the 

sample must be (1) shallow, hence rapidly cooled, and (or) 

(2) have unusually large olivine crystals. In deep (60–100m) 

samples, few crystals retain 1959 compositions. Two that 

do (fig.18A) are the megacryst in sample KI81-1-238 and 

the coarsest (3×6 mm) phenocryst in KI88-2-310, which is a 

sample of foundered crust (Helz, 1993). The preservation of 

Fo86–87 in the megacryst probably reflects its size, while the 

survival of Fo86 in the foundered crust sample may be due to 

initial rapid cooling of the crust in the upper part of the lake 

prior to its sinking to its present depth of >95 m, as well as 

the sample’s relatively coarse grain size. 

Figure 20 shows compositional traverses and 

photomicrographs for three megacrysts from Kīlauea Iki. 

KI76-2-20.5 (not shown in figure 18B) is the shallowest 

(6.2m) and most rapidly cooled, with a broad core of Fo86–87; 

its very narrow rims track changes in melt composition to 

the point where the olivine was armored by groundmass 

phases. The megacryst in KI81-1-238 (also shown in fig. 

18A) has shoulder regions that retain Fo86-87, but the inner 

core is Fo85, possibly because of reequilibration with large 

melt inclusions. It has wide rims, zoned out to Fo81–82, in 

contact with the interstitial melt, similar to in situ olivine 

compositions observed in other very hot samples (fig. 18A; 

A

B

500 μm

500 μm

Figure 19. Back-scattered electron (BSE) images of samples from 

Kīlauea Iki lava lake cores. A, BSE image of olivine and groundmass 

in sample KI79-1-203.4, showing that olivine is virtually unzoned at 

this level in Kīlauea Iki lava lake. The total compositional range in

this sample is 80.5–81.3 percent Fo, in crystals of all sizes. Other 

phases present include chromite (bright inclusions in olivine), melt 

(brighter than olivine), augite (small grains, slightly lighter than olivine), 

and plagioclase (dark laths). B, BSE image of the dunite inclusion in 

sample KI81-1-181.5, showing very slight zoning around the periphery, 

in contact with groundmass, and no discernible zoning within the 

dunite. The dunite also contains interstitial chrome spinel, augite, and 

plagioclase, but no melt. 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 265

70

75

80

85

90

0 2.5 5 7.5 10 12.5 15 17.5

Fo content, in mol percent

Position along traverse, in millimeters

KI88-1-190.4

KI76-2-20.5 KI81-1-238

KI76-2-20.5

KI81-1-238

KI88-1-190.4

Figure 20.  Compositional profiles showing forsterite (Fo) content along traverses in three olivine megacrysts from Kīlauea Iki lava lake. Traverses

shown by dashed lines in accompanying photomicrographs. The shallowest megacryst KI76-2-20.5 (red) was cooled fastest; it reached the solidus 

(~1,000 °C) in 1961, cooled to ~100 °C by 1964, and was recovered in 1976. Sample KI81-1-238 (blue) was partly molten when recovered from the hot

zone (T~1,130–1,135 °C) in 1981. Sample KI88-1-190.4 (green) reached the solidus in 1981, cooled to ~100 °C by 1986, and was recovered in 1988.

table 4). The third megacryst (KI88-1-190.4, also shown 

in figure 18B) is similar in composition to the phenocrysts 

in KI81-1-181.5 (table 4), so it is slightly more Mg-rich 

than the fine-grained dunite in that same sample (fig. 18B), 

perhaps because it is a single crystal rather than an aggregate. 

All three, however, retain compositions more magnesian than 

the in situ olivine in the same samples. By contrast, a fourth 

megacryst (in sample KI88-1-255), quenched from ~1,050 

°C, has reequilibrated to the same composition as the in situ 

olivine, as shown in figure 18A. Evidently, even the largest 

olivine crystals reequilibrate to groundmass compositions if 

they are in contact with melt for 29 years. 

The pattern of reequilibration of olivine in Kīlauea Iki 

compares closely with that described by Moore and Evans 

(1967) in the prehistoric Makaopuhi lava lake, although 

olivine compositions are more magnesian in Kīlauea Iki. 

The samples with the most uniform olivine in the prehistoric 

Makaopuhi lava lake were found within the olivine-rich 

zone, where phenocrysts reequilibrated to Fo64.3–68.3 (Moore 

and Evans, 1967), with ranges of 1.5–3.3 mol percent Fo, 

similar to those from the hottest zone within Kīlauea Iki 

(table 3). Samples higher in the section showed progressively 

wider ranges of olivine compositions, as observed in Kīlauea 

Iki (fig. 18A,B). Similar extensive reequilibration of olivine 

from chilled margin to the core of the Uēkahuna laccolith 

was reported by Wilkinson and Hensel (1988).

An important difference between the two lava lakes is 

that groundmass olivine, ubiquitous in Kīlauea Iki, was found 

only in the uppermost 6 m of the Makaopuhi lake, being 

replaced by low-Ca pyroxene at greater depths; however, the 

occurrence of orthopyroxene oikocrysts is very similar in the 

two bodies in both location and habit. In Makaopuhi (Evans 

and Moore, 1968), poikilitic orthopyroxene is found in the 

olivine cumulate zone (40–67 m below the surface), as it is in 

Kīlauea Iki (table 4; also Helz, 1987b). The oikocrysts enclose 

plagioclase and Fe-Ti oxides without displacing them but 

appear to have completely consumed both olivine and augite 

as they grew (fig. 21; also Evans and Moore, 1968). In both 

bodies, the oikocrysts are better developed below the thermal 

maximum, probably because the cooling rate in the lower crust 

is lower than in the upper crust (Moore and Evans, 1967). 

Summary and Future Directions

Lava lake studies, especially on closed-system lava lakes, 

have produced a wealth of information about how basaltic 

systems crystallize and fractionate. The list of fractionation 

processes observed or inferred is large (Wright and others, 

1976; Wright and Okamura, 1977; Helz, 1987b, 2009), as is the 

range of whole-rock and melt compositions analyzed. These 

results have been widely applied to (inter alia) the structure of 

mid-ocean ridge magma chambers (Barth and others, 1994) and 
266 Characteristics of Hawaiian Volcanoes

the behavior of melts in marginal zones in magma chambers 

(Tait and Jaupart, 1996). 

Results presented here that are relevant to the interpretation 

of basaltic systems include observations that (1) the actual 

efficiency of segregation/extraction of differentiated liquids 

is much lower than might be predicted from theoretical 

fractionation of bulk compositions; (2) the position of the last 

interstitial liquids (at Kīlauea Iki they are 80–100 m below 

the surface; see fig. 16B) is much deeper than the position of 

the most differentiated part of the lake (20–40 m; Helz, 2009), 

which solidified before 1975; (3) reequilibration of original 

olivine compositions from Fo86–87 to ≤Fo80–81 is extensive 

after 15 years and essentially complete after 30 years of slow 

cooling; and (4) a shift in cooling rates (from 2–3 °C/year above 

1,100 °C to 100°C/year below 1,100 °C) produces a distinct 

second generation of normal zoning in olivines from the most 

slowly cooled parts of the lava lake. 

The use of samples from Kīlauea lava lakes for studies 

of trace elements and isotopes is just beginning. Trace element 

compositions of bulk samples usually mirror processes inferred 

from major elements (Wright and Peck, 1978; Wright and Helz, 

1996; Helz, 2012) in Kīlauea lavas. Study of the platinumgroup elements plus rhenium (Re) in Kīlauea Iki (Pitcher and 

others, 2009) documented their behavior in the low-pressure, 

high-temperature igneous environment, unmodified by subsequent (including hydrothermal) processes. One result is that Re 

is progressively depleted in later drill core relative to samples 

quenched earlier, apparently because of loss of Re to the volatile 

phase. Recent work by Teng and others (2008, 2011) and Sio 

and others (2011), on both bulk samples and individual crystals of olivine, has found significant fractionation of Fe and 

Mg isotopes in olivines from Kīlauea Iki, apparently produced 

during diffusive reequilibration, whereas Fabbrizio and others 

(2010) have documented the preservation of phosphorus zoning 

in otherwise reequilibrated, unzoned olivines from Kīlauea Iki 

(KI81-1) samples. With expanded use of microbeam analysis 

for trace elements and isotopes, it should be possible to exploit 

samples from the various historical lava lakes, which include 

cores containing melts quenched from temperatures ranging 

from 980 to 1,140 °C, for many studies in the future. 

Other workers have taken advantage of the known cooling history and extensive sample collections from Kīlauea Iki 

and other lava lakes to investigate the evolution of mineral 

textures. Cashman (1993) used plagioclase data from the 

prehistoric Makaopuhi lava lake (Moore and Evans, 1967) 

to elucidate cooling rates in bodies of crystallizing basalt. 

Studies of Kīlauea Iki samples include crystal size distribution 

analysis of olivine (Vinet and Higgins, 2011) and variation in 

the intercrystalline angle at plagioclase-pyroxene-plagioclase 

contacts (Holness and others, 2012a,b), also proposed as a 

means of estimating cooling rates. 

Studies of Submarine Hawaiian Lavas

The collection of lava samples from the submarine slopes 

of Hawai‘i began in the early 1960s (Moore, 1965; Moore and 

Fiske, 1969). Such submarine erupted samples, quenched under 

high water pressure, preserve melt compositions as glass and are 

natural analogues to hand-dipped molten samples collected and 

quenched during eruptions at Kīlauea and Mauna Loa. These 

early samples from Kīlauea, Mauna Loa, and Mauna Kea were 

incorporated into many subsequent studies on the petrology 

and geology of Hawai‘i. An important attribute of submarine 

erupted and quenched basaltic glasses is that volatile components that would be lost to the atmosphere during subaerial 

eruption are largely trapped in solution in the glass and can be 

analyzed to determine preeruptive volatile contents. One of the 

most striking discoveries from study of submarine lavas was 

that Lō‘ihi Seamount is an active Hawaiian volcano, likely to 

become the next Hawaiian island (Moore and others, 1982). 

Equally striking was the discovery that Lō‘ihi lavas included 

alkalic basalts, as well as the expected tholeiitic shield lavas. 

Preshield Stage Lavas

Rocks recovered from Lō‘ihi Seamount in 1981 

included a range of alkalic lavas, such as basanite, alkalic 

basalt, and hawaiite (Moore and others, 1982), some of 

which contained mantle xenoliths (Clague, 1988). Thicker 

palagonite alteration of the alkalic glasses suggested they 

were generally older than the tholeiitic lavas and led to the 

conclusion that Lō‘ihi Seamount, in particular, and Hawaiian volcanoes in general, began their growth with an early or 

preshield alkalic stage (Clague, 1987a). Lō‘ihi is well studied (see review by Garcia and others, 2006, for references), 

500 μm

Figure 21. Photomicrograph showing orthopyroxene oikocryst (light 

gray) extending between several small olivine phenocrysts in sample 

KI81-1-306.7 (T=1,087 °C) from Kīlauea Iki lava lake. Nicols crossed.

Note the near-absence of both olivine and augite inclusions within the 

area of the oikocryst. 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 267

and detailed petrologic studies indicate that the composition of Lō‘ihi’s rocks represents one of the several mantle 

components currently thought to contribute to generation 

of Hawaiian magmas (see, for example, Dixon and Clague, 

2001; Huang and others, 2009). The discovery of early 

alkalic magmas constrains melting models, such that growth 

of Hawaiian volcanoes begins and ends with smaller-degree 

alkalic melts (Clague, 1987b; Clague and Dalrymple, 1987) 

separated from more enriched mantle source rocks at greater 

depth during the pre- and postshield alkalic stages, compared 

with the larger-degree tholeiitic melts that form the voluminous shield stage, separated at shallower depths from less 

enriched mantle sources (for example, DePaolo and Stolper, 

1996; Xu and others, 2007). 

The high 3He/4He isotopic ratios and other rare-gas 

compositions of Lō‘ihi (and some Kīlauea) lavas (for example, 

Kurz and others, 1983; Honda and others, 1991, 1993; Trieloff 

and others, 2000, 2001) indicate an ancient, undegassed mantle 

source and have been a primary argument for the deep-mantle 

source of the plume, although this view is not uncontested (see, 

for example, Anderson, 1998). Neon isotopes have been proposed to have a solar (Honda and others, 1991, 1993, Ballentine 

and others, 2001) or a meteoritic (Trieloff and others, 2000, 

2001) signature, indicating an ancient source undegassed during 

Earth’s history. In contrast, analyses of Lō‘ihi lavas for CO2, 

H2O, Cl, and S (Dixon and Clague, 2001) show that the Lō‘ihi 

mantle component is relatively dry and therefore derived from a 

recycled mantle component that preferentially lost H2O during 

subduction. This conclusion seems to contradict that based on 

rare gas components, but it may indicate that the rare gases are 

derived mainly from one mantle component, whereas water is 

mainly derived from another. 

The Early Growth of Kīlauea 

Lō‘ihi Seamount was the only known example of a 

preshield-stage volcano until 1998 to 2002, when a series of 

dive programs, using the Japan Agency for Marine Earth Science and Technology (JAMSTEC) remotely operated vehicle 

Kaiko and manned submersible Shinkai 6500, explored the 

deep rift zones and flanks of a number of Hawaiian volcanoes. 

These dives discovered extensive preshield-stage alkalic products from early Kīlauea (Lipman and others, 2002; Sisson and 

others, 2002; Sisson, 2003; Coombs and others, 2004, 2006; 

Calvert and Lanphere, 2006; Kimura and others, 2006; Hanyu 

and others, 2010), as well as small exposures of alkalic material from early Hualālai (Hammer and others, 2006).

Because of the Hawaiian Volcano Observatory’slocation on 

Kīlauea and emphasis on shield-stage Kīlaueastudies,the longterm development of Kīlauea is of particular relevance. Early 

Kīlauea eruptive products are exposed in the submarine region 

due south of the volcano’s summit. This region consists of a 

south-facing slope that descends steeply from the shoreline 

to a broad mid-slope “Hilina” bench at ~2,500–3,000 m 

depth, fronted by another south-facing scarp that descends 

steeply to the seafloor at ~5,000-m depth. The submarine 

Hilina bench is elongate northeast-southwest, parallel to the 

shoreline, with maximum dimensions of 15 by 55 km. It is 

bounded on the southwest by Papa‘u Seamount and tapers to 

the northeast, where it merges with the Puna Ridge. Because 

of its flat top and orientation parallel to the shoreline, 

the Hilina bench was initially interpreted as a prodigious 

rotational landslide block that dropped ~10 percent of 

Kīlauea’s subaerial tholeiitic shield to great water depths 

along normal faults similar to those of the subaerial Hilina 

Pali (Lipman and others, 1985; Moore and others, 1989; 

Smith and others, 1999). Initial JAMSTEC dive observations 

and recovered samples, however, revealed that the frontal 

scarp exposes bedded sandstones and debris-flow breccias 

(Lipman and others, 2002), consistent with seismic reflection 

profiles (Hills and others, 2002) that show the Hilina bench 

to be a fold and thrust wedge of marine volcaniclastic rocks 

being extruded from beneath growing Kīlauea as a result 

of edifice spreading. The upper slope, above the bench and 

toward the shoreline, is widely mantled by unconsolidated 

to weakly consolidated glassy sands derived by shattering 

of shoreline-crossing lava flows, but narrow rock ribs 

expose underlying in-place pillow lavas incised in landslide 

headscarps. 

The oldest preshield Kīlauea products are exposed along 

the 3,000–5,000-m deep frontal scarp of the Hilina bench, 

and Kīlauea’s submarine products become progressively 

less alkaline and younger, overall, both at shallower water 

depths and northeastward toward the Puna Ridge. Bedded 

sedimentary rocks of the frontal scarp are mainly volcanicglass sandstones, with glass grains dominated by degassed, 

relatively Si-rich, and Na- and Ti-poor tholeiitic compositions, 

similar to modern products of Mauna Loa, but accompanied by 

subordinate populations of alkalic glasses, most with elevated 

S, CO2, and H2O concentrations indicative of submarine 

eruption (Sisson and others, 2002, 2009; Coombs and others, 

2006). Similarly, clasts in bedded debris-flow breccias of the 

frontal scarp are Mauna Loa-type tholeiites accompanied by 

alkalic clasts, rarely with preserved glass that also has elevated 

volatile concentrations. Alkalic glasses encompass basanites, 

nephelinites, hawaiites, alkali basalts, tephriphonolites, 

phonotephrites, benmoreites, and mugearites (Sisson and others, 

2009). Alkalic clasts encompass alkali basalts, hawaiites,and 

basanites; a probable sill or thick lava flow of phlogopite 

nephelinite was also discovered. Judging from the wide range 

of eruption depths interpreted from dissolved volatiles, Kīlauea 

began as diverse alkalic vents spread across the south flankof 

the extant and active Mauna Loa shield. The sedimentary 

rocks of the frontal scarp are composed of fragmentation 

products shed mainly from shattering of shoreline-crossing 

Mauna Loa lavas and also from shattering, landsliding, and 

explosive eruptions of shallow-to-deep marine early Kīlauea 

vents (Coombs and others, 2006). Minor amounts of low-Si 

tholeiitic glasses with elevated S concentrations accompany the 

alkalic and degassed Mauna Loa products, possibly indicating 

that transitionally tholeiitic magmas appeared early in Kīlauea’s 
268 Characteristics of Hawaiian Volcanoes

history. Alkalic samples from the frontal scarp have 40Ar/39Ar 

ages spanning from 280±20 ka to 212±38 ka, including higher 

precision results on phlogopite from nephelinite of 234±9 and 

236±10 ka (Calvert and Lanphere, 2006). These results support 

an inception age of ~300 ka for Kīlauea.

A narrow rock rib above the midslope bench, directly 

toward Kīlauea’s summit, consists of in-place weakly alkalic 

and transitional basaltic pillow lavas (Kimura and others, 2006). 40Ar/39Ar ages of in-place, weakly alkalic pillow lavas from the 

upper slope span from 195±48 ka to 138±30 ka (Calvert and 

Lanphere, 2006), and two transitional basalt clasts shed from the 

upper slope onto the midslope bench as landslide debris have 

ages as young as 67±29 ka and 65±28 ka (Hanyu and others, 

2010). To the northeast, where the midslope bench gives way 

to the Puna Ridge, lava samples are tholeiitic, but their glasses 

are notably lower in Si and higher in Ti and Na than is typical of 

the subaerial Kīlauea shield. 40Ar/39Ar ages measured on these 

low-Si tholeiites (transitional basalts of Sisson and others, 2002) 

are imprecise because of low K concentrations, but are 228±114 

and 138±115 ka (Calvert and Lanphere, 2006). The location of 

these low-Si tholeiites toward the Puna Ridge, and the absence 

of alkalic materials from that region, suggest that they record 

an intermediate stage in Kīlauea’s development commencing 

at roughly 150–100 ka. At that time, volcanic output increased 

markedly, initiating growth of the East Rift Zone and the Puna 

Ridge and ending the period of alkalic magmatism, during 

which the diversity and degree of undersaturation of alkalic 

products diminished as the volcano matured. 

Subsequently, as volcanic output increased markedly, 

magma compositions came to be dominated by low-Si tholeiites. Cessation of this low-Si tholeiitic intermediate growth 

period was probably gradual, giving way to tholeiites typical 

of Kīlauea’s subaerial shield, as is shown by the presence 

of low-Si tholeiitic lavas among ordinary tholeiites exposed 

in fault scarps of the lower subaerial south flank, estimated 

to be close to 25 ka in age (Chen and others, 1996), and the 

local presence of similar relatively high-Ti and high-K glasses 

(440–640 C.E.; Fiske and others, 2009) in the Kulanaokuaiki 

Tephra. Although the record of Kīlauea’s development is 

complicated by posteruptive sedimentation, concurrent deposition of products from Mauna Loa, deformation resulting from 

island spreading, and incomplete exposures, it is the only 

Hawaiian volcano where products of all stages, from inception 

to mature shield stage, are accessible and have been dated by 

radiometric dating.

Occurrence and Characteristics of Submarine 

Picritic Lavas

The deep submarine rift zones of Hawaiian volcanoes 

consist of common to abundant picritic tholeiites, with less 

common aphyric or sparsely phyric lavas (Moore, 1965; Garcia 

and others, 1995b). Clague and others (1995b) studied Kīlauea 

samples reported in Moore (1965), plus more recently acquired 

samples, and found that all the lavas had fractionated glasses 

with MgO<7 weight percent, while containing abundant olivine 

and less abundant augite, plagioclase, and orthopyroxene. 

Many of the olivine crystals are either more primitive or more 

evolved than crystals in equilibrium with the host melts; on 

average, the data parallel an olivine control line of Fo87.9, 

which is more magnesian than the Fo86–87 observed in Kīlauea’s 

subaerial lavas. Clague and others (1995b; and unpublished 

data) find 13 of 47 sampled submarine sites have rocks with 

whole-rock MgO>15 weight percent, so picritic lavas are 

25–30 percent of the sampled flows and much more abundant 

than picritic lavas on land. The glasses are generally highly 

fractionated and contain between 4.3 and 7 percent MgO, with 

a fairly even distribution from 4.8 to 6.5 percent and a few 

samples with higher and lower values. Another suite of 65 

lava rind glasses recovered along the submarine rift (Johnson 

and others, 2002) includes four with MgO>7 weight percent 

(one at 7.2 and three from the same dredge at 9.2 weight 

percent MgO), but almost all are between 4.2 and 6.7 weight 

percent MgO. Thirty-five other tholeiitic glasses from the Puna 

Ridge and south flank of Kīlauea range from 5.3–7.1 weight 

percent MgO (Coombs and others, 2006). The combination of 

glass, mineral, and whole-rock compositions was interpreted 

by Clague and others (1995b) as having been formed by a 

complex series of processes including spinel+olivine and 

then multiphase (olivine+plagioclase+augite) fractionation, 

degassing, wall-rock stoping and assimilation (Kent and others, 

1999), magma mixing in the crustal reservoir, entrainment of 

olivine xenocrysts from a hot ductile olivine cumulate body 

(Clague and Denlinger, 1994), and disruption of gabbroic 

wall rocks in the rift zone. The rocks may thus be picritic, 

but the host melts are almost all highly fractionated and have 

undergone a complex evolution.

On Mauna Loa, 54 percent of 46 samples from the 

submarine south rift are picritic tholeiites containing 

>15 percent olivine crystals (Garcia and others, 1995b). The 

host lavas contain 5.8 to 8.4 weight percent MgO, while olivine 

compositions vary from Fo81 to Fo91. Moore and Clague (1992) 

report only 7 of 19 whole-rock pillow samples from this rift 

as having >15 weight percent MgO, and the host glasses of 20 

samples range from 4.5 to 8.0 weight percent MgO. Seventeen 

of these 20 samples have glass MgO between 5.5 and 6.9 weight 

percent. Wanless and others (2006a,b) report compositions for 

lavas from radial vents on the submarine west flank of Mauna 

Loa, but only 1 of 9 samples contains >15 weight percent MgO, 

with the remainder having whole-rock MgO contents of 6.0–7.6 

percent. Morgan and others (2007) report MgO contents of 

4.9–7.8 weight percent MgO, with 90 percent between 5.4 and 

6.9 weight percent MgO (median of about 6.3 weight percent 

MgO) in 259 glass grains and pillow rinds from the southwest 

flank of Mauna Loa. Only 5 of 54 whole-rock analyses contain 

>12 weight percent MgO, and none have more than the 15 

weight percent MgO that would be considered picritic. Thus, 

although common to abundant picritic lavas occur along Mauna 

Loa’s submarine rift zone, most have host melts that are highly 

fractionated and have undergone a complex evolution similar to 

those on Kīlauea’s Puna Ridge.
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 269

On Lō‘ihi Seamount, nine samples from the lower part of 

its south rift zone (depth 4,976–5,058 m) are picritic tholeiites, 

with all whole-rock analyses having >20 weight percent MgO 

(Matveenkov and Sorokhtin, 1998). Tholeiitic samples, too 

numerous to count, from the summit, upper south rift zone, and 

east flank are only rarely picritic (see, for example, Moore and 

others, 1982; Frey and Clague, 1983; Hawkins and Melchior, 

1983; Garcia and others, 1993, 1995b, 1998; Clague and others, 

2000, 2003; Dixon and Clague, 2001). 

The submarine part of Hualālai’s northwest rift zone is also 

characterized by abundant picritic lavas (Clague, 1987b; Moore 

and Clague, 1992). Fifteen of 22 whole-rock analyses have 

MgO>15 weight percent, and 11 of those have MgO=20–31.1 

weight percent. The host glasses contain 5.9 to 8.0 weight 

percent MgO (median about 6.8 weight percent), while the bulk 

compositions define an olivine control line of ~Fo89. Fully onethird of the glasses have >7.2 weight percent MgO, compared 

with only 9 percent of glasses from Mauna Loa and none from 

Kīlauea’s Puna Ridge. Tholeiitic samples from the submarine 

west flank of Hualālai (Hammer and others, 2006; Lipman and 

Coombs, 2006) include picritic lavas, but the glasses in rinds 

and fragments range from 6.0 to 8.8 weight percent MgO, with 

a median of about 6.5 weight percent MgO for the picritic pillow lavas from Kaiko dives 218 and 219 (Lipman and Coombs, 

2006), and a median of about 7.15 weight percent MgO for the 

glass grains in rocks from Shinkai dives 690 and 692 (Hammer 

and others, 2006). 

Although picritic lavas are common to abundant along the 

submarine rifts of Kīlauea, Mauna Loa, Hualālai, and Lō‘ihi, 

none represent primary or near-primary melt compositions, all 

indicating, instead, strongly to moderately fractionated melts 

with variable amounts of added olivine, mostly as deformed 

xenocrysts. Picritic lavas from these several volcanoes differ 

in two ways: (1) the ranges of MgO contents of the host melts 

differ, with Kīlauea having the most fractionated host melts and 

Hualālai having the least fractionated ones, and (2) the amount 

of olivine added to these host melts also varies, with Hualālai 

and Lō‘ihi having abundant olivine xenocrysts added to the 

melts and Kīlauea having the least added olivine. 

Glasses and Volatiles in Samples from Kīlauea’s 

Submarine East Rift Zone

 Sand grains recovered in a box core north of the Puna 

Ridge included a spectrum of glasses with MgO contents as 

high as 14.7–15.0 weight percent (Clague and others, 1991, 

1995b; Wagner and others, 1998). These glasses are Kīlauealike and allow reconstruction of primary melt compositions by 

adding in small amounts of olivine, with minimal uncertainty 

due to fractionation, mixing, or assimilation of wall rock. 

The primary melts are estimated to contain 13.4 to 18.4 

weight percent MgO, depending on their FeO contents; the 

volatile contents for a primary melt with 16.4 weight percent 

MgO are estimated to be 1,000 ppm S, 100 ppm Cl, and 

0.37 weight percent H2O (Clague and others, 1995b). The 

melts have calculated liquidus temperatures of 1,283 °C 

to 1,384 °C, with an average of 1,346 °C, and very low 

viscosities (0.3–0.6 Pa·s), so olivine would sink rapidly in 

these melts. They also have densities slightly greater 

(2.66–2.68 g/cm3) than more fractionated melts, so the 

magma reservoir should be stably stratified until the 

uppermost magma degasses; if no crystallization occurred, 

this would increase melt density and could cause turnover in 

the reservoir (Clague and Dixon, 2000). 

Additional high-MgO glass grains with as much as 11 

weight percent MgO were found in Kīlauea beach sand, recovered in the Hilo drill core (Beeson and others, 1996), and also in 

several ash deposits on Kīlauea (as much as 14.5 weight percent 

MgO; Clague and others, 1995a; Helz and others, 2014). These 

latter glasses were erupted subaerially or in shallow water and, 

unlike the deep submarine ones, have lost most of their volatiles. Lastly, extensive sampling along the Puna Ridge (Johnson 

and others, 2002) recovered glasses with as much as 9.2–9.3 

weight percent MgO; these were from the same shallow cone 

as three pillow fragments from picrites (glass MgO=9.2 weight 

percent) that were mentioned above.

Measurements of magmatic H2O and S in Hawaiian 

submarine lavas (some of the earliest such studies were Moore 

and Fabbi, 1971; Killingsley and Meunow, 1975; Meunow 

and others, 1979) yield varying results. Dixon and others 

(1991), using many of the samples from Kīlauea dredged 

by Jim Moore in 1962, found variable H2O and S contents. 

They attributed the variability to mixing of degassed and 

undegassed magmas, followed by equilibration of CO2 in the 

hybrid at the eruption pressure. The degassed end member 

required gas loss at very low pressures, perhaps even subaerial 

conditions. Such recycling of degassed lava back into the 

magma chamber, where it might mix with undegassed magma, 

was observed during the 1959 Kīlauea Iki eruption, which had 

repeated cycles of high fountains and drainback (Richter and 

others, 1970; Eaton and others, 1987). One component of the 

1959 eruption is known to have traveled out Kīlauea’s East 

Rift Zone at least as far as Puna, where it mixed with stored 

rift lava, and then was erupted in the latest 1960 lavas (Wright 

and Helz, 1996). It and earlier magmas could well have traveled farther, to erupt along the submarine east rift. 

Kīlauea and Mauna Loa: Insights into 

Primary Magma Compositions and 

Sources

Because of their isolation from interaction with continental crust, Hawaiian volcanoes, and especially Kīlauea 

and Mauna Loa, provide a direct look at basaltic magmas 

and their mantle sources. Similarly, the relative simplicity 

of each volcanic edifice has allowed a relatively clear look 

at magmatic plumbing systems. This final section will first 

summarize petrologic constraints on primary (or parental) 
270 Characteristics of Hawaiian Volcanoes

magma compositions, the nature of the mantle source(s), 

and the relationship between compositions of Kīlauea and 

Mauna Loa, both shield-stage and preshield. Then we will 

consider the nature of the currently active volcanic plumbing 

system and review the constraints imposed by petrologic data 

on magma storage, especially for the summit reservoirs of 

Kīlauea and Mauna Loa. 

Primary Magmas and Sources for Shield-Stage 

Lavas

One of the earliest papers to grapple with the subject of 

primary melt compositions, coupled with the need for mantle 

sources capable of producing large volumes of basaltic melt 

at closely spaced volcanoes, is Murata’s (1970). Given his 

assumption that primary Hawaiian magma had ~20 weight 

percent MgO, to be in equilibrium with peridotite, and given 

the 37±3 km spacing between the volcanoes, he estimated that 

Mauna Loa would need a source ranging from 60 to 125 km in 

depth. In view of the limited range of composition of olivine 

phenocrysts observed at Kīlauea and Mauna Loa, however, 

he suggested that this primary liquid would rise quickly, 

crystallizing significant olivine only in the summit reservoir. 

Mauna Loa basalts have higher SiO2 contents at a given 

MgO content (Tilley and Scoon, 1961; Wright, 1971), which 

led Murata (1970) to suggest that their source was richer in 

orthopyroxene than that of Kīlauea’s lavas. Work of Clague 

and others (1991, 1995b), discussed above, has documented 

the existence of glasses with MgO contents as high as 15 

weight percent, present as fragments in dredge samples from 

the submarine part of Kīlauea’s East Rift Zone. How this 

material leaks to the surface is not clear, but, even in trace 

amounts, it provides direct evidence that melts in Kīlauea’s 

plumbing may be highly magnesian, though not quite as 

magnesian as Murata (1970) suggested.

Wright (1984) revisited the question of the origin of 

Hawaiian tholeiite, again with a particular focus on Kīlauea 

and Mauna Loa. This approach, which involved modeling 

rare earth elements (REE) as well as major elements, 

established the need for a complex source, in which the 

large melt fractions (as much as 42 percent of the hybrid 

peridotite source) were supplemented by varying amounts of 

nephelinitic liquid plus amphibole, apatite, and ilmenite. The 

nephelinite provided the REE signature for residual garnet, 

a feature of both Kīlauea and Mauna Loa basalts (Hoffman 

and Wright, 1979; Leeman and others, 1980). Melting and 

melt segregation were considered to take place over a range 

of pressures. A dynamic model, with fresh mantle moving 

beneath the volcanoes, reduced the thickness of the melted 

zone needed from Murata’s (1970) estimate of 65 km to 

~27 km (Wright, 1984). In contrast to Murata (1970), Wright 

estimated the parental compositions of Kīlauea and Mauna 

Loa basalts as having 13–14 weight percent MgO, with 

olivine fractionating during rise of magma from the mantle 

to the base of the summit reservoirs. This is consistent 

with data of Helz and others (2014), who have found rare 

subaerial glasses in the Pāhala Ash at Kīlauea with as much 

as 14.5 weight percent MgO. 

More recent models also hypothesize complex mantle 

sources for Hawaiian shield-stage tholeiites. Hauri (1996) 

suggested that oceanic crust is a significant component, especially prominent in the Loa-line volcanoes, from Ko’olau 

onward to Mauna Loa. The model offers no explanation for 

why this component is so much lower in the Kea-line volcanoes (such as Kīlauea), and the isotopic data from Lō‘ihi 

(a Loa-line volcano) require still further complexities. More 

recently, Sobolev and others (2005), although acknowledging that the mantle is dominantly peridotite, suggested that 

the source for Hawaiian tholeiite is secondary pyroxenite 

produced by reaction between mantle peridotite and recycled 

oceanic crust. A new study by Putirka and others (2011), 

however, suggests that the Ni content of olivines in Hawaiian 

picrites requires a peridotitic source. Their proposed parental melts for Kīlauea and Mauna Loa have MgO=18 weight 

percent (versus 20 weight percent assumed by Murata, 1970, 

and 13–14 weight percent assumed by Wright, 1984). 

Commonalities among the various models include (1) 

that parental melts are highly magnesian, although rarely 

seen at the surface, whether the source is strictly peridotitic 

or not; (2) that there is much olivine removal as the magmas 

ascend higher in the volcanic plumbing (resulting in a gravity high associated with the magma conduit; see, for example, Wright and Klein, 2006); and (3) that separate sources 

(or different mixes of the various components) are clearly 

needed for volcanoes along the Loa and Kea lines.

Implications of Isotopic Signatures of Preshield 

Lavas for the Origins of Hawaiian Magmas 

To a first-order approximation, the succession of lava 

types and erupted volumes at Hawaiian volcanoes, from 

early alkalic to voluminous shield-stage tholeiitic to late 

alkalic, can be interpreted as an increase and then a decrease 

in extent of melting as each center passes over some portion of the leading edge, interior, and trailing margin of the 

upwelling source region (Moore and others, 1982; Clague, 

1987a; Clague and Dalrymple, 1987). Isotopic measurements 

show, however, that the types of materials undergoing melting also change with the stages of volcano development. Isotopes of Pb, Sr, Nd, Hf, and He can be interpreted in terms of 

four general end-member sources for Hawaiian magmas (see 

summary in Hanyu and others, 2010). These end members 

are represented most purely by the early and late products 

of the volcanic systems, and, notably, the sources for early 

magmas are not the same as those for the latest, suggesting 

an isotopically zoned and (or) mixed mantle source. 

Inception-stage products are currently well represented by 

abundant samples and isotopic measurements from Lō‘ihi and 

offshore Kīlauea (Kurz and others, 1983; Staudigel and others, 

1984; Stille and others, 1986; Kimura and others, 2006; Hanyu 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 271

and others, 2010). Most early Kīlauea magmas are distinguished 

by their high and uniform values of 206Pb/204Pb, defining what 

has been named the “Kea” or “Hilina” end-member source 

component. Lō‘ihi’s magmas do not have as high 206Pb/204Pb 

but are relatively high in 208Pb/204Pb for Hawai‘i, defining the 

“Lō‘ihi” end-member source component. An important further 

distinction is that Kea end-member samples from early Kīlauea 

have moderate values of 3He/4He (10–15 times the present 

atmospheric ratio [Ra]), whereas 3He/4He in the Lō‘ihi end 

member is characteristically high (to >30 Ra). A “depleted” 

end-member component, similar isotopically to many mid-ocean 

ridge basalts (low 87Sr/86Sr, high 143Nd/144Nd), is sampled most 

purely by low-volume alkalic magmas that erupt well after 

shield-building, as well as on the crest of the Hawaiian Arch 

(Chen and Frey, 1983, 1985; West and others, 1987; Frey and 

others, 2000). The fourth, “Ko‘olau,” end-member component 

is represented by anomalously high-SiO2 basalts or basaltic 

andesites that erupted late in the shield stages of Ko‘olau and 

Lāna‘i and that have high 87Sr/86Sr, but low 206Pb/204Pb and 143Nd/144Nd (Roden and others, 1994).

The progression in isotopic values of Hawaiian magmas 

gives insights into possible configurations of the region 

undergoing melting. Inception-stage magmas are dominated 

either by the Lō‘ihi or Kea components, whereas shield-stage 

tholeiites have isotopic values intermediate between those 

and the depleted or Ko‘olau components. There is also a gross 

geographic division in which tholeiitic shields on the northeast 

side of the Hawaiian Island chain have the Kea component 

as a persistent contributor, while shields on the southwest 

(Loa) side receive contributions from the Lō‘ihi component 

(Tatsumoto, 1978). The origin of this bilateral division is 

unknown, but it is consistent with the relative geographic 

positions of early Kīlauea and Lō‘ihi. An apparent paradox 

is that the inception-stage magmas of Kīlauea and Lō‘ihi are 

chemically heterogeneous, encompassing diverse alkalic and 

transitional compositions, but tend to be isotopically uniform 

within each locality, whereas the shield-stage tholeiites can be 

chemically monotonous but are isotopically quite variable, with 

contributions from the depleted and Ko‘olau components that 

do not appear in the inception-stage products. Explanations 

for this paradox have postulated (1) that the Hawaiian magma 

source is compositionally zoned with the Lō‘ihi and Kea 

components localized to the outer or leading margin, the 

Ko’olau component concentrated in the interior (Takahashi 

and Nakajima, 2002), and the depleted component either also 

in the interior or representing ambient upper mantle melted by 

the Hawaiian source, (2) that the Lō‘ihi and Kea components 

are readily melted materials widely distributed in the Hawaiian 

magmatic source region within a matrix of more refractory 

Ko‘olau and possibly depleted components that are only 

capable of melting in the central region where temperatures 

are high and upwelling is vigorous (Kimura and others, 2006; 

Hanyu and others, 2010), or (3) a counterintuitive view that 

inception-stage components are concentrated in the interior 

of the upwelling source but appear most purely in inceptionstage magmas because those components melt deep and early, 

whereas marginal, more refractory, components only melt with 

greater degrees of upwelling where the source has also been 

dragged “downstream” by motion of the overriding Pacific 

Plate (DePaolo and Stolper, 1996; DePaolo and others, 2001).

An issue with interpretations that would place the Lō‘ihi 

component, in particular, at the outer edge of the Hawaiian 

magmatic source is that this component is typified by high 

3

He/4He. 4He is an alpha-decay product of heavier radioactive 

elements, mainly uranium and thorium, whereas 3He is not, so 

if a region has undergone degassing early in Earth’s history, 

it will have low He/U or He/Th and subsequent radioactive 

ingrowth of 4He will lead to a low 3He/4He value. As an 

example, depleted mid-ocean ridge basalts have a 3He/4He ratio 

of 8±1 Ra (Hilton and Porcelli, 2005). Conversely, regions that 

escaped early degassing should have high 3He/4He values. It 

has been interpreted, therefore, that the high 3He/4He Lō‘ihi 

component originates from portions of the mantle that escaped 

degassing, or were less degassed than average, and that this 

source must reside in the deep mantle (Kurz and others, 1983). 

Dynamic models of mantle upwelling indicate that deeper 

materials end up in the interior of an upwelling zone or plume 

with shallower materials entrained along the margins (Hauri 

and others, 1994), so a simple expectation would be that the 

Lō‘ihi component would occupy the interior of the ascending 

Hawaiian mantle plume (DePaolo and Stolper, 1996). The 

interior of the plume would be the hottest and undergo the 

greatest decompression, so another expectation is that it 

would melt to the greatest amount, in which case the Lō‘ihi 

component might dominate the culminating shield-stage, not 

the inception-stage, of volcanism. One proposed resolution is 

that the He becomes mobile very early and deep in the melting 

process, possibly owing to high CO2 contents (Gerlach and 

others, 2002), and that the upwelling source is inclined as a 

result of shear by the overriding Pacific Plate, so that initial 

melts carrying the He signal from the interior of the source 

arrive early, supplying inception-stage volcanoes, whereas 

most of the melting of the rising source takes place at shallower 

levels and displaced laterally to a position beneath the shieldstage volcanoes (DePaolo and Stolper, 1996; DePaolo and 

others, 2001). An alternative resolution is the experimental 

discovery (Parman and others, 2005) that, during partial melting 

at elevated pressures, He is retained in olivine more readily than 

is U; therefore, in regions that have undergone high-pressure 

melting, melt loss may have elevated He/U, resulting in high 

values of 3He/4He over time. Under this interpretation, the 

high 3He/4He of the Lō‘ihi component allows for an ancient 

high-pressure melting event, consistent with the anomalously 

low H2O concentrations of Lō‘ihi magmas (Dixon and Clague, 

2001) and the low He concentrations of high 3He/4He magmas 

worldwide (Anderson, 1998). Although the Lō‘ihi source 

appears to be nonpristine, and diametrically different proposals 

for the spatial configuration of Hawaiian source components 

remain unresolved, the general association of high 3He/4He and 

ocean-island volcanoes supports the interpretation that their 

sources are mainly deeper than those that supply the mid-ocean 

ridge system (Hart and others, 1992).
272 Characteristics of Hawaiian Volcanoes

Petrologic Constraints on Summit 

Reservoirs at Kīlauea and Mauna Loa

This chapter has reviewed petrologic features of tholeiitic 

shield-stage magmas stored in, and erupted from, the summit 

reservoirs of Kīlauea and, to a lesser extent, of Mauna Loa. 

Here, these observations will be used to constrain a model of 

the summit reservoir systems, with particular emphasis on 

Kīlauea. Critical observations include:

(1) Individual magma batches at Kīlauea succeeded 

each other at intervals of 2–7 years from 1952 to 1982 (table 

1), with residence times of the batches within the volcano 

(where definable) of ~10 years (table 1; see also Wright and 

Klein, 2010). Pietruszka and Garcia (1999a) and Pietruszka 

and others (2001) have shown that trace element and isotope 

compositions also vary for the entire historical period at 

Kīlauea (1790 to the present), although not in parallel with 

batches defined by major-element chemistry. 

(2) In contrast to the magma batches, the residence time 

for crystals, as determined by Mangan (1990) for olivine in the 

1959 summit eruption, is one decade, and Pietruzka and Garcia 

(1999b) estimate magma storage times of several decades. The 

time interval between batches is thus shorter than the lifetime 

of each batch. Also, olivine may be retained in the reservoir 

somewhat longer than are the individual melt batches. 

(3) Variation in the composition of melts erupted 

effusively within Kīlauea’s caldera has been subtle from 1912 

to the present, with melts becoming slightly more magnesian 

(and, hence, slightly hotter, per Helz and Thornber, 1987) from 

1924 on (fig. 3). These melts, which are almost crystal-free 

and lie at the end of the olivine-control line (~7 weight percent 

MgO), presumably represent the magma at the top of the 

summit reservoir.

(4) The bulk composition of olivine in the summit 

reservoir is Fo86–87 at both Kīlauea and Mauna Loa, and all 

historical summit lavas exhibit olivine control (Wright, 1971; 

Rhodes and Hart, 1995). The results of Sparks (1990) show 

olivine control at Fo86.5 for young but pre-1790 subaerial 

Mauna Loa lava, while figure 11 shows that the same is 

true for Kīlauea’s Uēkahuna laccolith. The laccolith, which 

predates the formation of Kīlauea’s caldera at 1470–1490 C.E. 

(Swanson and others, 2012), provides evidence that olivine 

from Kīlauea’s summit reservoir has also had an average 

composition of Fo86–87 for many centuries. 

(5) Recent data document the range of melt compositions 

that may exist at any one time in the summit reservoir. 

Extracaldera summit lavas, such as those of the 1959 Kīlauea 

Iki eruption and the 1971 and 1974 summit eruptions, contain 

melts with a much wider range of MgO contents than observed 

in the intracaldera lavas, as have explosive eruptions, including 

the 1924 eruption and various prehistoric tephras (fig. 4). 

Calculations using the equilibrium constant for the 

exchange of Fe and Mg between melt and olivine (Roeder and 

Emslie, 1970), and using the most magnesian melt compositions 

(9.7–10.4 weight percent MgO) from the 1959 eruption, show 

that Fo86–87 is in equilibrium with those melts when Fe2O3 is 15±5 

percent of the iron present. This is somewhat higher than fayalitemagnetite-quartz (FMQ) conditions and is a widely observed 

level for Kīlauea basalts (Carmichael and Ghiorso, 1986) and 

gases (Gerlach, 1993). The fit suggests that melts similar to the 

most magnesian 1959 melts, which fall within the upper end 

of the MgO range of the various tephras, are appropriate to be 

interstitial to olivine (Fo86-87) at the base of the reservoir. Given 

the rate at which olivine reequilibrates as interstitial melts evolve 

(shifting from Fo86–87 to <Fo81 in 15–30 years in Kīlauea Iki), the 

fact that olivine composition is constant over centuries requires a 

continuous flux of MgO-rich melt into the base of the reservoir at 

both Kīlauea and Mauna Loa. 

The Structure of Kīlauea’s Summit Reservoir in 

Light of Petrologic Observations 

Kīlauea’s very active summit reservoir leaked lowMgO but olivine-controlled lava almost continually through 

the 19th and 20th centuries; the only significant breaks in 

summit eruptive activity have been the intervals 1934–52 and 

1982–2008. But why does the summit rarely erupt anything 

other than the limited range of melts illustrated in figure3? 

And why is the average olivine composition in more olivinerich lavas Fo86-87, when we rarely see intracaldera melts with 

the necessary MgO content? The uppermost and lowermost 

parts of Kīlauea’s summit reservoir have both been consistent 

in character for several centuries. They are also very different, 

raising the question of what physical arrangement is consistent 

with their persistent differences. 

Kīlauea’s summit reservoir was first envisaged as a plexus 

of sill-like magma bodies because of the complex pattern of 

summit inflation that preceded the 1967–68 summit eruption 

(Fiske and Kinoshita, 1969; Kinoshita and others, 1969). Other 

studies, depending on geophysical observations (Duffield and 

others, 1982; Ryan and others, 1983), led to similar models 

of a segmented magma reservoir, or even multiple discrete 

but interconnected reservoirs (Poland and others, this volume, 

chap. 5). Segmentation of the reservoir would allow different 

levels in the reservoir to contain persistently different melt 

compositions. It would also allow the distinctive magma 

batches observed at Kīlauea (table 1) to be preserved as they 

move through the summit reservoir. 

The variation of density of Kīlauea melts with temperature 

may also be relevant. Figure 22 shows densities of glasses from 

the 1959 eruption, together with glass densities for selected 

layers from the Keanakāko‘i and Pāhala tephras, calculated 

using the data of Lange and Carmichael (1990). The olivinecontrolled melts have temperatures of ~1,170–1,280 °C (Helz 

and Thornber, 1987) and their densities decrease as temperature 

decreases, as would be expected for a sequence of melts 

crystallizing olivine (Sparks and others, 1980; Stolper and 

Walker, 1980; Sparks and Huppert, 1984). 

The presence of dissolved water will lower the 

calculated melt density values considerably (Ochs and 
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 273

2.65

2.67

2.69

2.71

2.73

2.75

1,050 1,100 1,150 1,200 1,250 1,300

Melt density, in grams per cubic centimeter

Temperature, in degrees Celsius

1959 Keanakāko‘i

Pāhala

Experiments

Lange, 1999). Clague and others (1995b) have estimated 

the density of primary Kīlauea melt (MgO=16.5 percent) 

to be 2.66–2.68 g/ cm3 at 0.37 percent H2O, as cited above. The 

data in figure 22 stop at 13 weight percent MgO and T=1,280 

°C, so the uppermost densities are not directly comparable. 

Although the scale on the ordinate in figure 22 would change 

at such H2O contents, the slopes for olivine-controlled melts 

would be steeper, not flattened or reversed (Lange and 

Carmichael, 1990; Helz, 2009). Also, although stored magma 

loses CO2 early, H2O is largely retained in the melt to very 

shallow levels (Gerlach and Graeber, 1985; Gerlach and 

others, 2002), being released in significant quantities only 

during eruption (for example, Wallace and Anderson, 1998). 

Magma consists of crystals and gas bubbles, as well as 

melt—here, again, the expected gradients (more olivine at the 

bottom, less at the top; gas bubbles produced during storage 

either a uniform flux or perhaps slightly increasing toward the 

top) would result in higher bulk density at the bottom of the 

reservoir and lower bulk density at the top. The principal source 

of instability in the magma reservoir (between major summit 

eruptions) would appear to be the continual input of fresh 

magma from the mantle. 

As discussed in Gerlach and others (2002), the presence 

of CO2 bubbles is critical to the buoyancy of intruding primary 

magma when the crystal content of the stored magma is <5 

percent. Crystal contents in the deepest part of the reservoir are 

probably higher than 5 percent, but they are very low at the top, 

with the crystallinity at intermediate levels being unknown. 

A further consideration is that the specific molar volume of 

CO2 increases rapidly at pressures below 2 kbar (Bottinga 

and Javoy, 1989), which is roughly the pressure at the base 

of Kīlauea’s summit reservoir (Ryan, 1987); thus, a slow but 

steady input of primary magma into the lowest part of the 

reservoir should produce considerable turbulence in the deepest 

segment of that reservoir. The rapid increase in buoyancy of 

the CO2 bubbles once they enter the reservoir, however, should 

Figure 22. Graph of melt density versus temperature for four 

sets of glass analyses. Densities calculated using the procedure 

of Lange and Carmichael (1990), with H2O and Fe2O3 set to zero 

in the absence of data. Red dots, glasses from 1959 eruption 

samples (Helz, 2009; and R.T. Helz, unpublished data). Green 

triangles, glasses from selected layers in the Keanakāko‘i tephra

(Helz and others, 2014). Blue diamonds, selected layers in the 

Pāhala Ash (Helz and others, 2014). Black squares, glasses

from Kīlauea Iki sample KI75-1-143.8 experiments (Helz and

Thornber, 1987). Darker symbols indicate melts crystallizing 

only olivine+chromite; lighter symbols indicate melts crystallizing

augite±plagioclase±Fe-Ti oxides in addition to olivine. Lines 

above 1,170 °C are least-squares fits for the melts crystallizing

only olivine±chromite; dotted lines indicate subsequent path of 

density versus temperature for more differentiated samples.

facilitate their separation from the input magma. The resulting 

bubble plumes might entrain significant crystals+melt as 

they rise further in the reservoir (Sparks, 1978). The vorbs in 

Kīlauea Iki formed by a similar process (Helz and others, 1989) 

and have entrained abundant coarse olivine (fig. 12A), but, 

in the lava lake, the bubble plumes traversed a semicoherent 

crystal mush of olivine+augite+plagioclase. In the summit 

reservoir, by contrast, bubble plumes would encounter only 

olivine±chromite in rapidly diminishing quantities at higher 

levels. Efficient segregation of bubbles in magma of decreasing 

crystal content would tend to restrict turbulence/entrainment 

to the immediate vicinity of the bubble plume. Furthermore, if 

the summit reservoir is segmented, it is not clear how transfer 

of turbulence from the segment at the level of primary input to 

higher segments would occur. One observation that suggests 

limited disruption of the summit reservoir by the normal level 

of magma input is that the CO2 plume apparently carries no 

melt with it, and has been emitted steadily whether lava is 

erupting at the summit or not (Gerlach and others, 2002). 

To summarize, geophysical evidence favors a 

segmented summit reservoir, with septa of olivine-rich and 

(or) partly molten material surrounding melt-rich bodies. A 

cartoon version of such a reservoir (fig. 23) envisions small, 

discontinuous but connected bodies of melt lying within the 

hot aseismic region shown in Ryan (1988). Such a structure 

is consistent with the observed preservation of distinct 

magma batches at Kīlauea and the observed persistent 

differences between the top and the base of the reservoir. The 

eruption of more magnesian melts just outside the caldera 

(as in 1959, 1971, and 1974) also argues for some degree 

of separation or stratification within the reservoir. Internal 

segregation by weak septa would be supported or enhanced if 

the melt column at Kīlauea tends to be gravitationally stable. 

In any case, a segmented reservoir structure would inhibit 

complete overturn of the summit reservoir under normal 

conditions.
274 Characteristics of Hawaiian Volcanoes

The Summit Complex Disrupted 

Relatively young, explosive eruptions at Kīlauea’s 

summit include the 1924 eruption, the Keanakāko‘i eruptions 

(which occurred from 1460 to 1790 C.E.; Swanson and others, 

2012), and the Kulanaokuaiki eruptions (400–1000 C.E.; 

Fiske and others, 2009). These, together with the 1959 summit 

eruption and the still older Pāhala Ash, show wide ranges of 

melt compositions (fig. 4). 

The Keanakāko‘i eruptions are interpreted by Mastin 

and others (2004) to have resulted from rapid injection of 

magma to high levels, presumably overstepping the storage 

capacity of the summit reservoir. The frequency distributions 

for all eruptions in figure 4, however, are dominated 

volumetrically by melts at intermediate MgO contents, which 

cannot reasonably be explained by mixing between the sparse 

low-MgO and high-MgO tails of the distributions during 

fragmentation and rapid eruption. It seems more likely that 

most melt compositions (excepting those in the high-MgO 

tails) existed in the reservoir and were entrained during these 

explosive eruptions. By contrast, the 1924 tephra resulted from 

collapse of the upper part of the reservoir, with ejection of 

only traces of juvenile material, which nevertheless included 

melts with as much as 8.7 percent MgO. The presence of 

relatively magnesian liquids in the 1959, 1971, and 1974 

summit eruptions, as well, suggests that such melts are 

commonly present in the existing reservoir. 

The 1959 Kīlauea Iki Eruption

The 1959 eruption was closely monitored and observed, 

and many samples of scoria were collected in real time 

(Murata and Richter, 1966; Richter and others, 1970), thus 

providing a view into the reservoir during the eruption. One 

component of the 1959 magma was tracked seismically from 

below 60 km to the surface (Eaton and Murata, 1960). It 

intersected Kīlauea’s summit reservoir on its north side and 

interacted with stored magma related to the 1954 summit 

batch to produce a range of mixed magmas (Wright, 1973; 

Helz, 1987a). This juvenile (1959E) component was injected 

into the erupting volume (which was dominated by the stored 

1959W component) at several stages during the eruption 

(Eaton and others, 1987; Helz, 1987a). Unusually high fountain heights (300–580 m; Richter and others, 1970) showed 

that the eruption was exceptionally gas-rich. Studies of the 

olivine phenocryst population also suggest that part of the 

magma bypassed significant storage time (Harris and Anderson, 1983; Helz, 1987a; Anderson and Brown, 1993; Wallace 

and Anderson, 1998). 

 The observed succession of scoria erupted in 1959 

suggests that the summit reservoir contained discrete magma 

bodies, in addition to entrainable Fo86–87 at its base. The 

early lavas erupted were dominated by the juvenile (1959E) 

component, and include some of the hottest magmas of the 

eruption (9.7–10.4 weight percent MgO, T=1,210–1,220 °C). 

These were followed abruptly by lavas (erupted on November 

20–21 and again on November 26) that contained only 6–9 

percent by weight of the juvenile component (Wright, 1973). 

This group of samples provides a good look at one body of 

the stored (1959W) magma, which had been in the reservoir 

for at least five years but was still uniformly hotter and more 

magnesian (8.5–8.6 weight percent MgO, T=1,185–1,187 °C) 

than melts in the top of the reservoir. 

After a pause in the eruption, renewed activity on 

November 28–29 produced melts 10–30 °C cooler than those 

from the previously sampled body of stored (1959W) magma. 

The source body for these phase 3 lavas must have been 

separate from the body sampled earlier first, because (unlike 

any but the two earliest 1959 scoria) the lavas erupted from it 

were not hybrids (Wright, 1973) and second, because it was 

too cool to be parental to any of the later (phase 4–17) hybrids, 

which have glass MgO contents and temperatures similar to, 

or higher than, the first body of 1959W magma (Helz, 1987a, 

2009). 

The dominant melts (fig. 4) of the 1959 eruption should 

be in equilibrium with Fo83–84 at the redox conditions inferred 

for the deepest parts of the reservoir (Fe2O3 ~15±5 percent of 

the iron present, as discussed above). Nevertheless, the overall 

average for the eruption is Fo86–87 (fig. 11), and olivine <Fo85

is sparse in the observed phenocryst population (Helz, 1987a). 

This is consistent with the two stored magma bodies tapped 

during the eruption being isolated from the deeper olivine-rich 

zone of the reservoir, suggesting that three distinct levels of 

the reservoir were tapped during the 1959 eruption. 

Figure 23.  Cartoon showing Kīlauea’s summit reservoir in a northsouth cross section that intersects neither of the volcano’s rift zones,

modified from Fiske and Kinoshita (1969). The colored field shows

the seismically quiet volume described by Ryan (1988), which, by 

implication, delineates the summit reservoir as lying between 2 and 6 

km below the surface. The magma bodies, as drawn, are complexly 

segmented and lie within the seismically quiet volume. 

2 km

6 km

Halemaʻumaʻu N S
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 275

Evolution of the Summit Reservoirs Through Time 

The pattern of eruption at Kīlauea suggests that the 

distribution and extent of melt-rich bodies in the summit 

reservoir may change with time. First, the summit reservoir 

gradually recharged between 1924 and 1968, with high 

segmentation and gradually expanding magma storage. Later, 

as eruptive activity at Kīlauea shifted more strongly to the 

East Rift Zone, magma batches became less distinctive, 

although the entire system (summit and both rift zones) saw 

magma storage and activity during 1971 and 1974. Since 

1982, however, East Rift Zone activity has dominated. A 

reasonable consequence of the long-term, ongoing East Rift 

Zone eruption would be reconfiguration of the summit 

reservoir to facilitate transport to the eruption site on the East 

Rift Zone. This, in turn, could lead to atrophy of parts ofthe 

reservoir, so that its overall capacity would be less than in 

the 1960s and 1970s and the only melts in storage would be 

related to the ongoing East Rift Zone eruption. However, a 

recent microgravity study by Johnson and others (2010), 

covering the time span 1975–2008, suggests that magma has 

been accumulating at a depth of 1 km below the east edge of 

Halema‘uma‘u Crater over much of this same period. The 

summit vent that opened in early 2008 is almost directly over 

the inferred location of the accumulation zone, and Johnson 

and others (2010) suggest that this shallow body is the source 

of the current summit activity. The compositions of tephra 

from the current summit eruption are identical to coevally 

erupted material at Pu‘u ‘Ō‘ō (Rowe and others, 2009), 

suggesting that the extent of cooling en route to the shallow 

summit reservoir is similar to that which occurs in the main 

flux of magma out to Pu‘u ‘Ō‘ō. 

Mauna Loa has seen only limited activity in the past 60 

years, in contrast to earlier periods; nevertheless, the lavas 

erupted have been extremely uniform. If the magma in its 

reservoir is crystallizing plagioclase (not true at Kīlauea at 

present; fig. 9), there may be spontaneous mixing of melts at the 

top of the reservoir, driven by the reversal in melt density versus 

temperature seen at that point in basalt crystallization (fig. 22). 

Another explanation would be that Mauna Loa’s mantle source 

has been more uniform than Kīlauea’s over this period. 

Long-term variations in trace element and isotopic 

compositions of basalts at Kīlauea and Mauna Loa shed light 

on the variability of their magmatic behavior over the past two 

centuries. Rhodes and Hart (1995) found analytically significant 

variations in trace element levels of Mauna Loa lavas, evenin 

the absence of magma batches definable by major element 

chemistry. Pietruzka and Garcia (1999a,b) and Pietruzka and 

others (2001) show similar variability in trace element and 

isotopic signatures at Kīlauea, again decoupled from magma 

batches defined by major elements. This variability has been 

attributed to changes in the degree of partial melting and (or) 

variations in exact source composition. At Mauna Loa, shifts 

consistent with increasing melting and magma supply preceded 

the 1868 seismic crisis, with incompatible trace element levels 

rising (and magma supply dropping) afterward (Rhodes and 

Hart, 1995). At Kīlauea, the major event was the 1924 collapse 

of the summit magma column (Pietruzka and Garcia, 1999a,b; 

Pietruzka and others, 2001), with incompatible trace elements 

rising (and the extent of melting and magma supply decreasing) 

before the event and recovering afterward. Evidently, the entire 

magmatic system at both volcanoes is continuous enough that 

any changes in magma production and (or) supply from the 

source region have large and geologically rapid impacts on the 

upper reaches of the volcanic plumbing, as documented recently 

by Poland and others (2012). 

Summary and Conclusions

Study of the youngest Hawaiian volcanoes over the 

past 100 years has produced a flood of observations relevant 

to understanding basaltic volcanism, especially, but not 

exclusively, at shield volcanoes. The overarching, distinctive 

feature of this effort on the part of Hawaiian Volcano 

Observatory staff and others is the attempt to observe, 

document, and quantify basaltic eruptions as completely 

as possible. Highlights among the results reviewed here 

include:

(1) Early recognition of magma mixing at Kīlauea in both 

summit and rift zone lavas, which is now known to be pervasive 

at this and other volcanoes.

(2) New data that show that the range of melts erupted 

effusively within Kīlauea Caldera do not represent the entire 

range of melts in Kīlauea’s summit reservoir during the 20th 

century, as evidenced by the eruption of extracaldera summit 

lavas with glass MgO contents of 8–9 percent by weight in 

the 1959, 1971, and 1974 eruptions and in the rare presence of 

similar material in the 1924 tephra. 

(3) Recognition that young, prehistoric tephras at Kīlauea 

have similar or larger ranges of glass MgO contents. The 

frequency distributions of glass MgO contents suggest that 

most of the range of melts sampled during Kīlauea’s explosive 

eruptions existed in the summit reservoir before eruption. 

(4) Documentation of the reequilibration of olivine 

during closed-system crystallization of Kīlauea Iki lava lake, 

which shows that olivine compositions shifted from Fo86–87

to <Fo81 in 15–30 years. The Fo86–87 control lines widely 

observed in lavas from both Kīlauea and Mauna Loa therefore require that the lower parts of both summit reservoirs 

see a continual influx of melt of appropriate composition.

(5) Olivine control lines for subaerial rift zone lavas, 

whether olivine-poor or picritic, are also Fo86–87 at both 

volcanoes. Most rift zone eruptions thus appear to be fed 

from no deeper than the middle to lower parts of the summit 

reservoir (4–6 km depth at Kīlauea). 

(6) Olivine control lines for submarine picritic lavas are 

Fo88 at Kīlauea and Fo89 at Hualālai, implying that melts below 

the level of the summit reservoirs, including in the deeper parts 

of the rift zones, are, on average, more magnesian than the 

subaerially erupted melts shown in figure 4. 
276 Characteristics of Hawaiian Volcanoes

(7) The most magnesian submarine glasses on the 

submerged part of Kīlauea’s East Rift Zone (14.7–15.0 weight 

percent MgO) are close to inferred primary compositions 

for Kīlauea and Mauna Loa magmas (14–20 weight percent 

MgO). However, picritic lavas observed at these and other 

Hawaiian volcanoes—even though they have similar bulk 

MgO contents—have been interpreted, without exception, as 

cumulates and not as quenched ultramafic liquids. 

The accumulated observations of the past 100 years are 

consistent with the view that the magmatic systems of Kīlauea 

and Mauna Loa are extremely open, locally segmented but 

well connected, and continually refreshed with new magma 

from the mantle. 

Acknowledgments

We wish to acknowledge Mel Beeson’s work in 

analyzing the Halema‘uma‘u samples collected by Thomas 

Jaggar during the early 20th century; the resulting data 

are presented here for the first time. Similarly, we wish 

to acknowledge Tim Rose (Smithsonian Institution) for 

permission to include glass compositional data from the 1924 

eruption in this report. We also thank Jim Moore (USGS, 

Menlo Park, Calif.) for permission to use his sketch of 

the segregation veins exposed in the cliff face through the 

prehistoric Makaopuhi lava lake and Harvey Belkin (USGS, 

Reston, Va.) for his assistance in analyzing olivines from 

Kīlauea Iki lava lake. This chapter has been improved by 

reviews from Tom Wright and Mike Clynne.

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Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 289

Appendix. Microprobe Analyses of Glasses from 

Kīlauea Volcano

The following tables present microprobe analyses of glasses from Kīlauea 

Volcano. These include samples collected at Halema‘uma‘u from 1912 through 

1934 (table A1), tephra samples from the 1924 summit eruption (table A2), and 

samples from the summit eruptions of 1954, 1971, and 1974 (table A3).
290 Characteristics of Hawaiian Volcanoes

Sample ID Date 

erupted

No.

points 

analyzed

SiO2 Al2O3 FeO MgO CaO Na2O K2O TiO2 P2O5 MnO S Total Comments

95TAJ-3 1912 20 50.73 14.05 11.46 6.84 11.32 2.43 0.56 2.86 0.28 0.17 0.02 100.72 Pele’s hair

95TAJ-4 1913 16 50.67 13.92 10.77 7.46 11.72 2.29 0.49 2.71 0.26 0.17 0.01 100.47 Pele’s hair

95TAJ-5 Jan 

1917 16 50.77 14.06 11.01 7.19 11.80 2.34 0.51 2.72 0.26 0.16 0.01 100.83 Driblet from Halema‘uma‘u 

95TAJ-23 Jan 

1917 16 50.34 13.95 10.83 7.19 11.79 2.29 0.49 2.71 0.27 0.16 0.01 100.03 Overflow

95TAJ-24 Jan 

1917 16 50.56 13.86 10.86 7.19 11.83 2.31 0.51 2.69 0.27 0.15 0.01 100.21 Spatter lava Halema‘uma‘u 

95TAJ-2 Mar 16, 

1918 16 50.68 13.80 11.67 6.72 11.40 2.40 0.56 2.91 0.28 0.17 0.01 100.60 Lava spatter fragments

95TAJ-15 Jan 

1918 13 49.69 13.71 11.53 6.78 11.37 2.35 0.52 2.79 0.27 0.16 0.008 99.17 Spatter caught red-hot

95TAJ-16 Feb 

1918 15 50.29 13.57 12.35 6.44 11.01 2.49 0.56 3.07 0.28 0.17 0.005 100.24

Glassy Kīlauea 

lava collected 

with stick

95TAJ-21-3 Jan 

1918 12 51.11 14.20 11.72 6.99 11.49 2.52 0.54 2.77 0.25 0.18 0.005 101.78 Splash on ledge 8 ft above lake

95TAJ-21-4 Jan 

1918 4 49.64 13.76 11.55 6.74 11.52 2.42 0.52 2.80 0.26 0.15 0.007 99.37 Splash on ledge 8 ft above lake

95TAJ-29 1918 16 50.66 13.88 11.10 6.95 11.77 2.36 0.51 2.89 0.29 0.18 0.00 100.59 (no description)

95TAJ-18 Oct 20, 

1919 15 50.16 13.92 11.79 6.75 11.41 2.40 0.49 2.74 0.26 0.17 0.008 100.10 Splash over N (W?) bank

95TAJ-22 1919 16 50.64 13.76 11.64 6.78 11.47 2.40 0.55 2.90 0.29 0.17 0.01 100.61 Flow glassy 

selvage

95TAJ-28 Mar 17, 

1919 16 50.49 13.81 11.61 6.68 11.29 2.41 0.56 2.91 0.30 0.17 0.01 100.24 Splash are next to NE cone

95TAJ-34 Oct 6, 

1919 16 50.48 13.99 11.22 6.93 11.61 2.35 0.53 2.81 0.27 0.17 0.01 100.37 Flow from Jaggar Lake

95TAJ-6 Nov 

1920 16 50.43 14.01 10.98 7.16 11.70 2.34 0.49 2.68 0.26 0.17 0.01 100.23 Pele’s hair

95TAJ-7 Mar 2, 

1921 16 50.41 14.07 10.97 7.12 11.69 2.38 0.50 2.77 0.28 0.17 0.01 100.37 Glass lapilli etc. Halema‘uma‘u 

95TAJ-19 1921 14 50.32 14.00 11.30 6.86 11.48 2.44 0.52 2.83 0.25 0.15 0.012 100.16 Halema‘uma‘u 

95TAJ-8 Mar 21, 

1921 16 50.39 14.14 10.96 7.03 11.61 2.38 0.50 2.77 0.27 0.16 0.01 100.28 Halema‘uma‘u spatter

95TAJ-9 Aug 28, 

1923 16 50.33 13.83 11.40 6.73 11.35 2.44 0.57 3.02 0.30 0.18 0.01 100.16 Pele’s hair

95TAJ-32 1924 16 50.24 13.88 11.61 6.66 11.30 2.43 0.58 3.03 0.30 0.17 0.01 100.21 1924 lava

95TAJ-10 Jul 25, 

1929 16 49.82 13.53 11.15 7.49 11.72 2.32 0.55 2.72 0.27 0.16 0.01 99.74 Glassy cinders

95TAJ-11 1934 10 50.17 13.69 11.41 7.28 11.57 2.27 0.52 2.63 0.24 0.17 0.013 99.97 Glassy cinders

95TAJ-12 Sept 

1934 14 50.23 13.70 11.59 7.19 11.56 2.32 0.52 2.66 0.26 0.17 0.012 100.21 Glassy cinders

95TAJ-13 1868 16 51.16 13.60 12.24 6.27 10.73 2.47 0.49 2.72 0.27 0.18 0.012 100.14 1868 Kīlauea 

inner walls

Table A1. Microprobe analyses of glasses in Halema‘uma‘u samples collected by T. A. Jaggar from 1912 through 1934.

[Analyses by M.H. Beeson; “Comments” are as written on sample bag labels, per M.H. Beeson; oxide compositions in weight percent ]
Petrologic Insights into Basaltic Volcanism at Historically Active Hawaiian Volcanoes 291

Sample ID Glassy Pele’s

tear 36

Glassy Pele’s

tear 37

Glassy Pele’s

tear 38

Glassy 

shards, with 

skeletal 

plagioclase

Partly 

crystalline 

Pele’s tears

No. tears 1 1 1 11 5

SiO2 50.15 50.60 49.48 50.43 49.82

TiO2 2.65 2.60 2.69 2.92 3.61

Al2O3 13.41 13.46 12.96 13.63 12.73

FeO 10.93 11.37 11.45 11.41 13.15

MnO 0.14 0.19 0.14 0.16 0.20

MgO 7.39 7.61 8.72 7.02 5.85

CaO 11.59 11.55 11.49 11.70 10.39

Na2O 2.22 2.27 2.25 2.29 2.43

K2O 0.52 0.53 0.56 0.54 0.76

P2O5 0.26 0.25 0.29 0.26 0.39

Total 99.27 100.42 100.03 100.37 99.33

T (MgO) in 

degrees Celsius 1,163.00 1,168.00 1,191.00 1,154.00 1,132.00

[Analyses by T.R. Rose; oxide compositions in weight percent; T (MgO), glass quenching temperature using Helz 

and Thornber (1987) calibration] 

Table A2. Microprobe analyses of glasses from tephra of the 1924 summit eruption.
292 Characteristics of Hawaiian Volcanoes

Table A3. Microprobe analyses of glasses from the 1954, 1971, and 1974 summit eruptions. 

[Analyses by R.T. Helz; oxide compositions in weight percent; --, no data; T(MgO), glass quenching temperature using Helz and Thornber (1987) 

calibration; NMNH, Smithsonian National Museum of Natural History]

Sample ID 1954A 1954B Kil 71-3 Kil 71-5 Kil 71-6 Kil 71-8 Kil-774-2 Kil-774-3 Kil-774-6a Kil-774-4 Kil 74-v2

No. points 11 13 11 12 10 12 16 12 11 15 19

SiO2 50.18 49.89 50.25 50.38 50.73 50.72 50.73 50.22 50.84 50.89 50.51

TiO2 2.86 2.82 2.46 2.36 2.72 2.76 2.39 2.32 3.05 2.86 2.26

Al2O3 13.97 13.95 13.45 13.22 13.94 13.80 13.55 12.98 13.29 13.74 13.36

Cr2O3 0.01 0.01 0.07 0.07 0.02 0.02 0.07 0.07 0.01 0.02 0.08

FeO 11.17 11.19 11.29 11.36 11.12 11.45 11.19 11.28 12.14 11.90 11.19

MnO 0.19 0.16 0.17 0.18 0.17 0.19 0.18 0.19 0.21 0.20 0.17

MgO 6.68 6.73 8.05 8.87 6.99 6.42 7.90 9.24 6.28 6.43 8.18

CaO 11.29 11.39 11.38 11.10 11.47 11.15 11.36 10.96 10.68 10.82 11.26

Na2O 2.26 2.29 2.13 2.01 2.25 2.28 2.12 2.09 2.26 2.27 2.11

K2O 0.56 0.54 0.42 0.41 0.51 0.54 0.42 0.42 0.59 0.53 0.42

P2O5 0.26 0.26 0.20 0.21 0.24 0.28 0.21 0.22 0.28 0.27 0.21

Total 99.43 99.23 99.87 100.17 100.16 99.61 100.12 100.57 99.04 100.09 99.75

T(MgO) 

in degrees 

Celsius

1148 1149 1177 1193 1154 1145 1173 1203 1140 1143 1179

Eruption 

date

6/

1954

6/

1954

8/14/

1971

8/14/

1971

8/14/

1971

8/14/

1971

7/19/

1974

7/20/

1974

7/20/

1974 

7/19/

1974

7/

1974

Location caldera floor extra-caldera caldera floor northern vents caldera 

floor

southern 

vents

northern 

vents

Whole rock 

analysis in Wright, 1971 Duffield and others, 1982 Lockwood and others, 1999 --

NMNH 

sample ID 116859-20 116859-20 116095-3 116095-5 116095-6 116095-8 116125-2 116126-3 116126-7 116125-4 --
Hawaiian Volcano Observatory scientists Margaret Mangan (front) and Christina Heliker (back) use a steel cable with a hammerhead 

attached to collect a lava sample from a skylight above an active lava tube. USGS photograph by J.D. Griggs, December 5, 1990.
Thomas A. Jaggar, Jr., collects a gas sample from the surface of the Halema‘uma‘u lava lake at the summit of Kīlauea Volcano. The gas 

samples collected at this time, done with the guidance of E.S. Shepherd of the Carnegie Institution, eventually informed the Kīlauea volatile 

budget model that is arguably the single most important accomplishment of the first 100 years of Hawaiian volcanic gas studies. Note 

HVO Technology Station in the center background. The station is also shown at the front of chapter 1, but at the time of the photo above, 

Halema‘uma‘u had filled to overflowing. The building was disassembled and moved to the rim of Kīlauea Caldera a few months after this 

photo was taken. USGS photograph, March 3, 1918 (photographer unknown).
1

U.S. Geological Survey.

The smell of sulphur is strong, but not unpleasant 

to a sinner.

 —Mark Twain, Letters from Hawaii, 1866

Abstract

The first volcanic gas studies in Hawai‘i, beginning 

in 1912, established that volatile emissions from Kīlauea 

Volcano contained mostly water vapor, in addition to 

carbon dioxide and sulfur dioxide. This straightforward 

discovery overturned a popular volatile theory of the day 

and, in the same action, helped affirm Thomas A. Jaggar, 

Jr.’s, vision of the Hawaiian Volcano Observatory (HVO) 

as a preeminent place to study volcanic processes. Decades 

later, the environmental movement produced a watershed 

of quantitative analytical tools that, after being tested 

at Kīlauea, became part of the regular monitoring effort 

at HVO. The resulting volatile emission and fumarole 

chemistry datasets are some of the most extensive on the 

planet. These data indicate that magma from the mantle 

enters the shallow magmatic system of Kīlauea sufficiently 

oversaturated in CO2 to produce turbulent flow. Passive 

degassing at Kīlauea’s summit that occurred from 1983 

through 2007 yielded CO2-depleted, but SO2- and H2Orich, rift eruptive gases. Beginning with the 2008 summit 

eruption, magma reaching the East Rift Zone eruption 

site became depleted of much of its volatile content at the 

summit eruptive vent before transport to Pu‘u ‘Ō‘ō. The 

volatile emissions of Hawaiian volcanoes are halogen-poor, 

relative to those of other basaltic systems. Information 

gained regarding intrinsic gas solubilities at Kīlauea and 

Mauna Loa, as well as the pressure-controlled nature of gas 

release, have provided useful tools for tracking eruptive 

activity. Regular CO2-emission-rate measurements at 

Kīlauea’s summit, together with surface-deformation and 

other data, detected an increase in deep magma supply more 

than a year before a corresponding surge in effusive activity. 

Correspondingly, HVO routinely uses SO2 emissions to study 

shallow eruptive processes and effusion rates. HVO gas 

studies and Kīlauea’s long-running East Rift Zone eruption 

also demonstrate that volatile emissions can be a substantial 

volcanic hazard in Hawai‘i. From its humble beginning, 

trying to determine the chemical composition of volcanic 

gases over a century ago, HVO has evolved to routinely use 

real-time gas chemistry to track eruptive processes, as well 

as hazards. 

Introduction

Volcanic gases are a dominating force in eruptive 

activity. Although the thermal and density contrasts between 

magma in a volcanic system and the surrounding country 

rock result in the rise of melt from depth to the edifice, the 

exsolution, expansion, rise, and separation of volatiles are 

what drive both effusive and explosive eruptions.

Thomas A. Jaggar, Jr., founder of the Hawaiian Volcano 

Observatory (HVO), stated it well: “The observatory worker 

who has lived a quarter of a century with Hawaiian lavas 

frothing in action, cannot fail to realize that gas chemistry 

is the heart of the volcano magma problem” (Jaggar,1940). 

Despite the fervor of this statement, hazardous field conditions 

associated with the sampling of volcanic vents and the early 

evolutionary state of analytical techniques applied to gas 

samples conspired to produce only modest progress in the 

understanding of gas release processes by 1940, when Jaggar 

opined the above remark. 

In this chapter, we trace the progress made at HVO 

in understanding the role of volatile emissions in volcanic 

processes. Gas studies at HVO evolved episodically, not 

unlike the eruptive activity that researchers were tracking. In 

the early days, gas studies relied on volcanic action to provide 

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 7

One Hundred Volatile Years of Volcanic Gas Studies at the 

Hawaiian Volcano Observatory

By A. Jeff Sutton1 and Tamar Elias1
296 Characteristics of Hawaiian Volcanoes

sampling opportunities, as well as the often fitful progression of 

technology, to provide the tools to study volcanism. Because of 

the infrequency of sustained eruptions on Mauna Loa and the 

comparative difficulty of working there, gas studies techniques 

and progress have generally evolved on Kīlauea and then been 

applied to the larger mountain. Paul Greenland reviewed the 

major research on eruptive gases leading up to HVO’s 75th 

anniversary (Greenland, 1987a). In recognition of HVO’s 

centennial and, owing to the obscurity of some of the earlier 

documentation, here we emphasize the formative years of gas 

studies, as well as those years postdating Greenland’s (1987a) 

review. 

HVO gas studies, to date, can be reasonably grouped 

in three distinct epochs of progress, each of which produced 

milestones that helped advance the science. At HVO’s 

inception, Halema‘uma‘u lava lake within Kīlauea’s summit 

caldera provided an excellent environment to observe volcanic 

gas release first hand. As with so much of the progress achieved 

by all the disciplines at HVO, close proximity to eruptive 

activity—literally, right out the back door—and the relative 

passivity of that activity helped workers make considerable 

progress within a comparatively short time. This first epoch saw 

development of the rudiments of gas sampling and analysis, 

and, although the extreme temperatures and the corrosiveness of 

the gases were challenging, the first data from these experiments 

revealed the fundamental ternary chemical makeup of volcanic 

gas emissions.

Opportunities to study eruptive gases at Kīlauea ceased 

as magma withdrew from the summit in 1924, and eruptive 

activity was subdued or absent for nearly 40 years. During the 

gap, the second epoch of growth in knowledge about volcanic 

gases occurred, owing to demand-driven improvements in 

analytical technology. These improvements fueled advances in 

the ability to sample and analyze gases when vigorous eruptive 

activity returned to the summit in the 1960s. 

Beginning in 2000, a sea change occurred in the way 

HVO monitored eruptive activity, caused by a major increase 

in computing and networking capability and especially in 

the centralization of data and capability to visualize it in 

near-real time. These improvements made it easier than ever 

before to track geophysical and geochemical data streams 

contemporaneously during the course of Kīlauea’s longrunning (1983–present) East Rift Zone eruption. This rich, 

real-time analytical environment, in combination with an 

upwelling of new gas-measurement techniques, provided 

opportunities to test theories about magma movement during 

a surge in Kīlauea’s magma supply. Major changes in eruptive 

activity both at Kīlauea’s summit and along the East Rift Zone 

provided the crucial data needed to further test and extend 

previously constructed models. The persistent high level 

of eruptive activity during this third epoch, along with the 

watershed of gas-studies tools and techniques, also occasioned 

a better appreciation of volatiles as a volcanic hazard requiring 

assessment, response, and mitigation, rather than something that 

could be ignored or downplayed in comparison with other, more 

visible volcanic processes. 

Early and Middle Years, 1912–50s: GasSampling Methods and the Ternary 

Compositional Basics

At the time of HVO’s founding, several workers had 

already written about Kīlauea’s volatile emissions. In his book 

“Vestiges of the Molten Globe, Part II,” William Lothian Green, 

a self-taught volcano scientist-observer and longtime resident 

of Hawai‘i, hypothesized that water was not a fundamental 

component of Hawaiian volcanic gas emissions (Green, 1887; 

Day and Shepherd, 1913). Although Green did not actually 

collect or analyze gas samples in the conventional sense, he was 

a careful observer; amongst his many interpretations of Kīlauea 

and Mauna Loa plume clouds, he attributed the visible fume 

above eruptive vents and the Halema‘uma‘u lava lake as simply 

a convective interaction between the tremendous heat given off 

by molten lava and the moisture laden trade winds passing over 

eruptive sources. Years later, Albert Brun, a Swiss chemist who 

made actual field measurements of volcanic gas emanations 

on Kīlauea and other volcanoes, put forth a similar conclusion 

(Brun, 1913). Reginald Daly, a professor at Harvard who visited 

Kīlauea in 1909, was more open-minded on the issue of water 

as a volcanic gas. While acknowledging the conclusions of 

Green and Brun about Kīlauea, Daly believed that volatiles, 

possibly including water, played an important role in volcanism 

(Daly, 1911).

The writings of Green, Brun, and Daly formed the 

knowledge base that fueled the interest and experimental plans 

of A.L. Day and E.S. Shepherd of the Carnegie Institute’s 

Geophysical Laboratory, the first two investigators to make 

gas measurements at Kīlauea under the HVO banner. The 

water question was a popular topic of the day, but Day and 

Shepherd had accepted Brun’s conclusion that Kīlauea’s 

volatile emissions were anhydrous and planned their 

experiments accordingly. 

Day and Shepherd’s gas collection scheme consisted of 

a series of 20 half-liter flasks strung together serially (Day 

and Shepherd, 1913). One end of this sampling string was 

connected to a glass-lined iron sampling pipe, where the gas 

would enter the string; the other end was linked to a handoperated piston pump, used to draw gas through the string of 

flasks. 

Shepherd had spent the second half of 1911 working at 

Kīlauea alongside Frank Perret, conducting observations at the 

Halema‘uma‘u lava lake and assisting in the first in situ lava 

temperature measurements. This experience served him well 

when he returned to the volcano with A.L. Day the following 

year. On May 28, 1912, Day and Shepherd had the good 

fortune of happening upon a volcanic gas sampler’s dream 

vent. On the floor of Halema‘uma‘u, a pressurized opening in 

what might be best described as a hornito vented burning gas 

to the air, offering an opportunity to sample essentially pristine 

volatile emissions. These conditions were exactly what the 

two researchers had been hoping for (fig. 1). They set up 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 297

Figure 1. E.S. Shepherd (left) and F.B. Dodge sample pristine eruptive gases 

at Halema‘uma‘u Crater, May 28, 1912. Dome-shaped hornito just right of center 

accommodated a 0.3-m-long iron sampling pipe that was linked through glass 

tubing to 20 serially connected sampling bottles contained within the crate between 

the two men and, finally, to a hand-operated pump. 

Figure 2. Early 1900s vacuum-type gas-sampling 

bottles. More fragile lower and middle bottles are from 

Shepherd, Day, and Jaggar era, circa 1912–25. Elongate, 

thin glass bulb at right end of each bottle was melted 

or broken to admit gas when placed inside a hot vent, 

attached to end of a 3-m-long pole, then carefully resealed 

by melting after gas was sampled. Top bottle, by Stanley 

Ballard of the University of Hawai‘i, circa 1938, improved 

on earlier ones by being sturdier, with a tip that was 

broken open for sampling using an attached wire. This 

newest bottle was resealed after sampling by closing 

stopcock at base of bulb. 

their collection gear, placed the end of a sampling tube within 

the hornito just past the burning gas jet, and began to pump. 

Because of Brun’s earlier conclusion, Shepherd and Day made 

no provision to collect condensed moisture, because they 

expected no water. 

What happened next surprised them. Large amounts of 

fluid, later determined to be water, began to condense in their 

apparatus under ambient temperature conditions. Within 15 

minutes, they had collected 300 mL of this aqueous condensate, 

along with a substantial volume of noncondensable “dry” gas. 

Most of the condensation occurred in the first few flasks of 

the sampling string because of the large temperature contrast 

between the hot vent and the relatively cool first flasks in the 

string. Thus, instead of having 20 replicate samples of a single 

composition to analyze, thereby ensuring high-precision results, 

Day and Shepherd obtained 20 bottles with varying amounts of 

volatile constituents scattered amongst them.

Although their experimental plan did not anticipate water, 

Day and Shepherd made the best of it. Partial analyses carried 

out 4 days later at the College of Hawai‘i on O‘ahu confirmed 

the presence of CO2, SO2, and CO, along with the water and 

dissolved gases composing the condensate. Because of the 

high water solubility of SO2 and haloacids, the O‘ahu analyses 

were partial and preliminary. The rest of the samples were 

taken back to the Carnegie Institute’s Geophysical Laboratory 

in Washington, D.C., for detailed analysis of both the dry gas 

and the remaining aqueous condensate.

Day and Shepherd’s May 28, 1912, samples affirmed 

not only that water was part of the makeup of pristine volcanic gases but also that it was a principal component, along 

with CO2 and SO2. Thus, in spite of their compromised gascollection apparatus, they readily characterized the volatile 

emissions at Kīlauea as a ternary mixture (H2O, CO2, and SO2) 

with minor amounts of other species, including CO and H2. 

The copious amount of water in the May 28 samples propelled 

Day and Shepherd to develop an alternative gas-collection 

strategy because of the large, but incomplete, dissolution of 

acidic gases, like SO2, CO2, HCl, and HF, in the condensate 

fraction. The alternative approach sought to draw the gases, 

including water, into a newly designed, evacuated sampling 

bulb (fig. 2). Although Day and Shepherd recognized water as 

a substantial component of volcanic exhalations, they had yet 

to establish the source mechanism for the presence of water, 

and water’s importance in the physicochemical equilibrium of 

magma. 

A.L. Day returned to the Carnegie Institute’s Geophysical 

Laboratory to resume his post as its director, but E.S. Shepherd kept a close collegial relationship with HVO, returning to 

refine gas-sampling methods and to work with Jaggar in order 

to make this technology part of the fabric of HVO’s research. 

Thus, gas studies became another research tool at HVO for 

understanding volcanic processes. 

Shepherd and Jaggar watched for gas-sampling opportunities, and together they published 26 variously partial and 

full sets of analyses for gas samples collected during 1912–19 

(fig. 3; Shepherd, 1921, 1925; Jaggar, 1940), the bulk of 

which came from a systematic sampling campaign carried 

out by Jaggar in 1918 and 1919. Jaggar, in consultation with 

Shepherd, endeavored to obtain samples uncontaminated by 

atmospheric oxygen, and a subset of these samples was used 

many years later to infer Kīlauea’s magmatic source conditions (Matsuo, 1962; Nordlie, 1971; Gerlach, 1980). 

Shepherd and Jaggar’s sampling focus was on what 

Shepherd called “volcano gases,” or what we would today 
298 Characteristics of Hawaiian Volcanoes

call eruptive gases. Shepherd also acknowledged the potential 

value of studying fumarolic solfatara gases, although he saw 

these gases, as well as those from wells, springs, and mines, 

as a less straightforward path toward understanding volcanic 

plumbing and processes. Shepherd, back at the Geophysical 

Laboratory, was also interested in the gases released by rocks 

heated in a vacuum, seeing these products as the last stop in 

the gas-release process (Shepherd, 1925).

The withdrawal of magma from beneath Halema‘uma‘u 

that preceded the explosive 1924 eruption effectively began a 

hiatus in eruptive-gas sampling at Kīlauea that lasted nearly four 

decades. Shepherd and Jaggar used the interval to carefully consider the implications of the data they had collected (Shepherd, 

1925, 1938; Jaggar, 1940). With the lapse in eruptive-gas sampling, Jaggar adjourned with colleagues S.S. Ballard and J.H. 

Payne of the University of Hawai‘i, who were collecting and 

studying gases from a 21-m-deep well located at Sulphur Banks, 

on the north edge of Kīlauea Caldera (fig. 4). Ballard and Payne 

Figure 3. A. Lancaster (left) and E.S. Shepherd sample gases at 

Halema‘uma‘u in 1917. The tip of an evacuated tube gas-sampling 

bottle, wired to the end of a bamboo pole, was inserted through a 

small, flaming hole in the lava-lake crust. The soft glass window at the 

tip of the bottle melted, allowing magmatic gas to be sucked inside. 

Once filled, the bottle was resealed by folding the melted tip against a 

hot lava surface.

Figure 4. From left to right, T.A. 

Jaggar, Jr., J.H. Payne, and S.S. 

Ballard sample gas from a 21-m-deep 

geothermal well located at North 

Sulphur Banks (on north rim of 

Kīlauea Caldera) circa 1937. Ballard 

operates a hand pump at lower right, 

drawing gas from circular wellhead 

(apparent between Jaggar and 

Payne) through a hidden, ice-chilled 

flowthrough bottle and then through 

a flask and series of wash bottles 

containing KOH solution to remove 

acidic gases and water. Final “dry” 

gas collected in vertical flowthrough 

bottle is located middle right of frame 

between Payne and Ballard. 

were interested, among other things, in improving methods for 

collecting and analyzing volcanic gases. At Sulphur Banks, they 

took advantage of the site as a long-lived gas-emitting feature 

and established the first multiyear gas sampling program there 

(Ballard and Payne, 1940). At this site, they reported an “incident” of increased H2

S emission before the 1940 Mauna Loa 

eruption (Payne and Ballard, 1940). This inferred link between 

Sulphur Banks gas release and Mauna Loa eruptive activity, 

though unsubstantiated mechanistically and never observed 

again, was the first suggestion of eruptive-gas precursors in the 

Kīlauea literature.

Thomas A. Jaggar, Jr., retired in 1940 after publishing his 

paper “Magmatic Gases” in the American Journal of Science, 

where he challenged future volcanologists to consider a list of 

working hypotheses regarding volcanic gas studies (Jaggar, 

1940). Shepherd’s, Jaggar’s, Ballard’s, and Payne’s diligence 

notwithstanding, the available analytical techniques of the day 

were stretched to the limits of detection, pursuant to quantifying 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 299

the chemical constituents in the samples obtained. Further 

advances in sampling and analytical technology would be 

needed if volcanic gas studies were to move forward. 

Technological Renaissance and Eruptive 

Reawakening Fuel Progress in Gas 

Studies: The 1960s Through the 1990s

Much of the U.S.’s focus had been on nationalsecurity issues related to World War II and the Korean War 

rather than on volcanoes during the years that followed 

the Halema‘uma‘u eruptive shutdown. This emphasis on 

national defense, however, led to advances in analytical 

and computational technology throughout the 1950s and 

1960s, the benefits of which spanned civilian, as well as 

military, applications. For example, the Clean Water Act of 

1960, the Clean Air Act of 1970, and the establishment of 

the Environmental Protection Agency and the Occupational 

Safety and Health Administration in 1970 created widespread 

regulatory needs for rapid and accurate analytical methods 

to study water and gas samples (albeit not volcanic ones). 

Gas chromatography (GC), chemically selective electrode 

methodology, and analytical spectroscopy developed rapidly, 

in part to meet the growing environmental requirements. 

These technologies also helped scientists drastically improve 

detection limits and accuracy for analyzing the principal 

constituents of volcanic gases.

Assembling New Tools to Study Gas Release

One example of the leveraging of environmental-science 

capabilities was when Werner F. Giggenbach, a gas-studies 

worker in New Zealand, developed an innovative method to 

sample and analyze volcanic and geothermal gases, which 

revolutionized volcanic gas studies worldwide, as well as at 

HVO (Giggenbach, 1975). His innovation relied on the fact 

that volcanic emissions are composed chiefly of water vapor 

and the so-called acid gases, including SO2, CO2, H2S, HCl, 

and HF. For most volcanic and geothermal systems, these 

species compose more than 95 percent of the emitted volatiles, 

and all of them are highly soluble in caustic (basic) solution. 

As with other investigators studying volcanic emissions, 

Giggenbach wanted a complete gas analysis, including the 

non-acidic gases. Thus, he partially filled his collection flasks 

with caustic NaOH solution and then evacuated the headspace 

of the flask. A sampling tube, usually of an inert material, such 

as aluminum oxide ceramic or corrosion-resistant titanium, 

brought hot volcanic gas to the inverted flask, which, when 

opened, caused the acidic gases to be stripped out by the 

NaOH solution as they were drawn into the bottle (fig. 5; 

Giggenbach, 1975). The non-acidic gases, including H2, 

He, CO, and any unreactive atmospheric constituents, such 

as N2, O2, and Ar, passed unaffected through the solution 

and accumulated in the headspace. These headspace gases 

were analyzed by the newly available GC technique, and the 

acidic gases were determined by wet chemistry methods, 

ion-selective electrodes, and gravimetric analysis. This novel 

approach to sample collection put complete gas analysis 

within the reach of any researcher with access to a relatively 

modest GC, an analytical balance, a pH meter and electrodes, 

a manometer, and standard wet chemical supplies. This 

powerful new analytical technique rapidly became popular, 

and Kīlauea proved itself an excellent place to apply it for 

studying volcanic gases.

Another major gas-studies advance at HVO that was 

brought on by the rapid expansion of the environmentalscience field was adoption of the correlation spectrometer, 

or COSPEC, an instrument that became commercially 

available in the mid-1960s to measure the pollution released 

by power plants and other industrial smokestacks (Moffat 

and Millan, 1971). The remote capability of the technique 

was a remarkable aspect. The COSPEC, pointed at a 

gas-bearing plume, measures the proportion of scattered 

ultraviolet sunlight energy absorbed by SO2. The first volcanic 

COSPEC measurements were obtained in Japan at Mount 

Mihara in 1971 (Okita, 1971). Richard Stoiber and research 

colleagues from Dartmouth College began experimenting 

with the COSPEC at Kīlauea in 1975 (Stoiber and Malone, 

1975). During these early trials, they demonstrated that 

the SO2 emission rate could be measured by driving the 

vehicle-mounted, upward-looking instrument along the road 

beneath Kīlauea’s passive summit plume. The product of the 

concentration-pathlength cross section, recorded on a paper 

strip chart recorder, and the wind velocity yielded the SO2

Figure 5. W.F. Giggenbach’s volcanic-gas-sampling technique, 

illustrated here, relied on mechanically durable, chemically resistant 

titanium tubing to bring magmatic gas to an evacuated sampling bottle 

partially filled with NaOH solution. The inverted bottle is opened, allowing 

gas to bubble through solution. Acidic gases (SO2, H2S, CO2, HCl, HF) 

and H2O dissolve in the NaOH solution, and insoluble gases (H2, He, 

CO, N2, O2, Ar) pass through, and accumulate in, the headspace. 

Plug-in connection

Titanium tubes Sampling

bottle

with NaOH

solution

Butyl rubber

sleeves
300 Characteristics of Hawaiian Volcanoes

emission rate. Although the technique was novel, conversion 

of the chart-recorder trace to a concentration cross-sectional 

area and calculation of an SO2 emission rate were laborious 

multihour tasks.

The measurements by Stoiber and follow-on colleagues 

indicated that Kīlauea, even when not actively erupting, 

emitted as much as several hundred tons of SO2 each day—an 

amount equivalent to that of a small coal-fired powerplant 

(Stoiber and Malone, 1975). The COSPEC technique showed 

enough promise that regular measurement of SO2 emission 

rates became part of HVO’s routine monitoring by 1979 

(fig. 6). A favorable attribute of routine, campaign-style 

COSPEC measurements was building, for the first time, 

a volatile-emission time series that could be compared to 

2,000

1,500

1,000

500

0

1979 1984 1989 1994 1999 2004 2009 2013

SO2 emission rate, in metric tons per day

contemporaneous geophysical data. One of the first such 

investigations, for example, examined the relation between 

Kīlauea’s daily SO2 emission rate, caldera seismicity, and 

fortnightly Earth tides (Conner and others, 1988).

HVO Commits Full-Time Staff to Volcanic Gas 

Studies 

In 1978, Tom Casadevall joined the HVO research staff, 

inaugurating a permanent commitment by HVO to volcanic 

gas studies, with laboratory technical expertise from Bruce 

Furukawa, a University of Hawai‘i chemistry department 

graduate with a strong electronics background. In addition 

Figure 6. SO2 emission rates from Kīlauea’s summit caldera became part of the regular monitoring program at HVO in 1979. Measurements 

made by vehicle-based COSPEC revealed fitful degassing during runup to 1983 East Rift Zone eruption. The first 10 years of that eruption saw 

increased throughput of summit magma-storage complex and resulted in a general doubling of emissions before declining somewhat. The next 

large increase began in late in 2007 and continued through onset of 2008 summit eruption. Although this eruption began officially in March 2008, 

increased emissions had already forced long-term closure of the downwind part of Crater Rim Drive in the previous month. In this plot, each data 

point represents an average of several transects of the plume. Error bars are omitted for clarity, but uncertainty is generally within 30 percent of 

the measured value through 2007. Values shown beginning in early 2008 (shaded region) represent minimum constraint, possibly underestimating 

actual emission rates by factor of 2 to 5 (Elias and Sutton, 2012). 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 301

to establishing routine SO2 emission-rate measurements, 

Casadevall, along with Rick Hazlett of the University of 

Southern California, conducted the first detailed inventory 

of significant thermal features on Kīlauea and Mauna Loa 

volcanoes (Casadevall and Hazlett, 1983). Paul Greenland, 

an analytical geochemist who had worked for many years 

at the U.S. Geological Survey’s National Center in Reston, 

Va., joined HVO in 1980. Greenland immediately set to work 

building the first gas-analysis capability at the Observatory 

and began systematically characterizing the volatile species 

chemistry of Kīlauea’s many steam vents and solfataras.

In addition to mapping thermal features and initiating the 

use of the COSPEC on Kīlauea and Mauna Loa, Casadevall 

and Furukawa helped advance a new type of volcanic gas 

study: continuous, in situ, sensor-based monitoring. The 

electrochemical gas sensors used in this pursuit were another 

technology that HVO and its adjuncts adapted from the 

environmental movement. These sensors produce an electrical 

output proportional to the abundance of a specific gas species 

present in their immediate environment. A harsh volcanic setting 

like Kīlauea, however, presents analytical challenges to most 

commercially available sensors. 

Motoaki Sato, of the USGS in Reston, Va., had pioneered the adaptation of in situ chemical-sensing methodology on Kīlauea beginning in 1966, when he and Thomas 

Wright made the first field measurements of oxygen fugacity (f

O2

) on the surface of a newly emplaced lava lake at 

Makaopuhi Crater (Sato and Wright, 1966). They chose to 

study f

O2

 because of the importance of this intrinsic parameter to magmatic differentiation (Sato and Wright, 1966; 

Sato, 1978). Other gas-forming elements, such as carbon, 

hydrogen, and sulfur, react with oxygen and with mineral 

assemblages, changing the relative oxidation state of the 

melt during these interactions and during exsolution. Sato 

and Wright (1966) insisted that a better understanding of the 

amounts and types of these species could reasonably lead to 

a more thorough comprehension of magmatic differentiation 

as a part of the eruptive process.

One specific outcome of Sato and Wright’s Kīlauea f

O2 study was their recognition of the importance of molecular 

hydrogen (H2) in influencing the oxidation state of basaltic 

magma. Accordingly, Sato went on to develop a robust, field 

deployable H2 and reducing-gas sensor based on fuel-cell 

technology (Sato and McGee, 1981; Sato and others, 1986). 

A network of gas-monitoring sites with H2/reducing-gas 

sensors was established on Kīlauea, and a single site was 

installed on Mauna Loa. The Kīlauea network used a hybrid 

communications technology to make data available rapidly. 

An analog radio-telemetry system sent signals back to HVO 

for real-time display, and, simultaneously, the data were 

transmitted by way of geostationary satellite to Wallops 

Island, Va., where Sato’s group could access the data in nearreal time (McGee and others, 1987).

Casadevall’s, Greenland’s, and Furukawa’s initiation of 

regular SO2-emission-rate measurements and their pursuit 

of detailed fumarole gas chemistry greatly improved HVO’s 

capability to ask and answer fundamental questions related 

to gas release, further enhancing understanding of volcanic 

processes. In addition, the network of continuous, near-realtime gas sensors honed the emission-rate and chemical data 

to a fine temporal edge.

Kīlauea’s Renewed Restlessness Tests Volcanic 

Gas Studies’ Mettle While Researchers Synthesize 

a Volatile-Release Model

The growth and alignment of volcanic-gas-studies 

capabilities at HVO came just in time for an eruptive 

reawakening at Kīlauea. A near-surface intrusion on Kīlauea’s 

Southwest Rift Zone in 1980 and two one-day summit 

eruptions in 1982 helped stimulate investigation of the 

chemistry of directly sampled eruptive gases in the newly 

configured HVO laboratory (Greenland, 1987a). Increases in 

SO2 emission rates were measured by COSPEC during a 1979 

upper East Rift Zone eruption, a 1980 upper East Rift Zone 

intrusion, and the April 1982 summit eruption (fig. 6). These 

events were all short lived and occurred with little warning, 

further demonstrating the importance of a permanent volcanicgas-studies presence at HVO. 

This point was underscored in late 1982 when the continuous gas-monitoring network established by Sato’s group 

captured a widespread gas-release sequence that began at 

Kīlauea’s summit in the last days of December and, over the 

course of 3 days, propagated with the eruptive dike nearly 

30 km down the East Rift Zone before the January 3 onset of 

what became known as the Pu‘u ‘Ō‘ō-Kupaianaha eruption 

(fig. 7; McGee and others, 1987). Continuous monitoring of 

helium at Sulphur Banks showed a simultaneous decrease in 

gas release during this interval (Friedman and Reimer, 1987). 

Friedman and Reimer, however, did not suggest a cause for 

the decrease, which could reasonably have been associated 

with a pressure decrease and closure of caldera ring faults 

as magma left the summit region to feed the impending East 

Rift Zone eruption.

The early years of the Pu‘u ‘Ō‘ō-Kupaianaha eruption 

provided numerous opportunities for HVO and adjunct 

researchers to use the new analytical methods to expand 

and test models for how Kīlauea and similar volcanoes 

work. Initial gas samples from the East Rift Zone eruption 

showed a uniformly CO2-depleted condition that Greenland 

(1984) attributed to preeruptive magma storage in a 

shallow reservoir. Using the detailed chemistry of vent-gas 

samples and the first airborne CO2 and SO2 emission-rate 

measurements at Kīlauea’s summit and East Rift Zone, he 

refined this interpretation specifically to apply to magma 

storage beneath the volcano’s summit (Greenland and others, 

1985). 

Terry Gerlach, who joined the USGS Volcano Hazards 

Program in 1991, was also keenly interested in Kīlauea’s 

eruptive gases. At the onset of the Pu‘u ‘Ō‘ō-Kupaianaha 
302 Characteristics of Hawaiian Volcanoes

Pu‘u ‘Ō‘ō

East Rift Zone

Kīlauea

Caldera

Southwest Rift Zone

Mauna Loa

Pacific Ocean

Halema‘uma‘u

west

1971 fissure

Mauna Ulu

Pu‘u Kiai

1982 1983

354 360 001 007

Julian date, in UTC

Relative gas concentration

0 5 10 MILES

0 5 10 15 20 KILOMETERS

155°30' 155°

19°30'

155°20' 155°10' 154°50'

19°20'

19°10'

19°50'

19°40'

Figure 7. Continuous monitoring at Kīlauea in late December 1982 captured preeruptive gas release that began in summit area and spread 

down East Rift Zone. In all, 6 reducing-gas sensors at 5 sites (4 plots shown here), stretching 30 km across volcano, recorded event over course 

of 3 days and relayed collected data in near-real time by geostationary satellite. Pu‘u ‘Ō‘ō eruption began January 3, 1983 (red line). Adapted 

from McGee and others, 1987.

eruption, Gerlach, while still at Sandia National Laboratories, had become the latest researcher to reinterpret the 

famous Jaggar gas collections of 1917–19 from a thermodynamic perspective, identifying contamination by meteoric 

water as a principal cause of their apparent disequilibrium 

(table 1; Gerlach, 1980).

When episodic high lava fountaining began at the East 

Rift Zone in 1983, Gerlach and colleague Ed Graeber, in 

consultation with HVO researchers, sampled the eruptive 

vents and analyzed the gases as Greenland had done. In their 

landmark paper in the journal Nature, Gerlach and Graeber 

(1985) began by considering the parental volatiles trapped 

in olivine crystals deep beneath Kīlauea; then, progressively, 

they followed the degassing process by examining the 

volatile content of samples quenched at the somewhat-lower 

pressures of the submarine seafloor. Finally, they studied the 

volatiles remaining in subaerial eruptive spatter degassed to 

atmospheric pressure. These data, along with their careful 

analyses of gases collected close to where Pu‘u ‘Ō‘ō formed, 

and in HVO’s SO2 emission- and eruption-rate records, 

helped them construct a comprehensive volatile budget for 

Kīlauea Volcano spanning 27 years of intermittent activity 

(tables 2, 3; Gerlach and Graeber, 1985).

For their model of gas release at Kīlauea, Gerlach and 

Graeber (1985) identified two principle eruptive-gas types 

representing compositional end members. For a concerted, 

single-stage summit eruption, they proposed that as magma 

ascends directly to the surface, it erupts and degasses in a 

single stage and releases a CO2-rich “Type I” gas (fig. 8A). 

This eruptive configuration prevailed throughout Kīlauea’s 

continuous 19th-century Halema‘uma‘u lava-lake activity, 

which lasted until 1924. This interval also produced Jaggar 

and Shepherd’s classic gas collections of 1917–19 that had 

been reinterpreted by Gerlach (1980). The other compositional end member—“Type II” gas—was said to have characterized rift-zone eruptions. In this eruptive scenario, magma 

rises from the mantle and preeruptively degasses at 1- to 

6-km-depth from the summit magma reservoir, releasing 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 303

Table 1. Gerlach’s thermodynamic restoration of Jaggar and Shepherd’s best 1917–19 gas samples from Halema‘uma‘u lava lake testify to the care 

taken in the sampling by Jaggar, and the analysis of the gases by Shepherd’s lab. 

[Quality rating by Jaggar: E=excellent, G=good. Samples identified with an “R” designation denote Gerlach’s restoration (Gerlach, 1980). Both Gerlach’s 

restorations and the original reported analyses are expressed in mole percent. Gerlach’s restoration procedure provided the equilibrium temperature, f

O2

, and H2S 

concentration for each sample]

Sample Date Quality CO CO2 H2 H2O SO2 S2 H2S HCI T (°C) Log fO2 C/S Atomic H/C

J-8 3/25/1918 E 1.5 48.91 0.49 37.11 11.87 0.04 ND 0.08 4.12

J-8R 3/25/1918 E 1.51 48.9 0.49 37.09 11.84 0.02 0.04 0.08 1,170 –8.38 4.13 1.49

J-11 3/13/1919 E 0.62 21.89 0.33 64.38 12.52 0.26 ND 0 1.75

J-11R 3/13/1919 E 1.03 36.69 0.55 40.14 21.06 0.25 0.2 0 1,100 –9.33 1.74 2.17

J-13 3/15/1919 E 0.6 17.54 0.99 69.84 10.73 0.09 ND 0.21 1.63

J-13R 3/15/1919 E 0.62 17.82 1.01 69.29 10.93 0.03 0.08 0.21 1,175 –8.40 1.63 7.64

J-14 3/16/1919 E 0.48 15.26 0.18 79.16 4.82 0.1 ND 0 3.17

J-14R 3/16/1919 E 1.52 47.41 0.54 35.09 15.06 0.17 0.15 0 1,100 –9.30 3.15 1.46

J-16 3/17/1919 E 0.58 18.61 0.69 68.38 11.42 0.15 ND 0.17 1.63

J-16R 3/17/1919 E 0.74 23.21 0.87 60.42 14.31 0.07 0.14 0.21 1,140 –8.84 1.62 5.14

J-17 3/17/1919 G 0.37 11.76 0.59 80.36 6.57 0.24 ND 0.1 1.79

J-17R 3/17/1919 G 0.62 20.27 1.02 65.95 11.44 0.16 0.32 0.17 1,085 –9.65 1.77 6.45

Table 2. Gerlach and Graeber’s (1985) volatile budget factored the sequential degassing pathway of parental melt through summit reservoir 

equilibration and rift zone storage and eruption by considering gas release behavior during a 27-year period of activity (July 1956 to April 1983).

Volatile class H2O CO2 S Cl F

Parental volatiles 18(2)1,2 18(2)3 1004 38(5)1 39(5)3 1004 7.6(7)1 7.7(7)3 1004 0.51(5)1 0.52(5)3 1004 2.1(2)1 2.1(2)3 1004

Chamber gas 2(3) 2(3) 10 36(5) 37(5) 95 3.5(9) 3.6(9) 46 0.00(7) 0.00(7) ~0 0.0(3) 0.0(3) ~0

Stored volatiles5 16(2) 16(2) 90 2.0(2) 2.0(3) 5 4.1(5) 4.2(5) 54 0.51(5) 0.52(5) 100 2.1(2) 2.1(2) 100

Volcanic gas 3.5(7) 3.6(7) 20 0.4(1) 0.4(1) 1 1.1(2) 1.1(2) 15 0.02(2) 0.02(2) 3 0.01(9) 0.01(9) ~0.4

Residual 2.0(3) 2.1(3) 12 0.31(5) 0.31(5) 1 0.31(3) 0.31(3) 4 0.16(2) 0.17(2) 32 0.72(8) 0.73(8) 34.6

Non-erupted 10(1) 10(1) 58 1.3(2) 1.3(2) 3 2.7(3) 2.7(3) 35 0.33(3) 0.34(3) 65 1.3(1) 1.4(1) 65

1

Total g × 10–12 calculated for the period July 1956 to April 1983 and rounded to significant figures.

2

Parentheses enclose propagated error in preceding digit.

3

(g per day) × 10-8 calculated from total g × 10–12 per the 9,800-day period and rounded to significant figures.

4

Each category (before rounding) as percent of parental volatile supply.

5

Assumes equilibration of CO2 in reservoir melt at 2-km depth.

Volatile class H2O CO2 S Cl F

Parental1 0.30(2)4 0.65(6) 0.130(4) 0.0087(3)2 0.0354(20)5

Stored2 0.27(2) 0.034(3) 0.070(5) 0.0087(3)2 0.0354(20)

Residual3 0.10(1) 0.015(2) 0.015(1) 0.0080(3)2 0.0350(20)

Table 3. Gas composition of parental, stored, and residual volatiles based on the volatile 

budget of Gerlach and Graeber (1985; table 2).

1

Weight percent concentrations for parental melt oversaturated in CO2 and chamber conditions. All 

concentrations calculated except for S, which is based on glass inclusion data.

2

Weight percent concentrations for reservoir-equilibrated melt at 2-km depth. All concentrations 

calculated except for S, which is based on data for glassy East Rift Zone submarine basalts.

3

Mean values for CO2, S, Cl, and F in weak fountain spatter. H2O value for fountain spatter from the 

1983 middle East Rift Zone eruption.

4

Parentheses enclose propagated error in preceding digit.

5

Based on stored volatile concentrations.
304 Characteristics of Hawaiian Volcanoes

Figure 8. Gerlach and Graeber’s (1985) volatile budget diagram 

(adapted) showing two end-member eruptive gas types. A, Single-stage 

degassing of parental magma, with a plume rich in CO2 and including SO2

and H2O (Type I volcanic gas), attends sustained summit eruptions. B, 

Prolonged rift-zone eruptions are associated with two-stage gas release, 

involving preeruptive venting of CO2-rich chamber gas at summit and SO2- 

and H2O-rich gas release (Type II volcanic gas) from rift-zone eruption site. 

most of its CO2, along with a small portion of its SO2, as 

“chamber gas”; the remaining SO2- and H2O-rich, Type II 

gas is released through the process of eruption when magma 

reaches the surface in the rift zone (fig. 8B). 

Gerlach rigorously examined the progressive exsolution 

of volatiles in relation to pressure and depth, using analyses 

of the Type II gases that were prevalent during the beginning 

of the Pu‘u ‘Ō‘ō-Kupaianaha eruption. CO2, the least soluble 

gas, exsolved early during magma ascent at depths of tens of 

kilometers, followed by SO2, H2O, HCl, and HF (Gerlach and 

Graeber, 1985; Gerlach, 1986). Working with his 1983 East Rift 

Zone gas collections, Gerlach (1993) concluded that coexisting 

lavas effectively buffer f

O2

 from molten temperatures down to 

several hundred degrees below subsolidus temperatures.

By 1983, when the Pu‘u ‘Ō‘ō-Kupaianaha eruption 

began, the campaign-style SO2-emission-rate-measurement 

program started by Casadevall in 1979 had become an 

integral component of eruption monitoring at HVO. Although 

Casadevall and Furukawa left HVO to join the newly formed 

Cascades Volcano Observatory (CVO) in Vancouver, Wash., 

in 1981, Casadevall remained an active adjunct of HVO for 

years to come. The regular COSPEC measurements that he 

initiated and that Barry Stokes took over in 1982 captured 

a rapid doubling of Kīlauea summit SO2 emissions after the 

start of the East Rift Zone eruption, and airborne SO2 and 

CO2 emission-rate measurements at the summit and East Rift 

Zone affirmed the previous observations by Paul Greenland 

and Terry Gerlach (Greenland and others, 1985; Gerlach 

and Graeber, 1985)—that during rift zone eruptions, Kīlauea 

magma loses most of its CO2 through passive degassing at 

the summit, while most of the SO2 is released in the rift zone 

eruptive plume (Casadevall and others, 1987). 

The March 25 to April 14, 1984, Mauna Loa eruption 

provided an opportunity to make the first airborne SO2 and CO2

measurements at that volcano (Casadevall and others, 1984). 

The initial SO2 emission rate, ~70 metric kilotons per day 

(kt/d), exceeded the dynamic range of the COSPEC during the 

first week of the eruption. These rates were measured by the 

Total Ozone Mapping Spectrometer (TOMS) carried aboard 

the NIMBUS 7 spacecraft until they dropped below its 10-kt/d 

practical detection limit. After this interval, the COSPEC, 

reconfigured to measure higher SO2-emission rates, recorded the 

decline in emissions until the eruption ended.

Greenland continued to explore the volatile-release 

processes associated with Hawaiian eruptions by combining 

the gas-chemistry analyses with COSPEC-derived SO2

emission rates and geophysical data. He reported the first 

selenium and tellurium abundances in eruptive gases 

(Greenland and Aruscavage, 1986) and noted the similarity 

in CO2 depletion between Kīlauea and Mauna Loa summit 

and rift volatile release during the 1984 Mauna Loa eruption 

(Greenland, 1987b). At Kīlauea, he used tilt and emissionrate data to estimate the dimensions and rise rate of magma 

in the eruptive conduit at Pu‘u ‘Ō‘ō, along with pressure/

density profiles for the eruptive column (Greenland and 

others, 1988). Greenland also reconciled and summarized 

75 years of eruptive gas study as part of HVO’s Diamond 

Jubilee (Greenland, 1987a). In his conclusions, he noted 

that Hawaiian eruptive gases exist fundamentally as either a 

preeruptively degassed, CO2-depleted form with a molecular 

C/S ratio of ~0.2, or, in the case of magma arriving directly 

from the mantle, a CO2-rich form with a molecular C/S 

ratio of ~2.0. He also showed that the disequilibrium 

compositions of eruptive-gas samples could be accounted 

for by assimilation of crustal or meteoric water. Finally, he 

concluded that the total volatile content of magma stored 

in the summit reservoir is uniform and less than 0.5 weight 

percent. Paul Greenland retired in 1986, and although his 

position was not immediately filled, Barry Stokes kept up the 

high-quality, long-running SO2 emission-rate and summitfumarole chemistry databases until he left the USGS in 1991.

Tuning Technologies and Enhancing the TimeSeries Database

In 1993, Jeff Sutton joined HVO as staff geochemist. 

Sutton had worked intermittently at HVO in his role as a 

chemical-sensors specialist on Moto Sato’s project, studying 

the geochemistry of gas-forming elements in the 1980s, and 

he had helped build CVO’s gas-analysis and continuous 

gas-monitoring capability while a staff member there. In 

Parental magma

(parental volatiles)

Lava lake

(residual volatiles)

Halema‘uma‘u

Reservoir-equilibrated magma

(stored volatiles)

East Rift Zone

Type I

volcanic

gas

CO2-rich plume

Chamber gas

Type II

volcanic

gas Lava flow

(residual

volatiles)

Reservoir-equilibrated magma

(stored volatiles)

Parental magma

(parental volatiles)

B

A
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 305

conjunction with his appointment, Tamar Elias became a 

member of the volcanic-gas-studies group almost a year 

earlier, after overlapping with Stokes. Elias, a chemist, had 

been working at HVO in various capacities for several years, 

as well as with the National Park Service and other agencies, 

to monitor air quality.

In addition to the continuation of the Pu‘u ‘Ō‘ōKupaianaha eruption, the early 1990s were a time of geothermal energy development on Kīlauea’s lower East Rift Zone. 

In support of an environmental impact assessment related 

to this development, Sutton and Elias reported the effects of 

volcanic emissions on ambient-air characteristics through a 

literature review. A series of campaign-style field measurements of ambient air and rainwater at 15 sites at Kīlauea’s 

summit and along the East Rift Zone to the coastal lava entry 

documented the mostly localized environmental effects (Sutton and Elias, 1993a).

Because most of the SO2 being emitted from Kīlauea 

was coming from Pu‘u ‘Ō‘ō and other scattered East Rift 

Zone sources, Elias began making routine vehicle-based 

SO2-emission-rate measurements along Chain of Craters 

Road in 1992 (fig. 9). Before these measurements 

commenced, HVO had relied on sporadic, tripod-based 

COSPEC SO2 data collections near Pu‘u ‘Ō‘ō to measure 

rift zone emissions (Elias and others, 1998). Fixed-wing 

airborne measurements, though satisfactory for episodes 

of high lava fountaining, were expensive, time consuming, 

and less acceptable for measuring the low-level, groundhugging plumes that characterized the eruption, beginning 

with the continuous effusive Kupaianaha phase in 1986. 

Also, a comparison between contemporaneous tripod- and 

vehicle-based East Rift Zone COSPEC measurements 

indicated that plume-source geometry and light-propagation 

issues contributed to tripod data that were consistently lower 

than vehicle-based emission-rate measurements, and that 

the vehicle-based measures were probably more accurate 

(Andres and others, 1989; Elias and others, 1998).

The closing years of the 1990s brought an additional 

technological advance in HVO gas studies—the field-portable 

LI-COR Biosciences CO2 analyzer. This direct-sampling 

1992 1997 2002 2007 2012

12,000

10,000

8,000

6,000

4,000

2,000

0

SO2 emission rate, in metric tons per day

Figure 9. Vehicle-based SO2 emission-rate measurements were shown to be an effective and economical means of tracking activity at Pu‘u ‘Ō‘ō 

and other East Rift Zone eruption sites. These emissions dwarfed those of the summit by nearly 10:1 until several months after start of 2008 summit 

eruption, when preeruptive degassing at summit caused the SO2 emission rate from Pu‘u ‘Ō‘ō to decrease. In this plot, each data point represents an

average of several transects through plume. Error bars are omitted for clarity, but uncertainty is generally within 30 percent of the measured value.
306 Characteristics of Hawaiian Volcanoes

instrument applied the same nondispersive-infrared-spectroscopic approach to measuring CO2

 that had been used with the 

Miran analyzer during the early days of the Mount St. Helens 

eruptions (Harris and others, 1981); however, improved sensitivity and updated digital signal processing made the LI-COR 

analyzer superior both for vehicle-based and airborne gas 

studies (Gerlach and others, 1997). Gerlach and coworkers used 

the LI-COR analyzer at Pu‘u ‘Ō‘ō in an airborne (fixed wing) 

mode to simultaneously measure both SO2 emission rate, by 

COSPEC, and SO2 concentration, by Fourier transform infrared 

(FTIR) spectroscopy (Gerlach and others, 1998; McGee and 

Gerlach, 1998). The method they developed to measure the CO2

emission rate at Pu‘u ‘Ō‘ō used the molecular C/S ratio in the 

core of the gas plume. The product of the SO2 emission rate, 

the molecular C/S ratio, and the molecular-weight ratio yielded 

the CO2 emission rate. 

The CO2 emission rate measured in the Pu‘u ‘Ō‘ō 

plume, 0.300 kt/d, supported earlier assertions by both 

Gerlach and Graeber (1985) and Greenland (1984) that 

the magma being erupted at Pu‘u ‘Ō‘ō had already been 

degassed of most of its CO2. The airborne plume molecular 

C/S ratio reported was similar to that measured in eruptivegas samples from Pu‘u ‘Ō‘ō fumaroles during the first 12 

years of the eruption, providing support for the new methodology (Gerlach and others, 1998).

Eruptive Changes Lead to a Better Appreciation of 

Volcanic Gas Emissions as a Volcanic Hazard

The spectacular high lava fountains that occurred from 

1983 to mid-1986 released SO2 from Pu‘u ‘Ō‘ō at rates 

ranging from 5 to 32 kt/d (Casadevall and others, 1987). 

Typically, however, these episodes were brief, lasting 24 

hours or less (Heliker and Mattox, 2003). During the nearly 

month-long intereruptive pauses, SO2 emissions commonly 

totaled less than 0.5 kt/d for the summit and East Rift Zone 

combined (Casadevall and others, 1987; Chartier and others, 

1988). For the first several years of East Rift Zone activity, these multiweek pauses resulted in greatly reduced SO2 

release, giving ample time for high-fountaining gas emissions to dissipate (Sutton and Elias, 1993b).

The continuous lava effusion and consequent gas release 

that produced the Kupaianaha lava shield ~3 km downrift of 

Pu‘u ‘Ō‘ō, beginning in 1986, resulted in an accumulation of 

volcanic gases downwind of Kīlauea’s summit and East Rift 

Zone sources, especially on the leeward (west) coast of the 

island (fig. 10). By 1987, island residents from the Ka‘ū and 

Kona Districts, especially, began reporting negative health 

symptoms that included breathing difficulties, headaches, 

eye irritation, and general flulike symptoms. The term “vog,” 

an Island of Hawai‘i-coined portmanteau for “volcanic 

smog,” caught on and eventually became used worldwide by 

communities impacted by nearby volcanoes. As the resident 

source of data and hazards information related to eruptive 

emissions, HVO was consulted about volcanic air pollution 

extensively by County, State, and Federal agencies, as well 

as health studies researchers. The conversion of SO2 to acidic 

sulfate aerosol during transport and the subsequent rainout 

of these acidic eruptive products into water-catchment 

tanks, especially on the island’s east coast, caused leaching 

of substantial amounts of lead and copper into domestic 

drinking-water supplies. Clinical tests confirmed high levels 

of lead in the blood of some downwind residents (Sutton 

and others, 1997). The discovery of this secondary volcanic 

hazard resulted in a countywide ban on the availability of 

lead-containing roofing materials and plumbing supplies 

(Sutton and Elias, 1993b). 

As the eruption continued, HVO became an active voice 

in a Vog Task Force convened by Hawai‘i County to understand 

and manage the growing impact of volcanic gas emissions on 

island life. HVO also participated in vog symposia held on the 

Island of Hawai‘i and on O‘ahu (Casadevall and others, 1991; 

Chuan, 1991; Gerlach and others, 1991; Sutton and Elias, 

1996, 1997). HVO gas geochemists began making frequent 

presentations to schools and community groups and served 

as a source of review for health studies professionals seeking 

to understand the nature of volcanic-emission sources and 

the physical aspects of Kīlauea’s plume dispersion. With its 

knowledge of and perspective on emission-source parameters, 

along with the physical science governing vog distribution, 

HVO staff produced a USGS Fact Sheet that helped answer 

continuing requests for information about the topic (Sutton and 

others, 1997). 

The 2000s: Examining Gas and 

Geophysical Data Together on a 

Common Time Base

CO2 and SO2 Emissions Reflect Magma-Supply 

and Eruption Rates 

The technological explosion of the 1960s through 

1990s, and the eruptive activity at Kīlauea and Mauna Loa, 

produced a developmental watershed, both of volcanic gas 

study methods and of reasoned conceptual models for active 

volcanic processes in Hawai‘i and elsewhere. The growing 

database of SO2 emission rates at HVO provided insights into 

Kīlauea’s plumbing system, especially the relation between 

gas emissions, seismicity, and lava effusion rates. The rapid 

jump in summit SO2 emission rates in 1983 at the beginning 

of the East Rift Zone eruption and the decline in summit SO2

release in 1991 during shutdown of the Kupaianaha vent 

affirmed the close magmatic connection between Kīlauea’s 

summit and East Rift Zone. Summit SO2 emission rates 

were also observed to correlate with shallow, short-period 

seismicity beneath the caldera but not with long-period 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 307

23°

22°

21°

20°

19°

18°

161° 160° 159° 158° 157° 156° 155° 154°

0.0

0.1

0.2

0.3

0.4

0.5

Aerosol optical density (unitless)

Figure 10. Image of the Hawaiian Islands from the Moderate Resolution Imaging Spectroradiometer (MODIS) carried aboard the National 

Aeronautics and Space Administration’s Aqua and Terra polar-orbiting satellites, showing Kīlauea’s volcanic-aerosol plume, which results from 

primary particles emitted directly from Kīlauea’s vents, and secondary particles formed by chemical conversion of SO2 to acidic sulfate aerosol. 

Under prevailing northeasterly trade-wind conditions, these emissions are carried from Kīlauea’s eruptive vents southward, where wind patterns 

wrap around southern tip of island and send emissions northward along leeward coast. Warmer colors (yellow and orange) depict increasing 

aerosol loading as SO2 is gradually converted to submicron-size droplets of sulfuric acid and particles of acidic and neutralized sulfates. 

seismic events, as had been demonstrated at more silicic 

systems (Sutton and others, 2001).

The relation between East Rift Zone SO2 release and lava 

effusion reported by Sutton and others (2001) was subsequently 

used to derive a continuous lava-effusion rate record for the 

first 20 years of the Pu‘u ‘Ō‘ō-Kupaianaha eruption (Sutton 

and others, 2003). The long-running record of eruptiveSO2

-derived lava effusion rates were compared with direct 

geologic observations and with effusion rates determined by 

very-low-frequency (VLF) electromagnetic profiling of lava 

transport through master lava tubes. The geochemical and 

geophysical methods produced total volume estimates for the 

eruption that agreed with one another within 10 percent. Such 

a good agreement increased confidence in both of these remote 

techniques for effusion-rate monitoring and, simultaneously, 

provided an additional tool for obtaining this fundamental 

volcano-monitoring metric. East Rift Zone SO2 emission rates 

were especially important for derivation of lava effusion rates 

during 2003–08, when changes on the lava flow field resulted in 

the interruption of routine VLF measurements due to a lack of 

long-term accessible lava tubes with skylights (fig. 11).

By the late 1990s, it was also becoming possible to measure CO2

 emission rates on a regular basis, thanks to Gerlach 

and others (1998, 2002), who adapted the airborne molecular 

C/S ratio technique of CO2 emission-rate measurement that they 

had developed earlier at Pu‘u ‘Ō‘ō for vehicle-based traverses. 

At Kīlauea’s summit, they reported an average CO2 emission 

rate of 8.5±0.3 kt/d between 1995 and 1999. This value, several 

times higher than what Greenland and others (1985) reported 

earlier, possibly resulted because the earlier measurements 
308 Characteristics of Hawaiian Volcanoes

relied on cross-sectional profiling of a ground-hugging plume 

that was difficult to measure in an airborne mode. Using the 

estimated magma supply rate of Cayol and others (2000), 

Gerlach and others (2002) revised upwards the estimated CO2

content of the parental magma (see table 2), from 0.65 to 0.7 

weight percent, and they suggested that CO2 emission-rate 

measurements could routinely serve as an effective proxy for 

magma supply monitoring (a proposal that proved important; 

see section below entitled “A Case Study: The 2003–12 Eruptive Sequence”). Accordingly, HVO adapted the vehicle-based 

method of Gerlach and others (2002) and made it operationally 

practical by replacing the cumbersome FTIR spectrometer that 

had initially been used to measure ambient SO2 concentration 

with a similarly sensitive, but less complex, electrochemical 

instrument. Using this revised measurement strategy, HVO continued to update the CO2

 emission-rate database (fig. 12).

Gas emission-rate studies were evolving on other fronts, 

as well. Although analog emission-rate data acquisition and 

reduction had been converted to digital by the early 2000s, 

the growing availability of low-cost, miniaturized ultraviolet 

(UV) spectrometers provided an attractive alternative to the 

expensive and cumbersome COSPEC. In collaboration with Lava effusion rate, in cubic meters per day

1.5×107

1.5×106

1.5×105

1992 1996 2000 2004 2008

Figure 11. Kīlauea East Rift Zone lava effusion rates over time, integrated and computed from a combination of SO2 emission rates and 

flow mapping, have been shown to agree to within 10 percent of lava effusion rates estimated by flow mapping and very low frequency (VLF) 

electromagnetic measurements (Sutton and others, 2003). When absence of a master lava tube made VLF measurements impractical during 

2003–08, HVO relied upon SO2-derived effusion rates. These gas-based effusion rates captured refilling of Pu‘u ‘Ō‘ō in 1997 for the first time in 

more than 10 years. Similarly, eruptive surges in 2002 and 2005 were tracked with this remote technique. Error bars are omitted for clarity, but 

uncertainty is generally within 40 percent of the measured value.

colleagues Keith Horton and Harold Garbeil of the University 

of Hawai‘i, Mānoa, HVO tested and implemented a version 

of this technology designed and built by Horton and Garbeil 

and dubbed the FLYSPEC, owing to its small size (Elias and 

others, 2006; Horton and others, 2006). The FLYSPEC used 

the correlation spectrometry approach of standardizing field 

measurements in-situ by calibrating experimental data with 

known SO2 concentration cells measured in the field (fig. 13). 

Data acquisition and reduction software automatically merged 

SO2 concentration-pathlength data with time-synchronized 

Global Positioning System (GPS) data, and readily 

incorporated the LI-COR CO2 measurement parameters and 

the electrochemical SO2 and H2S concentration data needed to 

calculate CO2 and H2S emission rates.

Fourier Transform Infrared Spectroscopy Links 

Short-Term Eruptive Processes to Individual Gas 

Species 

A further advance in gas studies at HVO came in 2004 

with the arrival of Mendenhall Postdoctoral Fellow Marie 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 309

CO2 emission rate, in metric tons per day

40,000

30,000

20,000

10,000

0

1995 1999 2003 2007 2011 2013

Figure 12. Kīlauea summit CO2 emission rates over time are related to long-term changes in magma supply to volcano. These rates, steady 

from 1995 to 2003, increased drastically in 2004, signaling a magma-supply surge that peaked in 2005–06 and then began a steady decline for 

the next several years as magma supply to the volcano waned. Error bars represent the standard deviation of several transects through plume. 

Values shown beginning in early 2008 (shaded region) are underestimates due to complications in gas emission rate measurements associated 

with summit eruptive activity (Elias and Sutton, 2012).

Figure 13. Miniaturization of ultraviolet spectrometers resulted in 

improved SO2 emission-rate measurements. With colleagues Horton and 

Garbeil of the University of Hawai‘i, HVO codeveloped the FLYSPEC 

(left). For six months, beginning in late 2003, FLYSPEC measurements 

were made in parallel with the COSPEC (right) before COSPEC use was 

discontinued. 

Edmonds, who characterized the summit and rift zone emissions of HCl, HF, SO2

, CO2, CO, and other gases using openpath Fourier transform infrared (OP-FTIR) spectroscopy. In 

this field measurement technique, infrared (IR) energy from a 

hot lava source, an IR lamp, or the sun is measured as it passes 

through, and is absorbed by, the volcanic plume along an open 

atmospheric path. Individual gas species, including SO2, CO2, 

H2O, CO, HCl, HF, and COS, absorb light energy and can be 

quantified within different wavelength regions of the spectrum. 

An example of the utility of the OP-FTIR approach at Kīlauea 

is provided by Edmonds and Gerlach (2006), who studied the 

composition of the coastal-entry plume and determined that 

intense HCl generation could occur at moderate-size lava ocean 

entries, creating a significant local hazard.

Edmonds and Gerlach (2007) also examined short-term 

variations in gas emissions at eruptive vents within Pu‘u ‘Ō‘ō 

crater using OP-FTIR spectroscopy in 2004 and 2005. In their 

study, they characterized three distinct types of gas release: 

(1) persistent, continuous release of H2O-rich, CO2-poor gas 

that occurred when magma ascended with bubbles that burst 

at the surface; (2) “gas-piston” release, characterized by loud 

jetting and increased glow within small eruptive vents (but 

probably different from the more traditional gas piston events 

at Kīlauea later described by Patrick and others, 2011a) and 

(DOGSPEC)

FLYSPEC COSPEC
310 Characteristics of Hawaiian Volcanoes

accompanied by the release of CO2-rich gas, interpreted to 

be driven by gas slugs rising from depths of several hundred 

meters; and (3) gas release associated with lava spattering, 

due to quiescent coalescence of large, H2O-rich and SO2- 

and CO2-poor bubbles. In a follow-up study, Edmonds and 

others (2009) used the 2004–05 Pu‘u ‘Ō‘ō data to develop an 

exsolution model for halogen gas release, recognizing that 

this process occurs principally at very shallow depths (tens of 

meters or less) in the magma column. Using synchronous SO2

concentration and emission rates, they reported HCl and HF 

emissions from Pu‘u ‘Ō‘ō of 25 and 12 t/d, respectively.

Regular OP-FTIR-spectroscopic measurements quickly 

became an important part of the eruption-monitoring strategy at 

HVO. As a remote technique, OP-FTIR spectroscopy allowed 

HVO staff to make more detailed measurements of gas chemistry with more flexibility and safety than ever before. Although 

routine gas-bottle collections with subsequent laboratory GC 

analysis were still conducted at long-lived summit fumaroles 

(fig. 14), regular collection of uncontaminated samples at active 

eruptive vents was difficult. In combination with SO2 emissionrate measurements, OP-FTIR-derived gas concentrations could 

be compared in ratio fashion with SO2 concentrations in the 

volatile matrix and used to estimate emission rates of principal 

gases (H2O, CO2, HCl, HF, CO, COS) emitted at vents that were 

impractical to sample with gas bottles—for example, in lava 

lakes at both Kīlauea’s summit and Pu‘u ‘Ō‘ō.

The growing database of different types of volcanic 

gas measurements meant that HVO could study eruptive 

activity with unprecedented depth. Regular CO2 emission 

rates, tied to magma supply, and East Rift Zone SO2 emission 

rates, indicating lava effusion rate, were linked analytically 

via OP-FTIR spectroscopy with other principal volcanic 

gases, providing a view encompassing a wide range of gas 

solubilities and associated subsurface conditions. Such 

applications also facilitated comparison of gas data with 

other geologic and geophysical monitoring parameters. 

For example, the World Wide Web-based Volcano Analysis 

and Visualization Environment (VALVE) permitted rapid, 

simultaneous display, on a common time base, of many 

types of data collected by HVO (Cervelli and others, 2002). 

With the new tool, plotting of time-series gas emissions or 

continuous monitoring of concentration data, for example, 

Figure 14. Easy access to Kīlauea’s summit fumaroles and their close proximity to HVO’s analytical laboratory have enabled fumarole sampling 

as a part of routine gas monitoring. Inset, HVO geochemist Tamar Elias samples gases from a boiling-point solfatara using an evacuated bottle 

(the same technique has been in use for over for more than 30 years). Molecular C/S ratio (upper plot) and temperature (lower plot) over time for 

fumarole sampling sites on southwest (B-site) and north (C-site) rim of Halema‘uma‘u Crater. Emissions from these sites are sensitive to major 

changes in conditions in summit magmatic system, as well as to sulfur-gas scrubbing caused by seasonal rainfall events. Molecular C/S ratio Temperature, in degrees Celsius

600

200

400

105

100

95

90

85

80

1997 1999 2001 2003 2005 2007 2009 2011

25

20

15

10

C-site

B-site
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 311

against seismicity or surface deformation became possible. 

This new capability helped set the stage for interactive and 

ongoing interpretation of volatile emissions in the broader 

context of Kīlauea’s ongoing eruptive activity.

A Case Study: The 2003–12 Eruptive 

Sequence 

From 2000 through mid-2003, CO2 emission rates at 

Kīlauea’s summit approximately followed the long-term average of 8.5 kt/d, indicating an overall steady magma supply rate 

to the volcano. During the same interval, the absence of significant change along a caldera-crossing GPS baseline indicated 

little summit deformation and an eruption rate that was keeping 

up with, or slightly outpacing, magma supply (Cervelli and 

Miklius, 2003; Poland and others, 2012). Relatively consistent 

East Rift Zone SO2 emissions suggested a steady lava effusion 

rate, consistent with the inference of a constant magma supply. 

In late 2003, however, conditions began to change. The 

summit-crossing GPS baseline began to increase in length, and 

uplift began to occur within the summit caldera (fig. 15A). By 

mid-2004, the summit CO2 emission rate had increased by ~250 

percent (fig. 15B), indicating at least a doubling of the magma 

supply rate (Poland and others, 2012). At about the same time, 

an increase in the East Rift Zone SO2 emission rate (fig. 15C) 

indicated a near-doubling of the eruption rate as well. Other 

gas-studies measurements corroborated the increase in magma 

supply. Continuous CO2 concentration monitoring within the 

North Pit instrument vault, located near the north edge of 

Halema‘uma‘u Crater, began to show a marked rise in CO2

concentration at the beginning of 2004. The rise in subsurface 

CO2 to hazardous levels in the summit caldera caused Hawai‘i 

Volcanoes National Park to close caldera lava-tube caves to 

researchers and visitors alike in 2005.

Summit uplift continued into 2006 and began to extend 

outward into Kīlauea’s Southwest Rift Zone. With the supply of 

magma to the summit outpacing the East Rift Zone’s capability 

to erupt lava at the increased volume rate, the upper East Rift 

Zone plumbing failed in June 2007. This failure manifested 

as a dike that intruded the upper East Rift Zone, resulting in 

a small eruption on June 19. The Father’s Day eruption, as it 

Figure 15. Cross-caldera extension (A, measured between GPS sites that span caldera [red dots in the inset map]) and CO2 emission rate and 

ambient molecular C/S ratio (B) captured an increase in magma supply to Kīlauea beginning in 2003 that resulted in an increase in East Rift Zone 

effusion rate, reflected by an increase in East Rift Zone SO2 emission rates by 2005 (C). Although CO2 emission rates began to decline as early 

as 2006, Kīlauea’s summit continued to expand, with supply outpacing effusion until 2007, when an upper East Rift Zone intrusion and eruption in 

June relieved magmatic pressure somewhat. Magma withdrawal likely lowered summit-reservoir pressure, causing an increase in SO2 emission 

rate from the summit magma reservoir and a consequent decrease in molecular C/S as magma worked its way toward the surface in early 2008.

2000 2001 2002 2003 2004 2005 2006 2007 2008 2009

SO2 emission rate,

in metric tons per day

CO2 emission rate,

in metric tons per day

Distance change,

in meters

0

2,000

4,000

0

10,000

20,000

0

0.2

0.4

30,000

0.6

Molecular C/S ratio

0

100

200

300

400

Summit CO2 emission rate

Summit molecular C/S ratio

Summit

East Rift Zone

A

B

C
312 Characteristics of Hawaiian Volcanoes

became known, was volumetrically small—approximately 

1,500 m3 (Poland and others, 2008)—but the emplaced dike 

exceeded 15×106 m3 in volume (Montgomery-Brown and 

others, 2010). Gas measurements at the summit during the 

activity paradoxically recorded a spike in the SO2 emission rate, 

despite the fact that magma was clearly draining from beneath 

the summit caldera, as indicated by deflationary deformation. 

This spike was interpreted as an indicator of decompression of 

the summit reservoir due to magma withdrawal (Poland and 

others, 2009). 

The summit SO2 emission rate after the Father’s Day 

eruption became steady once more, though somewhat 

elevated, relative to the pre-Father’s Day eruption period. 

The summit CO2 emission rate, however, continued the 

linear decline that had begun in 2006, and the corresponding 

molecular C/S ratio in the summit gas plume dropped incrementally for the first time in more than 3 years (fig. 15B). 

As 2007 came to a close, summit caldera seismicity began to 

increase, and, after a large rainfall event in early December 

that effectively scrubbed summit SO2, those emissions began 

to climb, as well.

At the beginning of 2008, several gas-release effects were 

observed almost simultaneously. The summit SO2 emission rate 

began to increase more sharply and, by late January, intermittent 

signal saturation indicated that FLYSPEC retrievals had begun 

to underestimate actual SO2 emission rates (Elias and Sutton, 

2012). Even as underestimates, these values were the highest 

measured in more than 25 years (fig. 6). Meanwhile, the summit 

CO2 emission rate continued its monotonic decline. Analysis of 

gases from the two long-monitored summit fumaroles indicated 

that, although the CO2 concentrations in these solfatara had 

remained fairly steady during the previous year, the SO2

concentrations had risen significantly. The net effect was that by 

mid-January 2008, gases being emitted from these fumaroles 

had taken on a Gerlach and Graeber (1985) CO2-depleted, SO2- 

and H2O-rich, Type II eruptive gas composition (fig. 14). 

The composite gas signature of rising summit SO2

emissions and an eruptive molecular C/S-ratio signature in 

fumarole gases, along with a declining CO2 emission rate, 

indicated a shallow process at work, possibly even a shallowing 

of magma. Summit tremor also increased with the rising SO2

emission rate but, enigmatically, the overall summit deformation 

signal was one of deflation (fig. 15A). By mid-February, 

the summit SO2 emission rate was high enough to produce 

hazardous SO2 concentrations adjacent to Halema‘uma‘u Crater. 

Hawai‘i Volcanoes National Park personnel were responding to 

an increasing number of respiratory emergencies in the vicinity, 

and in late February, in consultation with HVO, they closed the 

area adjacent to, and downwind of, Halema‘uma‘u. 

The appearance of a new, vigorously fuming area beneath 

the Halema‘uma‘u visitor overlook was noted on March 12, 

2008, which, a week later, was followed by the explosive opening of a new eruptive vent at Kīlauea’s summit—the first summit eruption since 1982 (fig. 16; Houghton and others, 2011; 

Patrick and others, 2011b). With continued widening over time, 

a lava lake deep within the growing vent became visible from 

the crater rim, providing an opportunity to study eruptive gas 

chemistry directly, using OP-FTIR spectroscopy (fig. 17). The 

composition of emitted gases early in the eruption, similar to 

that of the East Rift Zone eruptive gases reported by Greenland 

(1986, 1987b), also fit the description of a Gerlach and Graeber 

(1985), CO2-depleted, SO2- and H2O-rich, Type II eruptive gas 

(Sutton and others, 2009). 

Discussion and Conclusions

The first hundred years at HVO began with academicstyle volcanic gas studies probing the chemical properties 

of volcanic exhalations. In 1912, the fundamental question 

about volcanic gas being asked, not only in Hawai‘i but at 

volcanoes worldwide, was “What’s in it?” Not surprisingly, 

this basic question was not so easy to answer a century ago. 

Sampling strategies were a speculative venture at that time; 

there was no procedure for sampling gases, because knowledge of volcanic gas content was lacking. Paradoxically, the 

reason why gas composition was unknown was that hazardous, hot, and corrosive field conditions made gas sampling 

more challenging than anything ever tried, even in industry. 

Once the hard-won 1912–24-era samples—as rare and precious as Moon rocks—were collected in fragile glass vials 

and transported by ship back to the Geophysical Laboratory 

in Washington, D.C., the laboratory analyses themselves 

had to be carried out at a higher level of sensitivity than had 

been attempted previously. First, however, analysts needed 

to examine the samples qualitatively to determine which 

chemical species were present before they could attempt to 

measure the amounts. Until these steps were taken, investigators could not begin to consider the volcanological implications of their assays.

The 40+-year-long near hiatus in eruptive activity at 

Halema‘uma‘u Crater, from 1924 until the beginning of modern chemical analysis in the 1960s, provided the time needed 

for instrumental analytical techniques to catch up with the 

evolution of conceptual models attempting to describe how 

volcanoes like Kīlauea work. Thermodynamic inferences 

of the evolving conceptual models—for example, speculations about what might be happening to gas emissions when 

magma moves—became better informed by improvements in 

gas-analysis methods. 

The newly refined gas analyses could reasonably be 

considered the equivalent of high-quality chemical portrait 

photographs taken essentially instantaneously in space and 

time. Analogously, the regularly collected COSPEC and 

FLYSPEC SO2 measurements could be symbolized as timelapse chemical photography, albeit for just a single species. 

These time-series gas data consequently signaled a major 

step at HVO toward a dynamic interpretation of volcanic 

processes, using geochemical data in what had formerly been 

solely a geophysical framework. The potential interpretative 

power of merging geophysical and geochemical data streams 

was evident. The addition of gas-geochemical time-series 
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 313

Figure 16. Summit SO2 emission rates increased sharply through 

January and February 2008, while fumarole molecular C/S took on a 

Gerlach and Graeber (1985) Type II eruptive gas signature. A, Heavy 

fuming occurs from vents beyond eastern part of Halema‘uma‘u on 

March 10, 2008. B, On March 12, a new, vigorously fuming area was 

noted near the base of Halema‘uma‘u’s east wall. Over the course of 

the next week, heat flow through this feature increased to the point 

where glow could be clearly seen at night. C, Most summit SO2 release 

was focused on what became known as the Overlook vent, after its 

formation early on the morning of March 19. 

Figure 17. After opening of the Overlook vent in 2008, OP-FTIR 

spectroscopy was routinely used to monitor summit eruptive-gas 

composition. A, Infrared energy emitted by summit lava lake surface 

is absorbed by H2O, SO2, CO2, HCl, HF, and CO that are present in 

light path between sensor and heat source. Absorption is proportional 

to concentration of each gas. Photograph taken February 3, 2012. B, 

Triangle plot of the first summit OP-FTIR analyses (black dot) after 

the formation of the Overlook vent showed gases had molecular C/S 

similar to those of East Rift Zone eruption (red square; an average of 

about 2,000 individual analyses), albeit more water-rich. Notably, this 

Type II, reservoir-equilibrated gas differed significantly from the Type I 

gases detected by T.A. Jaggar, Jr., for the lava lake that had persisted 

throughout the 19th century (blue diamond).

SO2

CO2

H2O

0.00 0.25 0.50 0.75 1.00

0.00

0.25

0.50

0.75

1.00

0.00

0.25

0.50

0.75

1.00

A

B

A

B

C
314 Characteristics of Hawaiian Volcanoes

data notwithstanding, the ability to make sense of the 

geophysical and geochemical changes that began subtly in 

2003 owes its success, in good measure, to observatory-style 

collection of the right volcanic-gas-studies database elements 

years earlier—“observatory-style” in this context meaning 

study of the volcano in a broad and integrated sense scientifically, spatially, and temporally. 

From the conceptual models of Eaton and Murata 

(1960) and M.P. Ryan and colleagues (Ryan and others, 

1981; Ryan, 1987), which constrained Kīlauea’s plumbing 

and, later, its volatile-release characteristics (Gerlach and 

Graeber, 1985), establishment and maintenance of a record 

of summit CO2 emission rates to track magma supply made 

sense. Simultaneously, another dataset measuring East Rift 

Zone SO2 emission rates could be used to track the eruption 

rate. These measurements, together with the long-term 

monitoring of deformation and seismic signals, proved to 

be a good sentinel for the eruptive changes that occurred 

in the mid-2000s. Additionally, the 20+ years of carefully 

recorded summit SO2 emission-rate measurements, along 

with more than 30 years of regular sampling of Kīlauea’s 

summit fumaroles for their molecular C/S ratios, captured 

the 2007–08 transition of summit gas release from passive 

magma-chamber degassing to active eruptive degassing.

The CO2 and SO2 emission-rate records and the longterm study of fumarole chemistry began in response to the 

stimulation provided by a reasoned and persuasive conceptual plumbing and volatile-release model. As obvious as all 

this seems, however, we note that in the early 1980s, when 

Ryan and others’ model was formulated and interpreted 

through Gerlach and Graeber’s volatile budget lens, it was 

unclear precisely how it could be tested operationally and 

refined. Moreover, practical technology to do the testing did 

not yet exist.

The eruptive restlessness at Kīlauea that began with an 

increase in magma supply in 2003 and continued through 

the formation of the summit eruptive vent in 2008 persists 

as of this writing (September 2014). Models constructed 

during the 1960s through the 1980s dealt with scenarios 

that included either a summit or rift-zone eruption. These 

degassing models, however, did not speculate what might 

happen if both summit and rift zone eruptions occurred 

simultaneously. Kīlauea’s eruptive activity since 2008 has, 

among other things, given us the opportunity to extend 

Gerlach and Graeber’s (1985) model that predicted degassing 

of either a summit-reservoir-equilibrated magma or a 

deeper, more primitive magma. Preliminary data from 2012 

indicate that (1) summit and rift zone gas CO2, H2O, and SO2 

emissions are compositionally similar—a condition not seen 

before; and (2) the molecular C/S ratio is not indicative of 

either a reservoir-equilibrated melt or a primitive magma, 

but rather a composition between these two end members. 

These two observations are consistent with a summit magma 

storage complex that at times is intimately connected 

with both the East Rift Zone eruption site and the actively 

convecting lava lake within Halema‘uma‘u Crater.

The simultaneous summit and rift zone activity that 

has provided an opportunity to test and refine earlier 

conceptual models of how Kīlauea works has also helped 

HVO improve its understanding of gas emissions as a 

volcanic hazard. The opening of the summit eruptive vent 

in 2008 substantially increased Kīlauea’s volatile output, at 

least doubling the SO2 emission rate during 2008, relative to 

the annual average of the previous 10 years. The proximity 

of the new summit emission source at Halema‘uma‘u 

to downwind communities exacerbated the detrimental 

effects on those communities, as well as on agriculture, 

ecosystems, and infrastructure. Among these effects was 

a marked increase in the rate of exceedance of air-quality 

standards on the island, and the proclamation of a Federal 

disaster designation for Hawai‘i County in mid-2008 

because of gas-induced agricultural losses.

One way that gas hazard assessments were improved 

in response to the new eruptive activity was through a 

collaborative project between HVO and the University of 

Hawai‘i to develop a vog-forecast model. Functionally, the 

Vog Measurement and Prediction (VMAP) project provides 

island residents and community officials with projections 

of gas and acidic-aerosol-particle exposure as much as 

three days in advance, so that appropriate steps can be 

taken to mitigate risk (Businger and others, 2011). Visitors 

to VMAP’s Web site (http://mkwc.ifa.hawaii.edu/vmap/

index.cgi) find a dynamic vog-forecast map with links to 

the Hawai‘i State Department of Health’s real-time SO2 and 

aerosol-particle data, as well as answers to frequently asked 

questions about vog.

Figure 18. Network of ground-based gas measurement stations 

on Kīlauea Volcano. Yellow areas on map show idealized plumes 

emanating from eruptive vents at summit and along East Rift Zone 

during normal trade-wind conditions. Blue lines, roads, with red sections 

denoting transects used to measure summit and East Rift Zone SO2

and CO2 emissions. Red square, temporary location of ultraviolet-lightsensitive camera system optimized to measure SO2

 and trained on 

Kīlauea’s summit eruptive plume. Top left photograph of ultraviolet (UV) 

camera system, and upper right image (with warm colors indicating 

increased SO2 concentrations), false-color photographic frame, taken 

on September 23, 2011. Lower left photograph, 1 in network of 10 

upward-looking UV spectrometers (green circles), positioned downwind 

of Kīlauea’s summit eruption site, measures SO2 emissions nominally 

every 10 seconds in real time during daylight hours. Lower right 

photograph, one in a network of continuous species-selective sensors 

(yellow triangles) measures gas concentrations and meteorological 

parameters every 10 minutes, telemetering data to HVO, where they 

are merged with emission-rate data and geophysical observations for 

real-time visualization of volcanic activity.
One Hundred Volatile Years of Volcanic Gas Studies at the Hawaiian Volcano Observatory 315

155 ̊25' 155 ̊05'

19 ̊25'

155 ̊20' 155 ̊10'

19 ̊20'

19 ̊15'

155 ̊15'

Island of Hawai‘i

SO2 column density, in parts per million

0

1,000

2,000

3,000

4,000

5,000

6,000

7,000

Paired UV-sensitive cameras

Upward-looking

UV spectrometer

Continuous gas and

meteorologic sensors

0

0

2.5 5 KILOMETERS

1 2 3 MILES
316 Characteristics of Hawaiian Volcanoes

Thomas A. Jaggar, Jr.’s, provocative vision of a human 

presence to continuously observe active volcanism from 

the brink of Kīlauea Caldera has served the volcanological 

and residential communities well, not just in Hawai‘i but 

worldwide. At HVO, volcanic gas studies have evolved 

considerably over the first hundred years of HVO’s 

existence. We now have a good understanding of volcanic 

gas in terms of “what’s in it.” As we begin the next hundred 

years of volcanic gas studies, we are more likely to be asking 

process-related questions, like “What do the near-real-time 

changes we’re measuring in species W and X tell us about 

conditions at depth Y in light of the contemporaneous tilt, 

seismic, and lava level data?” What we have already learned 

from the process side about gases has us probing hazards 

issues, as well, leading to such real-time questions as “What 

might be the implications of observed changes in species W 

and X for people, agriculture, and infrastructure located Z 

km downwind of gas-release points?” 

The future of volcanic gas studies in Hawai‘i is bright. 

HVO and its collaborators have already begun transitioning 

from vehicle-based SO2-emission-rate measurements to using 

the recently installed array of upward-looking spectrometers 

designed to continuously measure the SO2 emitted from the 

summit eruptive vent during daylight hours (green circles, 

fig. 18). Concurrently, a newly established network of 

continuous gas monitors distributed from near-vent to 

>60 km distant measures and transmits the downwind SO2

concentrations to HVO (yellow triangles, fig. 18). These data 

can be visualized in real time, along with geophysical and 

geologic data. With colleagues at CVO, HVO is testing the 

utility of an environmentally hardened SO2-imaging camera 

system that operates continuously, producing streaming 

video of plume-gas emissions from Kīlauea’s summit during 

daylight hours (red square, fig. 18; Kern and others, 2013). 

Refinements to CO2 measurement technology are extending 

the useful life of these sensors under rough field conditions, 

thus improving the quality of in situ molecular C/S ratios. 

These and other improvements will enhance our ability to 

detect changes in volcanic processes within the summit 

magma storage complex.

Just as fuel, air, and heat combine to make a fire, so 

the timing of eruptive activity, unwavering human curiosity, 

and an explosive growth of technology have combined to 

address the persistent question of how volcanoes, like those 

in Hawai‘i, work.

Acknowledgments

The opportunity to review a century of progress in any 

field provides a potent historical perspective on the subject. 

In relation to Hawaiian volcanic gas studies, this study has, 

for the authors, cultivated a deep respect for generations 

of workers through the years who struggled to keep telling 

the story, both accurately and in a way that would allow the 

data of their own era to be directly comparable with those 

obtained before and after. Accordingly, over the past 20 

years, we gratefully acknowledge the many hours of hard 

work contributed by USGS volunteers. We also thank Ben 

Gaddis, Ken McGee, and Barry Stokes of the USGS for their 

help with compiling the unwritten history, and Shaun Hardy 

of the Carnegie Institution’s Geophysical Laboratory library 

for photoarchival assistance. The publication process was 

greatly facilitated by our encouraging and supportive editors, 

and the manuscript benefited from helpful reviews by Tom 

Casadevall and Lopaka Lee. 

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2001, Implications for eruptive processes as indicated by 

sulfur dioxide emissions from Kīlauea Volcano, Hawai‘i, 

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pp1676.] 

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agu.org/meetings/2009/FM/sections/V/sessions/V52B/

abstracts/V52B-08.html. 
Hawaiian Volcano Observatory gas geochemists Tamar Elias (standing) and A. Jeff Sutton (sitting) collect 

spectra of gases emitted from Kīlauea’s East Rift Zone. USGS photograph by L. Lee, July 24, 2007.
A lava bubble bursts from the lava lake in Halema‘uma‘u Crater, within the summit caldera of Kīlauea Volcano. USGS photograph by T.A. 

Jaggar, Jr., May 17, 1917. 
Abstract

This chapter, prepared in celebration of the Hawaiian 

Volcano Observatoryʼs centennial, provides a historical lens 

through which to view modern paradigms of Hawaiian-style 

eruption dynamics. The models presented here draw heavily 

from observations, monitoring, and experiments conducted on 

Kīlauea Volcano, which, as the site of frequent and accessible 

eruptions, has attracted scientists from around the globe. 

Long-lived eruptions in particular—Halema‘uma‘u 1907–24, 

Kīlauea Iki 1959, Mauna Ulu 1969–74, Pu‘u ‘Ō‘ō-Kupaianaha 

1983–present, and Halema‘uma‘u 2008–present—have offered 

incomparable opportunities to conceptualize and constrain 

theoretical models with multidisciplinary data and to fieldtest model results. The central theme in our retrospective is 

the interplay of magmatic gas and near-liquidus basaltic melt. 

A century of study has shown that gas exsolution facilitates 

basaltic dike propagation; volatile solubility and vesiculation 

kinetics influence magma-rise rates and fragmentation depths; 

bubble interactions and gas-melt decoupling modulate magma 

rheology, eruption intensity, and plume dynamics; and pyroclast 

outgassing controls characteristics of eruption deposits. 

Looking to the future, we anticipate research leading to a 

better understanding of how eruptive activity is influenced by 

volatiles, including the physics of mixed CO2-H2O degassing, 

gas segregation in nonuniform conduits, and vaporization of 

external H2O during magma ascent.

Introduction

Tonight in presence of the firepit, with the glowing 

lava again gushing and pounding and shaking the 

seismographs, I am asked to tell you something of 

the scientific meaning of it.

 —Thomas A. Jaggar, Jr. (1945)

1

U.S. Geological Survey.

2

University of Bristol, U.K.

Of endless fascination to the observers of Hawaiian eruptions 

is the “marvelous . . . superlative mobility” of molten lava (Perret, 

1913a). This observation is fundamental, for ultimately the fluidity 

of basaltic magma is what distinguishes the classic Hawaiian-style 

eruption from other forms of volcanism. Eruption of free-flowing 

tholeiitic basalt feeds the iconic images of Hawaiian curtains of 

fire, lava fountains, lava lakes, Pele’s hair, and thread-lace scoria 

(fig.1). Moreover, thin flows of fluid basalt are central to Hawaiian 

shield building, requiring quasi-steady-state magma supply and 

throughput for hundreds of thousands of years.

Early observations of the eruptive activity at Kīlauea and 

Mauna Loa volcanoes demonstrated that their broad, shield-like 

shape results chiefly from rift-zone eruptions fed from a summit 

magma reservoir and that the summit reservoir itself comprises 

interconnected storage regions that fill and drain repeatedly 

(for example, Eaton and Murata, 1960; Eaton, 1962; Fiske and 

Kinoshita, 1969; Wright and Fiske, 1971). 

The classic Hawaiian-style eruption typically initiates with 

a fissure emanating from en echelon fractures formed as a dike 

propagates across the caldera floor or down the rift zone from 

the main magma-storage region. Within hours to days, parts of 

the fissure seal off and focus the eruption into a solitary lava 

fountain towering tens to hundreds of meters above a central 

vent. A rain of molten pyroclasts splashes down around the vent, 

feeding rootless lava flows that spread across the landscape. 

The incandescence of liquid rock is what draws the eye of the 

observer, but it is actually a subordinate player, with the volume 

of magmatic gas erupted exceeding the volume of lava by as 

much as 70 to 1 (Greenland and others, 1988).

Once established, a central vent may host many fountaining 

episodes. New episodes start with a column of lava quietly rising 

in the conduit. Its upward progression is punctuated by rhythmic 

rise and fall as trapped gas is released. Low dome fountains 

appear, building to a steady, high fountain over minutes to hours. 

Fountaining episodes may persist for hours to days, discharging 

lava at rates of a few hundred cubic meters per second. In contrast 

to their slow buildup, fountain episodes quit abruptly, ending in 

chaotic bursts of gas and dense spatter as lava ponded around the 

vent is swallowed back down into the conduit. 

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 8

The Dynamics of Hawaiian-Style Eruptions: A Century of Study

By Margaret T. Mangan1, Katharine V. Cashman2, and Donald A. Swanson1
324 Characteristics of Hawaiian Volcanoes

A

C

D

B

E

Figure 1. Fluidity of basalt at near-liquidus temperatures is revealed in fountains of molten rock, spattering lava lakes, delicate magmatic 

foams, and spun glass filaments of classic Hawaiian-style eruptions. A, 60- to 100-m-high fountains from a fissure eruption in 1983, showing a 

500-m-long segment of an eruptive fissure that extended 1 km along the East Rift Zone. Photograph by J.D. Griggs. B, 320-m-high lava fountain 

from first phase of the 1959 Kīlauea Iki eruption at Kīlauea’s summit. Molten fallback from fountain feeds a river of lava 90 m wide. Photograph 

by J.P. Eaton. C, Mauna Ulu lava lake in September 1972. Lake circulation is from left to right, with 5- to 10-m-high spattering occurring at lake 

margin, where pliable lava crust circulates downward. Lava-lake circulation patterns were vividly described by Jaggar (1917): “. . . crusts form, 

thicken and stream across the lava lake to founder with sudden tearing, downsucking, flaming and violent effervescence. . . .” Photograph by R.I. 

Tilling. D, Delicate structure of thread-lace scoria (reticulite), a term coined by J.D. Dana after visiting Kīlauea in 1887 (Dana, 1890). View through 

a 5X binocular microscope reveals an open honeycomb of polygonal cells. Bubble walls (films) thin to the point of rupture as gas expands, leaving 

a rapidly quenched skeleton of glass struts, trigonal in cross section (plateau borders) and ~0.05 mm thick, that meet to mark original foam 

structure created during high lava fountaining. Photograph by M.T. Mangan. E, Pele’s hair (spun volcanic glass) as viewed through a 5X binocular 

microscope. Molten drops of basaltic lava are stretched into long glass filaments, <0.5 mm in diameter and as much as 0.5 m long, in turbulent 

updrafts surging above active vents and skylights in lava tubes. Photograph by M.T. Mangan. 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 325

An epoch of episodically occurring lava fountains 

eventually closes with diminished fountain heights and 

declining discharge rates. The end of fountain activity may 

signal the end of an eruption or, alternatively, the beginning 

of steady, effusive activity from an open vent or lava lake.

Early accounts of lava-lake dynamics used vivid language, 

describing “terrific ebullitions,” “splashing sinkholes,” 

“floating islands,” “thin skins,” and “plastic crusts” as the stuff 

of Halema‘uma‘u lava lake. In today’s vernacular, the salient 

features include a lake-bottom feeder conduit driving sporadic 

surface spattering and (or) low fountains at the surface; 

complex, lake-wide circulation patterns; and rhythmic rise-fall 

cycles associated with accumulation of gas bubbles. 

In this chapter, we provide the reader with a historical 

perspective on how we came to know what we know of 

the dynamics of Hawaiian-style eruptions, ending with 

an eye toward what is yet to be learned. As the site of 

the Hawaiian Volcano Observatory (HVO) and home to 

frequent, accessible eruptions, Kīlauea Volcano (fig. 2) has 

played a pivotal role in shaping research and technological 

innovation. It is thus fitting that we start with an overview 

of milestone Kīlauea eruptions that have fed our evolving 

understanding of how volcanoes work.

Halema‘uma‘u 1907–24

From 1907 until 1924, scientific research at Kīlauea 

Volcano centered on the persistent lava lake within 

Halema‘uma‘u, laying the groundwork for the first models 

of magmatic convection and open-system degassing. Daly 

(1911) and Perret (1913b) envisioned cellular convection 

in the lake, with buoyant upwelling of bubble-rich magma 

and dense, degassed lava sinking with much spattering and 

sloshing at the opposite end. These early workers realized 

that exchange flow was required to supply the heat necessary 

for persistent activity, but they differed in their views of the 

geometry and composition of exchange flow—whether gascharged magma or gas alone was exchanged, and whether 

convection cells were confined to the lake or penetrated deeper 

into the conduit. Perret (1913b), in particular, emphasized 

open-system degassing as critical to lake circulation, arguing 

that downwelling of degassed lava creates a “powerful siphon 

effect that is the mainspring of the circulatory system” and 

suggesting that descending lava is “re-heated and re-vivified” 

by infusion of minute gas bubbles rising from the depths of the 

conduit. Geophysical and geochemical investigations of recent 

lava lakes (Swanson and others, 1979; Tilling and others, 1987; 

Figure 2. Sketch 

map of southeastern 

Hawai‘i, showing summit 

calderas and rift zones 

of Mauna Loa and 

Kīlauea volcanoes and 

locations of the Hawaiian 

Volcano Observatory 

(HVO) and Kīlauea’s 

Halema‘uma‘u, Kīlauea 

Iki, Mauna Ulu, and Pu‘u 

‘Ō‘ō-Kupaianaha eruptive 

vents. Modified from 

Babb and others (2011). 

Lōʻihi

Hilo

Kapoho

Hawaiian Volcano

Observatory

Mauna Loa

summit caldera

(Mokuʻāweoweo) Kīlauea Iki

Mauna Ulu 

Puʻu ʻŌʻō

East Rift Zone

Southwest 

Rift 

Zone

Northeast Rift Zone

Southwest Rift

Zone

Hualālai

Mauna Kea

Kohala

HAWAI‘I

Area

of map

Halemaʻumaʻu

Kīlauea

summit 

caldera 

Kalapana

Kīlauea

Kīlauea

Mauna Loa

155°45' 155°30' 155°15' 155°

19°

19°15'

19°30'

19°45' Mauna Kea

Mauna Loa

0

0

10 20 KILOMETERS

5 10 MILES

Kupaianaha
326 Characteristics of Hawaiian Volcanoes

Johnson and others, 2005; Patrick and others, 2011a; Orr and 

Rea, 2012) have pursued the role of open-system degassing 

in controlling lake activity and longevity and have opened the 

door to new questions on how and where gas accumulates in 

shallow plumbing systems and on the mode of gas release.

Kīlauea Iki 1959

The Kīlauea Iki summit eruption marked a major 

milestone in HVO’s history (Eaton and Murata, 1960; Murata, 

1966; Murata and Richter, 1966; Richter and Moore, 1966; 

Richter and Murata, 1966; Richter and others, 1970; see Helz 

and others, this volume, chap. 6). This spectacular eruption was 

preceded by an intense swarm of deep (45–60km) earthquakes 

and followed by weak tremor in the middle crust and then 

inflation at the summit. Within 3 months of the onset of deep 

seismicity, a fissure erupted high on the wall of Kīlauea Iki pit 

crater. Over 36 days, 17 unusually high and hot fountaining 

episodes filled the pit crater below with 37millionm3 of picritic 

basalt. Fountaining episodes terminated abruptly when the level 

of the lava lake reached, then flooded, the vent.

The arrival of new staff and modernization of the monitoring 

network just before the Kīlauea Iki eruption gave HVO an 

unprecedented opportunity for comprehensive monitoring of the 

eruption. For the first time, the course of magma—from mantle 

source to surface vent—could be quantitatively charted by 

migrating earthquakes, ground deformation, and lava extrusion. 

A systematic program of lava sampling and measurements of 

fountain height, discharge rate, and lava temperature provided data 

critical to constraining later eruption dynamics models. 

Mauna Ulu 1969–74

The nearly 5-year-long Mauna Ulu eruption gave HVO 

staff their first opportunity to study long-lived rift-zone 

volcanism (Peterson and others, 1976; Swanson and others, 

1979; Tilling and others, 1987). Mauna Ulu’s repertoire 

was diverse, including 12 episodic lava-fountain events, 

near-continuous lava-lake activity, shield building, lava 

tube formation, and littoral lava-seawater interactions along 

Hawai‘i’s south coastline. The final 2 years of the eruption were 

punctuated by brief outbursts from fissures opening uprift of the 

Mauna Ulu vent, including a 3-day eruption at the summit.

Throughout the eruption, seismicity and deformation at 

Kīlauea’s summit correlated with events at the Mauna Ulu vent, 

providing new insight into how summit-reservoir overpressure 

is balanced by waxing and waning eruptive output. The term 

“gas pistoning” was coined during the Mauna Ulu eruption 

to describe cyclic degassing bursts that produced slow rise 

and rapid fall of the lava column (Swanson and others, 1979; 

Tilling and others, 1987). Rise-fall cycles typically occurred 

over tens of minutes to a few hours. The surface of the lava 

lake is relatively calm during the rise in a cycle but erupts into 

vigorous bubbling and spattering as gas escapes and the lava 

level drops. These observations, which have fueled numerous 

laboratory experiments and numerical models for two-phase 

flow, provide a conceptual link between the dynamics of 

Hawaiian and Strombolian eruption styles.

Pu‘u ‘Ō‘ō-Kupaianaha and Halema‘uma‘u 1983–

Present 

The present Pu‘u ‘Ō‘ō-Kupaianaha eruption marks the 

most voluminous outpouring of lava on Kīlauea’s East Rift Zone 

since the 15th century and is the longest East Rift Zone eruption 

ever recorded (see descriptions by Wolfe and others, 1987, 

1988; Heliker and Wright, 1991; Heliker and Mattox, 2003; 

Poland and others, 2008). Five main epochs capture the diverse 

and sometimes destructive characteristics of this extraordinary 

eruption. The first epoch began in January 1983 with 24 hours 

of migrating earthquakes and ground deformation as a dike from 

Kīlauea’s summit propagated down the East Rift Zone. The 

dike, traveling 550 m/h laterally and 70 m/h vertically, breached 

the surface in a linear series of fissures 8 km long (Wolfe and 

others, 1987). Within 5 months, the fissure eruption had localized 

to a single vent, Pu‘u ‘Ō‘ō, that hosted more than 3 years of 

remarkably regular episodic lava fountains. 

In July 1986, after 44 episodes of high fountaining, 

the conduit beneath Pu‘u ‘Ō‘ō ruptured. Thus began a 

second epoch characterized by quasi-steady effusion from 

Kupaianaha vent, 3 km downrift of Pu‘u ‘Ō‘ō. The next 51⁄2 

years saw construction of a broad shield with a summit lava 

lake and a well-developed lava tube system that efficiently 

delivered flows from Kupaianaha to the coastline, igniting 

dramatic littoral activity and overrunning much of the 

community of Kalapana. Starting in mid-1990, Kupaianaha 

entered a protracted stage of diminishing output and frequent 

pauses (Kauahikaua and others, 1996). Kupaianaha’s 

slow decline revitalized Pu‘u ‘Ō‘ō. By February 1992, 

Kupaianahaʼs activity had ceased, and the eruption was back 

at Pu‘u ‘Ō‘ō, starting the third epoch.

Over the next 15 years, effusion of lava from Pu‘u ‘Ō‘ō 

was nearly continuous, interrupted only temporarily by nearby 

fissure eruptions and East Rift Zone intrusions (particularly in 

1997, 1999, and 2007). In July 2007, the eruption moved down 

the rift zone again to a site between Pu‘u ‘Ō‘ō and Kupaianaha, 

starting the fourth epoch. Effusion from this vent lasted 

until March 2011 and again sent lava into Kalapana, causing 

additional destruction. Like Kupaianaha, effusion from that vent 

eventually waned, ending in March 2011, when another brief 

East Rift Zone fissure eruption occurred. The fifth, currently 

ongoing (as of September 2014) epoch has been marked by a 

return to effusion from Pu‘u ‘Ō‘ō after that fissure eruption.

In March 2008, an eruptive vent also opened at Kīlauea’s 

summit (Houghton and others, 2011; Swanson and others, 

2011). Pu‘u ‘Ō‘ō still pumps out lava on the rift zone, but in 

addition, a lava lake circulates within the recesses of a crater on 

the east margin of Halema‘uma‘u Crater within Kīlauea Caldera 

(Patrick and others, 2011a, 2013). First appearing as a fuming, 

incandescent crack, the new vent evolved by way of a series of 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 327

magmatic gas blasts during 2008 and 2011 into an active lava 

lake more than 100 m in diameter. As of this writing (September 

2014), neither vent shows any sign of shutting down.

The East Rift Zone eruption, which celebrated its 31st year 

in January 2014, has seen the introduction of new, sophisticated 

technology and a steady influx of scientists from diverse 

disciplines. Major efforts have gone into interpreting integrated 

datasets from state-of-the-art geophysical and geochemical 

sensors. Perhaps most significant in this period of rapidly evolving 

capabilities, at least from the perspective of eruption dynamics, 

is the initiation of systematic gas monitoring just before the start 

of the Pu‘u ‘Ō‘ō eruption. Valuable volcanic gas samples were 

collected during the 1912–24 Halema‘uma‘u activity (Jaggar, 

1940; reanalyzed by Gerlach, 1980), but thereafter, only a 

few sporadic samples were obtained. Not until the installation 

of HVO’s gas-analysis laboratory in 1980 did monitoring of 

volcanic gases take a prominent role, alongside monitoring of 

ground deformation and earthquakes, in building the conceptual 

framework upon which models of magma transport and 

eruption are built (see sections below entitled “Geochemistry of 

Degassing” and “Overview of Existing Eruption Models,” as well 

as Sutton and Elias, this volume, chap. 7). 

Magma Transport 

The continuance of eruption at any point depends on 

victory in the struggle with cold.

 —Reginald A. Daly (1911)

Early 20th century volcanologists recognized that magma 

came to the surface through dike propagation. Observing relict 

gas cavities in the tops of exhumed laccoliths exposed in the 

western United States, Daly (1911) postulated a “blowpipe” 

mechanism for dike propagation in which hot, caustic gases 

digested rock strata in advance of rising magma. Perret (1913c) 

similarly proposed gas-driven “trepanning” and “stoping” of 

a “vertical tunnel” in overlying strata through which buoyant 

magma could pass (figs. 3A,B).

Later field observations in Hawai‘i, particularly on 

the Island of O‘ahu, supplied data on the size, shape, and 

orientation of dikes, providing the first quantitative framework 

for modeling the delivery of basaltic magma to the surface 

(Stearns and Vaksvik, 1935; Stearns, 1939; Wentworth and 

Jones, 1940; Wentworth and Macdonald, 1953). Blade-shaped 

dikes a few meters thick emerged as “typical” of Hawaiian 

volcanic systems. Stearns and Vaksvik (1935), working in the 

eroded Ko‘olau shield on O‘ahu, may have been the first to use 

the term “dike complex” in Hawai‘i to describe the recurrence 

of near-vertical preferred pathways, citing traverses with as 

many as 100 dikes in 1 mile.

At mid-century, a dynamic picture of magma transport 

emerged from seismic and deformation networks installed just 

before the 1959 Kīlauea Iki eruption. The results were described 

by J.P. Eaton, twice director of HVO, who used patterns of 

earthquake and ground deformation to construct a working 

model of Kīlauea that charted the “course of magma as it 

accumulates in a shallow reservoir within the volcano and as it 

migrates from this reservoir into the rift zones” (Eaton, 1962). 

Eaton’s conceptualization provided the framework needed 

to apply theoretical fracture mechanics, fluid dynamics, and 

thermodynamics to the problem of magma transport in Hawai‘i. 

Mechanics of Dike Propagation

Time scales of shallow dike injections are sufficiently 

short to be observed directly (hours to days), so that as HVO’s 

monitoring networks became more robust, geophysical data 

fostered many studies of dike propagation. Studies coupling 

classical fracture mechanics with the physics of fluid flow were 

conducted, integrating the driving forces of magma pressure and 

buoyancy with the resisting force of rock tensile strength.

Hydraulic fracturing of rock during shallow dike 

propagation at Kīlauea requires magma pressures of 2 to 

10 MPa (Aki and others, 1977; Rubin and Pollard, 1987). 

Stress concentration at the dike tip is the leading wedge 

for crack initiation (fig. 3C). Studies by Rubin and Pollard 

(1987), Lister and Kerr (1991), and Rubin (1993) suggested 

that a vanguard of gas is as important as magma pressure in 

dike propagation, but not as the blowpipe that Daly and his 

contemporaries postulated at the beginning of the century. 

Instead, microfractures in the damage zone at the dike tip are 

infiltrated by low-viscosity aqueous fluid (either exsolved 

magmatic volatiles or external pore fluids). Pressurized 

microcracks lower the effective tensile strength of the rock and 

coalesce, creating larger cracks into which the more viscous 

fluid (magma) can penetrate to extend and widen the dike. 

Ryan (1988) used a finite-element model to show that a tensile 

stress regime exists at the dike tip and a compressive stress 

regime at the margins where the dike is widening. Seismicity 

associated with dike propagation was attributed to stress 

release in zones of compression (Aki and others, 1977; Klein 

and others, 1987). As the surface is approached, overburden 

pressure decreases, causing torsion of the local principal 

stresses and segmentation of the dike tip (Pollard, 1973; Hill, 

1977; Delaney and Pollard, 1981; Pollard and others, 1982; 

Rubin and Pollard 1987; Ryan, 1988) and resulting in an en 

echelon set of eruptive fissures—the signature opening stage 

of Hawaiian-style eruptions (fig. 3D).

Thermodynamics of Magma Flow in Dikes and 

Fissures

Ryan (1988) described the thermodynamics of dike 

propagation at Kīlauea as a process of natural selection by 

which only the “fortunate dikes” traversing the hot, central 

core of the magma complex make it to the surface before 

freezing. His assertion is supported by the dikes exposed 

in the Ko‘olau complex, which generally intrude along the 

margins, or median axis, of earlier, but still hot, dikes (fig. 4). 
328 Characteristics of Hawaiian Volcanoes

Dike segment

Trend of dike

at depth

Propagation direction

Least-principal-stress 

direction

A B

a

θ

y

x

r << a

σxy σyy

σxx

C D

Figure 3. Past (A,B) versus present (C,D) conceptualizations of basaltic dike propagation. Past conceptualizations (modified 

from Perret, 1913c, and Daly, 1911), show lava columns capped with hot, compressed gas (gray shading) fluxing (A) and stoping 

(B) paths to surface. Present conceptualizations C and D (modified from Rubin and Pollard, 1987, and Delaney and Pollard, 1981, 

respectively) illustrate application of modern fracture mechanics. In figure 3C, σ is the orthogonal (xx,yy) and oblique (xy) stress 

component at dike tip along ray path, r, at an angle θ from vertical dike axis for r<<a, where a is dike half-length. Gray shading 

represents a vanguard of low-viscosity fluid (volatiles). Part D illustrates near-surface, en echelon segmentation of a rising tabular 

dike due to rotation in a plane perpendicular to propagation direction (least-principal-stress direction).
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 329

0

0 2 MILES

5 KILOMETERS

21°30'

21°25'

157°50' 157°45'

8 25

20

50

5

8

40 20

25

22

2 2

55 75

70

95

30 80

90

85

80

7

5 75

75

80

55

90

60

60

50

70

100

120 85

60 33

13

24

37

6

4

1

7

7 130

125 8 3

75 50

85

40

17

7

K O ‘ O L A U R A N G E

12

EXPLANATION

Number of dikes per 100-m 

traverse normal to 

predominant dike trend

>70 dikes per 100 m

20 to 70 dikes per 100 m

Crest of pali and ridges

0 1 2 METERS

0 2 4 6 FEET

Northwest Southeast

2 3 4

1

1

2 4 3

1

1

1

1

1

1

1 2

4 2 3

3

 EXPLANATION

Country-rock 

basalt

Earliest

intruded 

dikes

Later dikes

Kailua Bay

Kāne‘ohe Bay

158°

21°30’ O‘ahu Area

of map

Likewise, dike intensity (percentage of host rock occupied by 

dikes) drops precipitously outside the thermal core of the dike 

complex—in the Ko‘olau complex it decreases from 50–70 to 

<5 percent over a few hundred meters (Walker, 1987). Helz 

and others (in press and this volume, chap. 6) used olivineglass geothermometry of rapidly quenched pyroclasts to infer 

ambient temperatures of 1,145–1,235 °C within the active 

dike-and-sill complex at Kīlauea’s summit.

The first thermodynamic models avoided the 

complexities of multiple dike injection and treated the problem 

of a single dike invading cold country rock under conditions 

of steady magma supply (Delaney and Pollard, 1982; Hardee, 

1987; Bruce and Huppert, 1989, 1990). The initial heat 

exchange creates a layer of congealed magma along the walls 

of the dike, confining flow to its central region. The growth 

of the congealed zone and eventual freezing of the magma 

hinges on the balance between heat conducted out of the 

walls of the dike and advected in the direction of flow. Using 

typical values for dike thickness, magma pressure, magma 

temperature, viscosity, and magma-flow rates, models for 

Kīlauea predict that magma flowing a few kilometers from its 

source should freeze within hours of emplacement, assuming a 

normal geothermal gradient. Successful propagation of a dike 

to the surface therefore requires that it follow a thermal “hot 

zone” built from decades, if not centuries, of intrusive activity. 

Thermodynamic models of Kīlauea’s upper East Rift Zone 

(Hardee, 1987) suggest that an intrusive flux of ~10-3 km3/y 

over a span of 2–3 decades is required to form hot pathways 

for aseismic magma transport without brittle fracture.

The dikes that do breach the surface to feed fissure 

eruptions face additional heat loss through radiation, which 

increases as the fourth power of absolute temperature. Initial 

fissure sections shut down within hours to a few days of 

opening as the pressure driving magma flow diminishes, or 

because the magma-transport pathways focus at one or a few 

central vents. Numerical models (Delaney and Pollard, 1982; 

Bruce and Huppert, 1989, 1990; Wylie and others, 1999) 

indicate that focusing the flow of magma from an initial linear 

Figure 4. Sketch map showing high concentration of basaltic dikes composing the “paleo hot zone” in eroded core of Ko‘olau volcano 

in eastern O‘ahu. G.P.L. Walker mapped 7,400 subparallel dikes totaling 3 to 5 km in thickness in widest part of complex. Inset (area 

denoted by star) shows part of a succession of 14 intrusions across an outcrop 13.6 m wide. Modified from Walker (1987). 
330 Characteristics of Hawaiian Volcanoes

geometry to one or a few central vents strongly depends on 

fissure geometry. When the fissure varies in width, advective 

heat is reduced within narrow sections, owing to lower flow 

rates. Reduced heat advection promotes solidification, which 

contributes to further narrowing, slowing, and advective 

heat loss. This self-destructive feedback gradually closes 

off narrow fissure sections, although dikes may fail to close 

completely when cooling at the thin edges is combined with 

inelastic deformation (Daniels and others, 2012). If the magma 

pressure driving the eruption remains fixed, increased flow 

rate in the open sections can widen dikes by melting the layer 

of congealed magma formed earlier at the wall (Bruce and 

Huppert, 1989, 1990). Attainment of a stable fissure width 

reflects thermal equilibrium between conducted, radiative, 

and advective heat. The duration of flow is then controlled 

by magma pressure. A particularly well documented field 

example is afforded by the shutdown of the Kupaianaha 

vent at Kīlauea after nearly 6 years of quasi-steady magma 

flow. Here, the temperature of erupted lava was constantly 

1,152±3 °C over a 10-month period in which magma supply 

dropped linearly with time from 2.5×105 to 5.4×104 m3/d. The 

linear decay allowed accurate prediction of vent closure and 

supported a cessation model based on pressure loss rather than 

cooling (Kauahikaua and others, 1996). 

Magma Degassing

All . . . workers agree that rising gas achieves the 

work known at the surface of the earth as volcanic 

activity. 

 —Thomas A. Jaggar, Jr. (1917)

That gas is a driving force for subaerial volcanic 

eruptions was well known to early 20th century 

volcanologists (fig. 5A; Hitchcock, 1911; Day and Shepherd, 

1913; Perret, 1913b; Jaggar, 1917). Gas played a central 

role in the vivid accounts of Halema‘uma‘u pyrotechnics 

from 1907 to 1924, but the origin and constitution of the gas 

were unresolved, fueling nearly a half-century of vigorous 

debate. Resolution of contentious issues was beyond reach 

until methods of measuring gas species and fluxes were 

improved, and the concept of thermodynamic equilibrium of 

volatile species was adopted. Studies of the physics of gas 

exsolution and transport have lagged behind geochemical 

inquiries, awaiting better constraints on melt viscosity, 

melt-vapor surface tension, and volatile diffusivity. As 

experimental data on the relevant physical properties 

amassed, investigation of the kinetics and mechanisms of 

phase transformation commenced through application of 

classical nucleation theory and quantitative textural analysis 

of pyroclasts. Adaptations of engineering models for twophase flow (bubble and melt) fueled quantitative analysis 

of how gas propels and fragments magma during the final 

stages of ascent. 

Geochemistry of Degassing

From the early days of HVO, Halema‘uma‘u lava lake 

provided scientists with direct access to volcanic gases (see 

Sutton and Elias, this volume, chap. 7). Gas samples collected 

between 1912 and 1919 established carbon and sulfur as 

important components in volcanic gas, but the role and 

origin of water in Hawaiian eruptions remained an enigma. 

Some workers proposed that percolating meteoric water (or 

seawater) was ingested by magma at depth (Green, 1887; 

Dana, 1890; Brun, 1911). In opposition, Day and Shepherd 

(1913) argued that “[water] is entitled to be considered 

an original component of the lava with as much right as 

the sulphur or the carbon,” whereas Jaggar (1917) stated 

unequivocally that gases other than water (sulfur and carbon) 

“operate the volcanic engine.” 

As time went on, earlier maxims were shed. Shepherd 

(1938) discounted his early work on Halema‘uma‘u gas 

samples, lamenting the impossibility of determining original 

compositions and concentrations of juvenile magmatic gas 

from the types of samples that were collected. By that time, he 

recognized that the gas samples were invariably contaminated 

by atmospheric and organic components, regardless of 

the volcanic setting or the collection methods used. After 

reevaluating data from samples collected in 1917–19, Jaggar 

(1940) also reversed his original assertion and conceded 

that H2 is a primal magmatic volatile, although he remained 

convinced that molecular H2O is not primary. 

Progress on gas geochemistry stalled until midcentury, 

when Ellis (1957) showed that “modern” chemical 

thermodynamics could be used to establish relevant volatile 

equilibria. Moreover, he demonstrated that balanced reactions 

could be used to deduce the primary volatile species in 

samples at conditions other than those of collection, thus 

providing a means to correct for unavoidable atmospheric and 

organic contamination. His work stimulated a flurry of studies 

that used a thermodynamic-equilibrium approach, firmly 

establishing C-O-H-S as the primal volatile species dissolved 

in Hawaiian basaltic magma (Matsuo, 1962; Heald, 1963; 

Finlayson and others, 1968; Gerlach and Nordlie, 1975a-c; 

Volkov and Ruzaykin, 1975). Thermodynamic calculations 

were facilitated by increasing use of digital computers (Heald 

and Naughton, 1962) and new sampling techniques (for 

example, Naughton and others, 1963, 1969; Giggenbach 1975; 

see Sutton and Elias, this volume, chap. 7).

A critical link between gas thermodynamics and eruption 

dynamics was forged in papers that placed C-O-H-S geochemical 

studies in the context of the geometry of Kīlauea’s plumbing 

system and magma-supply rates (fig.5B; Gerlach and Graber, 

1985; Greenland and others, 1985; Gerlach, 1986; Gerlach 

and others, 2002). These papers demonstrated that at mantle 

levels, ~0.3 weight percent H2O, 0.7 weight percent CO2, and 

0.1 weight percent S are dissolved in primary Kīlauea magma. 

About 90 weight percent of the CO2 exsolves during ascent to 

the base of the summit reservoir at ~7 km (180 MPa lithostatic 

pressure, 4 volume percent gas). The newly arrived magma is 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 331

CO2-rich plume

(non-eruptive)

H2O-rich

plume

(eruptive)

Parental magma

(parental volatiles wt% CO2>H2O>S)

Reservoir equilibrated magma

(stored volatiles wt% H2O>S≥CO2)

Active

Extinct Extinct

Juvenile gas

A

B

strongly buoyant, relative to the resident magma, and should 

rise from the point of entry as a turbulent plume that mixes 

thoroughly and equilibrates to chamber conditions (75–180 

MPa; Gerlach and others, 2002). In recent times, CO2-rich 

gases are emitted passively from the summit reservoir at 

a rate of 9,000–25,000 metric tons per day (Gerlach and 

others, 2002; Poland and others, 2012; see Sutton and 

Elias, this volume, chap. 7). Reservoir-equilibrated magma 

contains dissolved volatile concentrations of ~0.3 weight 

percent H2O, ~0.1 weight percent S, and from 0.02 to 0.1 

weight percent CO2, depending on the depth of equilibration 

(Greenland and others, 1985; Gerlach and others, 2002). 

These residual volatiles exsolve at very low pressures 

(<10MPa) as the reservoir-equilibrated magma approaches 

the surface during eruptions.

Physics of Gas Bubbles

In comparison with the geochemistry of gas, the physics 

of vesiculation and its bearing on eruption processes received 

little quantitative attention until the second half of the 20th 

century because data constraining the physical properties 

of basaltic melt were few. Recognizing the importance of 

quantitative physical characterization of magma, Macdonald 

(1963) compiled existing early measurements of temperature, 

bulk viscosity, dissolved-gas content, and melt density 

of erupting Hawaiian basalt. In the 1970s and 1980s, 

the application of analytical equipment capable of highresolution, direct measurements provided additional data, 

including the poorly known, but essential, properties of meltvapor surface tension and volatile diffusivity (for example, 

Bottinga and Weill, 1972; Shaw, 1972; Murase and McBirney, 

1973; Khitarov and others, 1979; Watson and others, 1982; 

Zhang and Stolper, 1991).

Kinetics of Bubble Nucleation and Growth

A pivotal paper by Sparks (1978) provided an early, 

physics-based model for the kinetics of bubble growth during 

magma ascent. The crux of his paper is a numerical model for 

diffusional and decompressional bubble growth under various 

magma-ascent rates, melt viscosities, and H2O concentrations, 

solubilities, and diffusivities. Diffusive growth, modeled as 

rB(Dt)

-1/2 (Scriven, 1959), where r is the bubble radius, B 

is the rate constant, D is the volatile diffusivity, and t is the 

growth time, was shown to be important in the early stages 

of magma ascent but to trail off with the dwindling reserve 

of dissolved volatiles. Higher diffusivity leads to more rapid 

bubble growth. For a given diffusivity, instantaneous growth 

rates rise with increasing supersaturation because B is large 

in supersaturated melts. At first, this result may appear 

counterintuitive, but B is independent of D and is related only 

to the “yield” of volatile molecules vaporizing at the gas-melt 

interface. Interestingly, B buffers the kinetics of phase growth 

in Sparks’ model so that, all other things being equal, the final 

bubble size (and time-averaged growth rate) is independent of 

initial supersaturation. Decompressional growth—the simple 

expansion of a bubble due to a pressure drop—was determined 

to be most important in the later stages of magma ascent, with 

bubble size modeled as r(Po/P)

1/3 over decompression from 

pressure Po to P. 

The first comprehensive models for bubble nucleation in 

basaltic melts came a decade later, most notably in two papers 

by Bottinga and Javoy for submarine eruption of mid-oceanridge basalt (1990) and for subaerial eruption of Kilauean 

tholeiite (1991). Their work applied classical nucleation 

theory (for example, Frenkel, 1955), in which the free energy 

of formation (∆G(i)) of a bubble nucleus from a cluster 

containing i volatile molecules is given by the relation

 ∆G(i)= RTln(Cn/Ce)+A(i)s, (1)

Figure 5. Conceptualizations of magmatic degassing, then (A) and 

now (B). A, A visit to Kīlauea in 1909 led R.A. Daly to postulate that 

buoyancy-driven flow of C-N-Cl-S-H juvenile “fluid” to highest point 

in “feeding chamber” dictated position of eruption sites. Neighboring 

vents, robbed of juvenile gas by “the advantage of the middle vent,” 

are extinct. Daly reasoned that gas exsolution commenced at depths 

of several kilometers, while admitting that it was an assumption 

“necessarily difficult to test by the facts of field geology.” Modified from 

Daly (1911). B, Better understanding of Kīlauea’s plumbing system 

and, in 1980, start of systematic gas monitoring by the Hawaiian 

Volcano Observatory made a critical link between the geochemistry 

and depth of magma degassing and the dynamics of eruptions. Sketch 

shows passive CO2-rich degassing of parental magma at level of 

summit reservoir and shallow degassing of H2O-rich volatiles during 

rift-zone eruptions. Modified from Gerlach and Graber (1985). 
332 Characteristics of Hawaiian Volcanoes

where R is the gas constant, T is the temperature, A(i) is the 

surface area of a bubble nucleus at the pressure of interest, σ

is the surface tension, and Cn/Ce is the supersaturation ratio, 

where Cn is the concentration of dissolved volatiles at the 

onset of nucleation and Ce is the equilibrium concentration at 

ambient pressure. Creation of a stable bubble nucleus (∆G(i)<0) 

requires a level of supersaturation sufficient to offset the energy 

required to form a bubble wall. Consequently, higher degrees 

of supersaturation are needed at low pressure because of the 

larger surface area of incipient bubbles. In contrast to early 

degassing models, wherein nucleation was assumed to occur at 

small supersaturation (Sparks, 1978; Wilson and Head, 1981; 

Gerlach, 1986), the analysis by Bottinga and Javoy (1991) 

predicted significant “bubble-nucleation difficulties at low 

pressure,” requiring Cn/Ce ≥3 to trigger nucleation of H2O or 

CO2 bubbles at pressures below 20 MPa (or a depth of <800m, 

assuming a lithostatic pressure gradient). Once triggered, 

however, nucleation is rapid because of high supersaturation, 

with the nucleation rate, J, proportional to t-1exp(Cn/Ce), where 

t is the time needed to accrete a stable nucleus from collisions of 

randomly fluctuating volatile molecules.

In tandem with theoretical investigations, researchers were 

exploring the use of bubble-size distributions for empirical 

derivations of bubble nucleation and growth kinetics (for 

example, Toramaru, 1989, 1990; Sarda and Graham, 1990). Our 

own papers (Mangan and others, 1993; Cashman and Mangan, 

1994; Mangan and Cashman, 1996) specifically addressed 

shallow, H2O-rich degassing during subaerial eruptions along 

Kīlauea’s East Rift Zone. We adapted methods from studies 

of magmatic crystallization (Cashman and Marsh, 1988), with 

bubble-size measurements from rapidly quenched pyroclasts 

plotted as cumulative number distributions. On log-linear plots, 

the intercept at “size zero” gives the number density of bubble 

nuclei, and the slope determines the dominant bubble size of the 

population. If the time scale of vesiculation is known, timeaveraged bubble nucleation and growth rates can be calculated 

directly from the intercept and slope, respectively.

Using measured magma-ascent rates to infer vesiculation 

time scales, we determined that bubble-growth rates of 

~10-4 cm/s characterize a wide range of eruption intensities, from 

low-level effusion (mass eruption rates of ~103 kg/s) to highenergy lava fountaining (mass eruption rates of ~105

–106kg/s), 

with a dominant bubble size for both populations of ~ 0.01 cm. 

That time-averaged (not to be confused with instantaneous) 

bubble-growth rates and final bubble size are similar across 

the energy spectrum is consistent with the modulating effects 

of the rate constant B (Sparks, 1978). Nucleation rates, in 

contrast, increased by several orders of magnitude as a function 

of mass eruption rate: rates of ~10 events/cm3 per second were 

calculated for low-energy effusive eruptions, whereas rates 

of ~104 events/cm3 per second were obtained for high-energy 

lava fountains. Such extreme nucleation rates imply an intense 

vesiculation burst at high degrees of volatile supersaturation.

Collectively, the theoretical and empirical results summarized 

above highlight a largely overlooked conclusion by Verhoogen 

(1951, p. 729) about the controls on explosive volcanism:

Logarithm of mass eruption rate, in kilograms per second

Logarithm of bubble number density,

in bubbles per cubic centimeter

2 3 4 5 6 7 8 9 3

4

5

6

7

8

9

Etna

Kīlauea

Fontana

EXPLANATION

The most important single factor appears to be the 

number of bubbles which form per unit volume of 

time. The problem is similar to that of nucleation of 

crystals; and it is argued that differences in behavior of 

erupting volcanoes may depend more on the kinetics 

of the processes involved than on original differences 

in composition, gas content, depth, etc. 

A plot of bubble number densities (a proxy for nucleation 

rate) of pyroclasts from several Hawaiian-style eruptions as 

a function of mass eruption rate (fig. 6) illustrates the truth of 

Verhoogen’s statement.

Bubble Interactions

As the percentage of gas increases, the physics of bubble 

interactions becomes important. Chemical engineering studies on 

the behavior of polymer bubble suspensions (<74 percent porosity) 

and foams (≥74 percent porosity) suggest that bubbles begin to 

“feel” the effects of their neighbors when their separation distances 

are nearly equivalent to their radius (Ivanov and Dimitrov, 1988). 

In a homogenous distribution of similar-size bubbles, geometric 

considerations give the nearest-neighbor distance as bubble number 

density to the -1⁄3 power (Underwood, 1970). Bubble number 

densities characteristic of basaltic eruptions (102–107/cm3) suggest 

that interactions between bubbles as small as 10-3 to 10-1 cm cannot 

be neglected.

Figure 6. Plot of bubble number densities (BND) of rapidly 

quenched pyroclasts from Hawaiian-style eruptions, showing a positive 

correlation with mass eruption rates (MER), according to the expression 

log(BND)=0.73log(MER)+1.98 (correlation coefficient r2=0.85). Hawai‘i BNDs 

from Mangan and others (1993), Mangan and Cashman (1996), Stovall and 

others (2011, 2012), and Parcheta and others (2013), with MER calculated 

from data of Richter and Murata (1966), Swanson and others (1979), Heliker 

and others (2003), assuming a density of basaltic melt of 2,600 kg/m3; Etna 

data from Polacci and others (2006), Sable and others (2006), and Rust and 

Cashman (2011); Fontana data from Costantini and others (2009, 2010).
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 333

Of considerable relevance in basaltic systems is the 

physics of bubble coalescence, expansion, and ripening. 

Expansion dominates as bubbly melt approaches the Earth’s 

surface because of the large increase in gas volume at low 

pressure. At a fixed gas volume, both coalescence and ripening 

(redistribution of gas from small to large bubbles) increase 

the characteristic bubble size and decrease the bubble number 

density. Understanding the processes that change how a 

specific volume of gas is partitioned in the melt—whether 

as many tiny bubbles or as few, very large bubbles—bears 

significantly on the rheology of magma. 

Verhoogen’s (1951) study is one of the earliest on 

the physics of coalescence in basaltic bubble suspensions 

(<74percent porosity). He reasoned that bubbles rise 

buoyantly through a column of magma at rates proportional 

to their size, allowing larger, faster bubbles to overtake, 

collide with, and “swallow up” smaller, slower bubbles. Later 

modeling of coalescence in Hawaiian lavas defined a “linear 

collection efficiency” term given by the radius ratio of small 

to large bubbles (Sahagian, 1985; Sahagian and others, 1989). 

If this ratio is very small (large size difference), coalescence is 

unlikely because larger bubbles are swept past smaller bubbles 

without collision. Similarly, a ratio near unity (equal sizes) 

is also unlikely, because there is little chance that one bubble 

will overtake the other.

In circumstances where bubbles do collide, bubble 

size continues to be important. At least one of the colliding 

pair must deform to ingest the other. Small bubbles resist 

deformation by virtue of surface tension and high internal 

pressure, as is deducible from the Young-Laplace law, which 

gives internal bubble pressure (Pinternal) as a function of surface 

tension (σ), bubble radius (r), and ambient pressure (Pexternal): 

Pinternal = Pexternal+σ/r. (2)

Theoretical and experimental studies of basaltic systems 

(melt viscosity, 10–103 Pa-s) suggest that bubbles must be at 

least several millimeters in diameter before the wavelength of 

a disturbance (collision) creates significant deformation of the 

bubble wall (Manga and Stone, 1994; Suckale and others, 2010). 

The net result of collision-based models is that the 

smallest bubbles have the lowest probability of coalescing, 

as demonstrated empirically by the perturbations observed in 

bubble-size distributions of Kīlauea pyroclasts, which show 

the signature of bubble coalescence at sizes >0.1 cm (Mangan 

and others, 1993; Cashman and Mangan, 1994; Mangan and 

Cashman, 1996; Herd and Pinkerton, 1997; Stovall and others, 

2012; Parcheta and others, 2013) and in experimental studies 

using analog fluids (for example, Pioli and others, 2012).

In magmatic foams (≥74 percent porosity), bubble 

coalescence hinges on interstitial melt flow and typically 

involves ingestion of many bubbles simultaneously. Closest 

packing of uniform-size spherical bubbles (74 percent 

porosity) or deformed polyhedral bubbles (>74 percent 

porosity) provides a basic conceptual model for magmatic 

foams that is consistent with the structures observed in 

basaltic pyroclasts (Mangan and Cashman, 1996). The 

areas of contact between bubbles form lamellae, or films. 

The geometry of packing requires that three films meet at 

dihedral angles of 120°, which form channels, or plateau 

borders, that are trigonal in cross section. Four plateau 

borders intersect at ~109o to form tetrahedral vertices. Also 

important are geometric constraints on close packing of 

uniform polyhedra. Dana (1890) first described the geometry 

of basaltic foams, which he described as “thread-lace 

scoria” (now known as reticulite with ≥90 percent porosity) 

that formed in high lava fountains (Mangan and Cashman, 

1996). The regularity of the polyhedral cells in these clasts 

approaches that required by energy minimization (reduction 

of melt-gas interfacial area to gas volume), as shown by 

comparison of observed cell shapes with those predicted 

mathematically by Weaire and Phelan (1994).

Drawing insight from chemical engineering studies 

(for example, Hass and Johnson, 1967; Ramani and others, 

1993; Sonin and others, 1994), models of coalescence in 

basaltic foams assume that interstitial melt flows from bubble 

films (film thinning) to plateau borders, which, in turn, 

drain downward under the influence of gravity (Jaupart and 

Vergniolle, 1988; Toramaru, 1989; Proussevitch and others, 

1993; Mangan and Cashman, 1996). Mangan and Cashman 

(1996) equated the rate of film thinning, Vf, to the driving 

pressure for flow, DP: 

Vf = DP(δ3/3ηrc

2

), (3)

where δ is the film thickness, η is the melt viscosity, and rc

is the radius of the contact surface between bubbles. The 

formulation of DP may relate to the force of expansion 

(pressure exerted on melt in the interstices between expanding 

bubbles), capillary forces (interfacial pressure gradient 

due to bubble wall curvature), and (or) gravitational forces 

(downward drainage through melt channels). Measurements 

on rapidly quenched pyroclasts from Kīlauea eruptions 

suggest that basaltic melt films may become extremely thin 

(δ ~10-4 cm) before rupture and coalescence occur (Mangan 

and Cashman, 1996). In engineering applications, bubblewall thinning and rupture distinguish open-celled (high 

permeability) from closed-celled (low permeability) foams. In 

magmatic foams, the degree to which these processes occur, 

relative to quenching, distinguishes basaltic pumice (quasiclosed cell) from scoria and reticulite (quasi-open cell).

Gas partitioning in basaltic foams is also influenced by 

Ostwald ripening, a process first described by the German 

chemist Wilhelm Ostwald in 1896. Ripening involves diffusive 

transfer of gas from small to large bubbles due to pressure 

excess inside bubbles, which, from the Young-Laplace law, is 

high for small bubbles and low for large bubbles. Since bubbles 

in magmatic foams vary in size, gas will diffuse across bubble 

films from regions of high to low pressure. Large bubbles will 

grow, while small bubbles will shrink. Mangan and Cashman 

(1996) borrowed an expression for ripening time, t, from 

deVries (1972), in which bubble size changes as 

Δr= t(4RTDCeσ/ΔPδ), (4)
334 Characteristics of Hawaiian Volcanoes

where R is the gas constant, T is the temperature, D is the 

diffusivity, Ce is the equilibrium dissolved-gas concentration, 

σ is the surface tension, P is the external pressure, and δ is 

the film thickness. Results indicate that basaltic foams can 

ripen significantly over time scales of a few tens of seconds 

at δ≤ 0.005 cm. Even at porosities as low as 10–20 percent 

(δ ~0.01 cm), experimental studies have shown that significant 

ripening occurs, albeit on time scales of hours to tens of hours 

(Lautze and others, 2011).

Effect of Gas Partitioning on Magma Viscosity and 

Magma Flow Regimes

Attempts to determine the viscosity of molten basalt 

date back to the late 1800s and early 1900s in Hawai‘i, with 

hydrodynamic measurements in lava channels and lava lakes 

(for example, Becker, 1897; Palmer, 1927; Kinsley, 1931; 

Macdonald, 1954, 1955; Shaw and others, 1968). Most 

influential were the rotational viscometry measurements by 

Shaw and others (1968) in the 1965 Makaopuhi lava lake 

on Kīlauea’s East Rift Zone (Wright and Okumura, 1977). 

Combining the lava-lake data with those from laboratory 

experiments using bubbly silicon oils, Shaw (1972) was 

one of the first to recognize that bubbles variously influence 

magma viscosity. He observed a viscosity increase at low 

shear but a viscosity decrease at high shear that he attributed 

to bubble deformation (Shaw and others, 1968, p. 252): 

This behavior at high shear rates appears to be related 

to the distortion of the bubbles, which become drawn 

out into annular zones effectively decreasing the shear 

resistance. The cause of stiffness at low shear rates is 

complex, but qualitatively the surface tension of the 

bubbles opposes distortion. 

The effect of bubble size on viscosity was also 

recognized by Shimozuru (1978), who contrasted the behavior 

of large, deforming bubbles with that of small, spherical 

bubbles that act as rigid “obstacles” to increase shear viscosity. 

With insight from Hawaiian lava lakes, Shimozuru was also 

one of the first researchers to conceptualize the rheology 

of bubble+melt mixtures rising in eruptive conduits using 

engineering terms, specifying bubbly or slug flow regimes, 

depending on bubble size. The application of two-phase flow 

mechanics has since become entrenched in volcanology.

Bubbles and Viscosity 

Many theoretical and laboratory studies followed 

Shaw’s (1972) pivotal viscosity experiments in Makaopuhi 

lava lake (for example, Stein and Spera, 1992; Crisp and 

others, 1994; Manga and others 1998; Spera, 2000; Manga 

and Loewenberg, 2001; Rust and Manga, 2002; Pal, 2003; 

Llewellin and Manga, 2005; Harris and Allen, 2008). 

Crystals are well known to increase the shear viscosity of a 

magma+crystal suspension by an amount, 

 η(1− ) ø

øc

-øc

, (5)

where η is the melt viscosity, φ is the volume fraction of crystals, 

and φc is the volume fraction of crystals at maximum packing. For 

the near-liquidus, low-crystallinity magmas of most Hawaiianstyle eruptions, however, the shear viscosity of a melt+crystal 

suspension is less than a factor of 2 above that of the melt phase 

alone (1,200–1,150 °C and ≤20 percent crystallinity).

The impact of bubbles on shear viscosity, in contrast, 

is a complex interplay between bubble deformation, magma 

shear rate, and porosity (Manga and others, 1998; Manga and 

Loewenberg, 2001; Pal, 2003; Llewellin and Manga, 2005). A 

single dimensionless number—the capillary number, defined as

Ca = ηγr/σ, (6)

where η is the melt viscosity, γ is the shear rate, r is the bubble 

radius, and σ is the surface tension—represents the balance 

between two opposing forces: shear stress promoting bubble 

deformation and surface tension resisting bubble deformation 

(fig. 7). Pal (2003) showed that for Ca<<1, spherical bubbles 

have the same effect as rigid particles and increase shear 

viscosity, with the viscosity of the suspension, ηs, given by

ηs = η (1− ) ø

øm

-øm

| , (7)

where φ is the bubble volume fraction and φm is the volume 

fraction of spherical bubbles at maximum packing (~0.64 for 

bubbles of uniform size). For Ca>>1, deformed bubbles aligned 

in the direction of flow reduce the shear viscosity, with

ηs = η(1− ) ø

øm

1.67øm

| . (8)

Melt viscosity does not vary widely across the spectrum 

of Hawaiian-style eruptions, and so the product γr is a key 

determinant of Ca. 

Bubbles and Flow Regimes

Four two-phase flow regimes are generally used to describe 

the flow of magma-gas suspensions in eruptive conduits: bubbly 

flows, containing discrete, small gas bubbles; slug flows, in 

which large pockets of gas rise through the magma; annular 

flows, comprising a central gas jet sheathed in a coherent melt 

layer flowing along walls of the conduit; and dispersed flows, in 

which gas carries suspended melt droplets (fig. 8; see reviews by 

Jaupart, 2000, Houghton and Gonnerman, 2008). The behavior 

of bubbly flows is strongly controlled by buoyant and viscous 

forces. Laminar flow conditions are a good approximation, with 

the rate of flow decreasing near the conduit walls, owing to 

viscous drag (wall friction) and increasing inward to the medial 

axis. At the other end of the spectrum are dispersed flows, which 

are largely controlled by gas expansion and inertial forces, 

owing to the low viscosity of gas. Dispersed flows are turbulent, 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 335

0.01 0.1 1 10 100 1,000

Strain rate, in reciprocal seconds

0.00001

0.000001

0.0001

0.001

0.01

0.1

1

Bubble radius, in meters

Decrease in shear viscosity

Increase in shear viscosity

Effusion

Lava fountain

EXPLANATION

Ca < 0.64

Ca > 0.64

Ca = 0.64

A

0.01 0.1

0.1

1 10 100 1,000

10

1

Capillary number

Relative viscosity 

Ф = 0.1

Ф = 0.5

Ф = 0.3

0.001

Effusive eruptions bubble radius, 0.01–0.1 cm

Ca = 0.64

Lava fountains bubble radius, 0.01–0.1 cm

Suspension viscosity higher

than pure melt phase

Suspension viscosity lower

than pure melt phase

B

A B C D

with a constant mean velocity across the width of conduit, 

owing to low wall friction.

Between these two end members, flow behavior is 

complex, because buoyant, expansive, viscous, and inertial 

forces are all-important. In slug flows, melt moves around the 

slug at a rate controlled by viscous forces at the wall, rather 

than by interfacial forces at the melt-slug boundary. Mixed 

viscous and inertial forces complicate the ascent of the gas 

slug, with differential gas expansion causing the slug nose 

to accelerate, relative to the slug base (James and others, 

2008). A melt layer along the conduit walls is characteristic of 

annular flows, with the direction of flow either concurrent or 

countercurrent to the flow of gas, depending on the velocity of 

the gas phase and the viscosity of the liquid. Interfacial forces 

at the gas-melt boundary are consequential in this regime, 

creating instabilities in the melt film (waves) and making 

the melt rheology (laminar versus turbulent, steady versus 

unsteady) exceedingly difficult to ascribe. Melt droplets detach 

from unstable regions and become entrained in the gas phase. 

At a high proportion of droplets, the flow becomes dispersed. 

Bubbly flows and slug flows also transition into dispersed 

flows upon fragmentation.

Figure 7. Graphs illustrating how bubbles influence shear viscosity 

of basaltic magma during Hawaiian-style eruptions. A, Bubble radius 

(r) versus strain rate (γ) for which capillary number (Ca)=0.64. To 

right of Ca0.64 curve, shear viscosity decreases because of bubble 

deformation; left of the Ca0.64 curve, shear viscosity increases 

because bubbles, acting as rigid particles, resist deformation. 

Calculations assume a melt viscosity of 100 Pa-s and a melt-vapor 

surface tension of 0.1 J/m2 (Khirtov and others, 1979) for hydrous 

basaltic melt. Strain-rate estimates for low-energy effusive eruptions 

and high-energy lava fountains were calculated by using γ=2u/3R, 

where u is average flow velocity in a vertical, cylindrical conduit of 

radius R (for example, Mastin and Ghiorso, 2000). A range of strain 

rates was examined, assuming conduit radii of 0.5 to 2.5 m and flow 

velocities of 0.1 to 320 m/s, which were derived from eruptive fluxes 

reported for Pu‘u ‘Ō‘ō lava fountaining (Heliker and Mattox, 2003, 

table 1, episodes 20–39) and Kupaianaha effusive activity (1991 very 

low frequency [VLF] determinations in Sutton and others, 2003, table 

1). B, Effect of porosity (Φ) on relative shear viscosity of basaltic 

magma (viscosity ratio of melt+bubbles/pure melt). For Ca>0.64, 

bubbles decrease shear viscosity of bubble-melt suspensions, relative 

to that of pure melt; for Ca<0.64, bubbles increase shear viscosity. 

Plot is modified from Pal (2003) for conditions relevant to Hawaiianstyle eruptions. Black solid lines span capillary numbers for bubbly 

conduit flow, terminated by open circles (effusive eruptions) and 

filled circles (lava fountains). Strain rates were estimated, as in figure 

7A (effusive eruptions γ~0.03-3.0 s-1; lava fountains γ~2-425 s-1, 

assuming conduit radii of 0.5–2.5 m).

Figure 8. Diagram showing two-phase flow regimes proposed for 

mixtures of gas and low-viscosity silicate melt rising in a vertical 

conduit during Hawaiian-style eruptions. Black arrows within red 

fields (melt) indicate laminar flow of melt, with arrow length indicating 

relative velocities; black curved arrows in white fields (gas) indicate 

turbulent flow of a continuous gas phase. In natural volcanic systems, 

bubbly flows (A) may transition to slug flows (B) by way of bubble 

coalescence, or to dispersed flow (D) by way of fragmentation. Lava 

fountains have characteristics of dispersed flows. Annular flows (C) are 

attributed to large-scale, violent foam collapse. Shredding of bounding 

melt sheath may cause transition from annular to dispersed flows.
336 Characteristics of Hawaiian Volcanoes

Fragmentation of Gas-Charged Magma

Verhoogen (1951) was the first researcher to apply 

classical hydrodynamics to the fragmentation of basaltic 

magma, arguing that fragmentation results from violent 

“coalescence of a large number of bubbles expanding 

radially faster than they can rise and escape at the 

surface....” Later, McBirney (1963), McBirney and Murase 

(1970), and Sparks (1978) argued that coalescence-driven 

fragmentation is controlled by the geometric limits to 

bubble packing, which was assumed to be ~75 percent by 

analogy to maximum packing of uniform-size rigid spheres. 

Later, Mangan and Cashman (1996) pointed out that high 

vesicularity does not necessarily lead to fragmentation of 

bubbly flows, because some effusively erupted basalts have 

porosities of >75 percent. 

More recent experiments and theoretical treatments 

show that rapidly accelerating, low-viscosity magma may 

undergo shear-induced liquid breakup irrespective of a 

critical porosity (Mader and others, 1997; Papale and others, 

1998; Cashman and others, 2000; Namiki and Manga, 2008; 

Rust and Cashman, 2011). The experimental results of 

Namiki and Manga (2008) emphasize that fragmentation of 

vesiculating low-viscosity basaltic magma fundamentally 

differs from that of more viscous, silicic magma. In basaltic 

systems, inertial forces stretch and “pull apart” rapidly 

rising, expanding magma. Silicic fragmentation, in contrast, 

is characterized by brittle failure at high tensile stress. 

Mangan and Cashman (1996) argued qualitatively that 

an intense burst of bubble nucleation and growth at high 

volatile supersaturation could cause the accelerations and 

shear rates leading to liquid breakup during Hawaiian lava 

fountaining. Rust and Cashman (2011) have since found 

evidence for liquid breakup in basaltic lava fountains in the 

poor correlation between bubble size and pyroclast size. In 

Strombolian-style eruptions, which have been interpreted 

as reflecting breakage of individual large bubbles in the 

slug flow regime, liquid breakup is controlled by the nearsurface dynamics (James and others, 2008). As the vent 

is approached, the nose of the slug is expanding rapidly, 

relative to the slug base. The lag in momentum causes the 

slug base to rebound, inducing an upward-directed pressure 

transient that causes bursting at the nose.

Post-Fragmentation Evolution of Molten 

Pyroclasts

The physical evolution of a bubbly magma does not 

stop at fragmentation. Pyroclasts continue to expand in the 

conduit and fountain as they adjust to ambient pressure, 

and they may continue to do so upon deposition until the 

glass transition temperature is reached. Post-fragmentation 

coalescence and ripening also increase the mean bubble size 

and reduce the bubble number density. In their treatise on 

the features of basaltic rocks, Wentworth and Macdonald 

0 0.1 0.2 0.3 0.4 0.5 0.6 0.7

2

4

6

8

10

12

14

16

Bubble diameter, in centimeters

Natural logarithm of bubble number density,

in bubbles per cubic centimeter

Effusive eruption Lava fountaining

Scoria

Porosity 78%

Scoria

85%

Reticulite

97%

Vent lava

74%

2 mm

Scale

Scoria

Scoria

Reticulite

Vent lava

(1953) alluded to the continuous evolutionary nature of 

molten pyroclasts in their statement “In a genetic series 

reticulite lies between pumice and ash.” The evolution of 

magmatic foams is plotted in figure 9. Bubble expansion, 

coalescence, deformation, and ripening transform a 

suspension of multisize spherical bubbles into a polygonal 

foam with a preponderance of pentagonal and hexagonal 

faces, increasing bubble size by an order of magnitude (from 

0.01 to 0.10 cm), decreasing the bubble number density 

by three orders of magnitude (from 105 to 102 cm-3), and 

reducing the interfacial surface area per unit volume of melt 

by ~75 percent (55 to 13 cm-1).

Figure 9. Graph of natural logarithm of cumulative bubble number 

density as a function of bubble diameter for Kīlauea lava fountain 

pyroclasts and effusively erupted lava, with binary thin-section images 

of samples. On such plots, log-linear distributions are expected only for 

steady-state bubble nucleation and growth due to volatile exsolution 

and decompressional expansion. Data distribution is distorted from 

log linearity by coalescence (deflection upward at large size fractions; 

scoria and effusive lava) and ripening (concave downward; reticulite). 

Thin-section images illustrate continuous structural evolution from 

spherical foam (≥74 percent porosity) to polyhedral foam (>97 percent 

porosity). See Mangan and Cashman (1996) for details.
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 337

The preserved porosities of fountain pyroclasts reflect 

a competition between post-fragmentation expansion, 

permeability increase (interconnected bubbles), and 

outgassing (gas flowing out of the pyroclast; not to be 

confused with degassing, which involves volatile exsolution 

from melt into bubbles). Rust and Cashman (2011) 

demonstrated that the permeability of mafic pyroclasts 

increases by two orders of magnitude as porosity rises from 

45 to 70 percent. They suggested that once pyroclasts begin 

to outgas, further expansion is limited, “much like a hole in a 

balloon can make it impossible to further inflate.” Expansion 

dominates over outgassing for k/R2<ηgas/ηmelt, where R is 

the pyroclast radius (length scale for gas flow), k is the 

pyroclast permeability, ηgas is the viscosity of the gas, and 

ηmelt is the viscosity of the melt (fig. 10). In contrast to silicic 

pyroclasts, the low viscosity of near-liquidus basalt and the 

coarseness of fountain ejecta favor significant expansion. In 

further contrast, a condition of dynamic permeability exists 

in low-viscosity magma that is unlikely to exist at higher 

viscosity—that is, transient apertures open gas channels 

between bubbles, only to close again once the gas is released 

from the pyroclast and the fluid melt relaxes. Because of 

their fluidity, basaltic pyroclasts may oscillate across the 

boundary between outgassing- and expansion-dominated 

behaviors. Sporadic gas loss through dynamic permeability, 

which is basically gas-melt decoupling on short length and 

time scales (see subsection below entitled “Models Framed 

by H2O-Rich Degassing in the Conduit”), may explain 

the relatively wide range of vesicularities (Houghton and 

Wilson, 1989) and the regularity of bubble shapes (Moitra 

and others, 2013; Parcheta and others, 2013) that distinguish 

fountain pyroclasts from their more viscous counterparts.

0 5 10 15 20 25

Clast diameter, in centimeters

-13

-12

-11

-10

-9

-8

Logarithm of permeability, in square meters

Clast outgassing

Clast expansion

100 Pa-s

1,000 Pa-s

Numerical Modeling of Hawaiian-Style 

Eruptions

The Observatory will measure everything for a 

century to come.

 —Thomas A. Jaggar, Jr. (1945)

For uniform conduit geometry, four intertwined and changing 

parameters influence eruptive behavior: driving pressure, melt 

viscosity, crystal content, and gas content. This interrelation is seen 

most simply in the proportionality Q∝ΔP/ηf, where Q is the mass 

flux, ηf is the bulk magma viscosity (melt+crystals+bubbles), 

and ΔP is the pressure driving the flow (buoyancy+reservoir 

overpressure–weight of the overlying magma column). In nature, 

though, conduits narrow, kink, or flare; reservoir overpressure 

waxes and wanes; and degassing with or without crystallization 

causes changes in magma viscosity, buoyancy, and mass. In 

light of these complexities, Jaggar’s call for the observatory 

to “measure everything” shows much foresight. Direct field 

measurements are essential to choosing appropriate model input, 

constraining boundary conditions, and evaluating model results. 

Not coincidentally, the first numerical models of basaltic eruption 

dynamics drew heavily from the data archives of the Hawaiian 

Volcano Observatory.

Overview of Existing Eruption Models

The first computer-assisted numerical models of Hawaiianstyle eruptions employing classical conservation equations for 

mass, momentum, and energy were introduced in the 1980s. 

The earliest lava-fountain model, constructed by Wilson and 

Head (1981), considers magma ascending and degassing as 

bubbly flows (fig. 8A), with a transition to dispersed flows 

(fig. 8D) once the porosity reaches 75 percent. Shallow, H2Orich degassing is the driver in this model. CO2, which is both 

less abundant and less soluble, is assumed to have exsolved 

and escaped from the system at great depth (see for example, 

Gerlach and Graber, 1985).

Vergniolle and Jaupart (1986) provided an alternative 

lava-fountain model, with CO2-rich degassing as the driving 

force. In their model, deeply exsolved CO2 bubbles are not 

lost to the system; instead, they accumulate as a foam layer 

at the roof of the subvolcanic reservoir. Lava fountains result 

from violent annular flows (fig. 8C) triggered by spontaneous, 

wholesale collapse of the foam, once a critical thickness is 

reached. Secondary H2O-rich degassing accompanies the 

depressurization associated with collapse but is assumed to be 

inconsequential to the physics.

Extensions and refinements to both models, based largely on 

data and observations from the 1983–86 Pu‘u ‘Ō‘ō fountaining 

episodes, were presented by Parfitt and Wilson (1994, 1995, 

1999), Parfitt and others (1995), Parfitt (2004), Jaupart and 

Vergniolle (1988), Vergniolle and Jaupart (1990), Vergniolle 

(1996, 2008), Woods and Cardoso (1997), and Seyfried and 

Figure 10. Plot showing conditions for pyroclast outgassing (field 

above each curve) versus pyroclast expansion (field below each 

curve), calculated for melt viscosities (ηmelt) of 100 and 1,000 Pa-s. 

Curves obtained from relation k/R2=ηgas/ηmelt. Outgassing is favored by 

high permeability (k), low melt viscosity, and small pyroclast size (R). 

Calculations assume gas viscosity (ηgas) of 10-5 Pa-s. Modified from 

Rust and Cashman (2011). 
338 Characteristics of Hawaiian Volcanoes

Freundt (2000). These later treatments focused on constructing 

a continuum model in which transitions between persistent lavalake activity, low-energy effusive eruptions, and high-energy 

lava fountains are modulated by styles and rates of gas release.

Models Framed by H2O-Rich Degassing in the 

Conduit

The underlying assumption in the H2O-rich degassing 

model is that bubble-free magma rises steadily in the conduit until 

reaching the pressure-depth of H2O saturation. Bubbles then begin 

to nucleate and grow by volatile exsolution, decompressional 

expansion, and, depending on the rise rate of magma, bubble 

coalescence. Before reaching the saturation depth, magma 

rise is controlled by reservoir overpressure. Once vesiculation 

commences, however, the evolution of H2O-rich gas is what 

propels magma upward to the surface. 

Whether or not the magma is fragmented before reaching the 

vent depends on how strongly bubbles are coupled to the parcel 

of melt from which they exsolve. Under Stokes’s law, the rise rate 

of a single bubble is determined by the ratio of bubble buoyancy 

to viscous drag. At high magma rise rates, bubbles are carried 

along in the flow, with little independent motion or interaction 

(limited coalescence or ripening). Wilson and Head (1981) and, 

later, Parfitt and Wilson (1995) showed that for the melt viscosities, 

10

0.1

1

0.001

0.001 0.01 0.1 1 10

0.01

Magma rise rate, in meters per second

Bubble diameter, in meters

 Bubble rise/magma rise =1

Closed-system 

degassing

Hawaiian-style 

lava fountaining

Open-system 

degassing

Effusive 

eruption

A

B

C

Open-system degassing 

Strombolian-style spattering

dissolved H2O contents, and H2O saturation pressures typical of 

Hawaiian tholeiite (10–103Pa-s, 0.2–0.4 weight percent, and <10 

MPa, respectively), magma rise rates of >>0.1 m/s result in strong 

coupling between bubbles and “parent” melt in the mode of closedsystem degassing (fig. 11, location A). The ascending gas-charged 

magma eventually reaches sufficient porosity (~75 percent) to 

trigger fragmentation (original model of Wilson and Head, 1981) 

or, as more recent models convincingly demonstrate, shear-induced 

liquid breakup (Namiki and Manga, 2008; Rust and Cashman, 

2011). The sharp pressure drop and reduced wall friction instigated 

by fragmentation causes pronounced acceleration of the dispersed 

flow, which eventually emerges from the vent as a jet of molten 

pyroclasts and gas. Fountaining ceases when the reservoir pressure 

returns to a stable state. For quasi-uniform magma supply from 

depth, waxing and waning reservoir overpressure produces episodic 

fountains that are regulated by a magmatic “valve” of cooler, highyield-strength magma that clogs narrow segments of the feeder dike 

between successive events. 

At the other end of the spectrum, magma rise rates of 

<<0.1m/s create conditions for open-system degassing with minor 

eruption of magma (fig. 11C). When the magma rise rate is very 

low, larger bubbles gradually decouple from their parent melt. 

With differential rise of varying-size bubbles, coalescence can 

occur. Wilson and Head (1981) suggested that “a runaway situation 

may eventually develop” in which fast-rising, large bubbles beget 

faster, larger bubbles until slug flows prevail. Conduit-filling gas 

slugs burst at the vent, generating Strombolian-style spattering of 

molten pyroclasts. Slug flows formed by cascading coalescence 

were observed in large-scale two-phase flow experiments, using 

analog fluids in a long vertical tube (6.5×0.25 m) scaled to match 

persistent degassing from a standing column of basaltic magma 

(Pioli and others, 2012). Though not directly analogous because no 

bubbles nucleate in the column (gas fluxes from below only), the 

experiments and related theoretical modeling show that the process 

of collision-induced coalescence, slug formation, and quasi-steady 

bursting of slugs at the surface creates circulation cells within 

the conduit that homogenize and stabilize the magma column. 

This type of pulsed slug degassing and magma recirculation with 

negligible discharge of lava is suggested as a plausible mechanism 

for the rise-fall cycles observed in active Hawaiian lava lakes 

(Ferrazzini and others, 1991; Parfitt and Wilson, 1994, 2008; 

Parfitt, 2004; Johnson and others, 2005). 

Effusive Hawaiian eruptions lie between the end members 

above. The continuous extrusion of bubbly lava results from 

intermediate magma rise rates of ~0.1 m/s (fig.11, location B). 

Here, although open-system gas loss occurs because of partial 

decoupling of bubble and melt, bubbly flow is maintained over the 

entire length of the conduit. A fraction of larger, coalesced bubbles 

rising through, rather than with, the magma during effusive 

eruptions may accumulate as a foam layer at the top of an effusive 

vent that has skinned over as the result of radiative cooling. 

Recent field studies at the active Pu‘u ‘Ō‘ō and Halema‘uma‘u 

vents demonstrate convincingly that bubbles collecting under 

a viscoelastic lid at the top of the conduit can lead to sporadic 

gas-release events and episodic rise-fall cycles without slug flow 

(Patrick and others, 2011a,b; Orr and Rea, 2012). 

Figure 11. Plot showing curve for maximum bubble diameter achievable 

as a function of magma rise rate for H2O-rich degassing during flow in 

a vertical conduit. Model assumes melt viscosity of 100 Pa-s and initial 

dissolved H2O content of 0.5 weight percent. Bubbles are assumed to grow 

by mass transfer (exsolution), decompressional expansion, and, for magma 

rise rates <0.1 m/s, coalescence due to bubble collision. For magma rise 

rates of >0.1 m/s (A), Stokes rise rate of bubbles is less than magma rise 

rate, and closed system degassing leads to lava fountaining (bubbles and 

melt coupled; no coalescence; not depicted in A, bubbly flows transition 

to dispersed flows near the surface). When Stokes rise rate of bubbles 

exceeds that of magma (B,C), open-system degassing leads to effusive 

eruptions (modest numbers of bubbles coalesce; bubbly flows maintained, 

B) or pulsed, Strombolian-style spattering (significant numbers of bubbles 

coalesce resulting in slug flow, C). Modified from Parfitt and Wilson (1995). 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 339

100

1

10

0.01

10-6 10-5 10-4 10 10-3 -2

0.1

Bubble diameter, in meters

Critical foam thickness, in meters

Reservoir roof Reservoir roof

Stable foam layer Unstable foam layer

Size range of 

bubbles in

 subvolcanic

reservoirs

Models Framed by CO2-Rich Degassing in the 

Subvolcanic Reservoir

Fountaining has been suggested to be triggered by sudden 

release of CO2-rich bubbles trapped at a reservoir roof. Here, 

the roof-conduit connection is assumed to lie at a pressuredepth interval below that of H2O saturation. The foam layer 

is generated either internally, by bubbles rising through a 

static, gas-saturated reservoir, or externally, by quasi-steady 

supply of bubbly magma to the reservoir from the mantle. 

With quasi-steady magma supply, magma throughput from 

reservoir roof to conduit base must be sufficiently impaired 

by drastic narrowing at the roof-conduit connection, such that 

bubbles segregate at the roof with only a little magma leaking 

into the conduit. In either case, an increasing buoyant force is 

exerted on bubbles pressed against the roof by those arriving 

from below. Bubbles flatten until the force of surface tension 

is insufficient to maintain the integrity of bubble walls. The 

critical thickness, hc, leading to collapse of the foam layer is 

modeled as

hc = (2σ/φgρmgr), (9)

where σ is the surface tension, φg is the gas-volume fraction 

in the foam (by definition, ≥0.74), ρm is the melt density, g is 

the gravitational constant, and r is the bubble radius (fig.12). 

The minimum size of bubbles accumulating at a roof at 

ambient pressures of ~10 to 100 MPa will range from 0.1 to 

1.0 mm in diameter (assuming no coalescence). We derive this 

constraint from the sizes of CO2 bubbles measured in basalts 

erupted along the Mid-Atlantic Ridge (Sarda and Graham, 

1990) corrected to the relevant pressure range and from 

decompression experiments using basalts saturated with mixed 

H2O-CO2 fluids at 400 MPa (Mangan and others, 2006). 

Thus, as shown in figure 12, the critical thickness of the foam 

layer (φg ≥0.74) ranges from ~0.1 to 1 m over the spectrum of 

expected bubble sizes. 

For a given conduit dimension, the vigor of the 

ensuing eruptive activity depends on the volume of CO2 gas 

spontaneously released by the collapsing foam (given by 

hcS, where S is the surface area of the roof). If the volume of 

CO2-rich gas expelled is large (~106 m3 for lava fountains), 

the eruption is assumed to be driven by annular flow, with the 

onset of H2O-rich degassing only a secondary contribution 

to upward acceleration. The turbulent core of expanding gas 

streaming through the conduit shreds the bounding magma 

layer at the walls to produce the dispersed jet of molten 

pyroclasts associated with fountaining. Once the volume of 

trapped foam is expended, fountaining ends. If the flux of 

bubbles from below is continuous, the cycle begins again with 

the foam layer building and failing at time scales controlled by 

the gas-flux rate.

At the other end of the spectrum, effusive eruptions 

represent circumstances of low gas flux to the roof and (or) 

a roof-conduit connection that does not appreciably hinder 

throughput of magma to the conduit. The critical foam 

thickness is never attained, and a steady stream of bubbly 

magma is erupted without fragmentation. Transitional between 

end members are those eruptions driven by buildup and 

spontaneous collapse of small foam volumes. Slugs of CO2-

rich gas cause activity ranging from passive degassing to mild, 

Strombolian-style spattering of molten pyroclasts.

Dynamic Considerations for Historical Eruptions 

Based on Reservoir Geometry and Gas Chemistry

Fundamental to the two lava-fountain models described 

above are the assumed compositions of the gases (CO2 versus 

H2O) propelling magma to the surface and, as a result, the depth 

(subvolcanic reservoir versus conduit) at which fragmentation 

occurs. Gas-emission data (composition and volume) and 

probable reservoir dimensions for Kīlauea’s historical lava 

fountains fit most neatly into the H2O bubbly-flow model. 

For the Pu‘u ‘Ō‘ō fountaining era, for example, scaling 

the foam-collapse model to satisfy both measured gas 

volumes (106 m3 of gas erupted per episode; Greenland, 

1988) and surface area of the subvolcanic reservoir roof 

(~400 m2; Dvorak and others, 1986; Wilson and Head, 1988; 

Hoffman and others, 1990) is difficult to reconcile with model 

predictions of critical foam thickness (fig. 12). On a smaller 

scale and at a shallower level, however, the critical foam 

thickness predicted by the model shows parity with field data 

Figure 12. Plot of critical foam thickness versus bubble diameter. 

Calculation uses formulation of Vergniolle and Jaupart (1990) for 

critical thickness (hc=2σ/φgρmgr), with surface tension σ=0.1 kg/s2, gas 

fraction of bubbles in foam layer φg=0.74, bubble radius r, and melt 

density ρm=2,600 kg/m3. In absence of coalescence, expected size 

of bubbles at subvolcanic reservoir pressures (0.1–1 mm; ~10 MPa) 

suggests that foam layers ~1 m thick are unstable. In Hawai‘i, a typical 

lava-fountaining episode releases ~106 m3 of gas (Greenland, 1988), 

suggesting unrealistically large reservoir roof areas (~106 m2) for historical 

eruptions of Kīlauea. We note, however, that formulation for foam stability 

is consistent with rise-fall cycles documented by Patrick and others 

(2011b) and Orr and Rea (2012), in which a surface foam layer collapses, 

releasing 10 to 104 m3 of H2O-rich gas amid mild spattering.
340 Characteristics of Hawaiian Volcanoes

obtained during recent rise-fall cycles at Pu‘u ‘Ō‘ō, albeit 

for H2O-rich foam (Patrick and others, 2011b; Orr and Rea, 

2012). Here, measured “roof” (viscoelastic skin of lava at the 

surface) dimensions and volumes of gas release are consistent 

with near-surface collapse of a foam layer ~1 m thick (fig. 12).

Also problematic for the CO2 foam-collapse model are 

the gas-emission data from Pu‘u ‘Ō‘ō fountaining episodes, 

which suggest low CO2 concentrations (CO2/S = 0.2; Gerlach, 

1986; Greenland and others, 1988). We note, however, that 

CO2-rich degassing could be important if parental magma 

bypasses Kīlauea’s summit reservoir or shoots through it 

without substantial residence in the shallow crust. Though 

not requiring foam collapse, the high-fountaining episodes of 

the 1959 Kīlauea Iki eruption are candidates for syneruptive 

mixed H2O-CO2 degassing. Seismic, tilt, and petrologic data 

all suggest that rapidly ascending parental magma bypassed the 

main summit reservoir, intruding into a small, separate storage 

compartment north of Halema‘uma‘u, less than 2 months 

before the onset of eruption (Wright and Fiske, 1971; Helz, 

1987). Helz (1987) observed that “the 1959 lava was hotter and 

more gas-rich than typical Kilauea summit lava,” and although 

no measurements were made to corroborate gas compositions, 

the extreme H2O supersaturation implied by the high bubble 

number densities in Kīlauea Iki scoria led Stovall and others 

(2011) to propose that “upward forcing” by expanding, 

previously exsolved CO2 bubbles led to extreme acceleration 

and rampant secondary nucleation of H2O-rich bubbles. 

With application of new field-based spectroscopic 

techniques, instances of eruptive activity driven by mixed 

H2O-CO2 degassing are being documented in basaltic systems. 

Open-path Fourier transform infrared spectrometry (OP-FTIR) 

measurements at the peak of fountaining on Mount Etna 

(Italy) give CO2/S ratios of ~10, some 2 to 4 times greater 

than the time-averaged CO2/S ratios (Allard and others, 2005). 

OP-FTIR measurements made during sporadic degassing 

bursts at Pu‘u ‘Ō‘ō in 2004–05 yield CO2/S weight fractions 

of ~18 (Edmonds and Gerlach, 2007), close to an order of 

magnitude higher than those reported by Greenland (1988) 

for Pu‘u ‘Ō‘ō fountaining episodes. We therefore conclude 

that vigorous fountaining episodes at Kīlauea can be driven 

by a combination of CO2 and H2O degassing, although the 

dynamics of mixed-gas eruptions remains poorly understood.

Eruption-Column Models

The structure of a Hawaiian-style eruption column can 

be diagrammed dynamically as a lower ballistic region of 

centimeter- to meter-size molten clots dispersed in a gas 

jet, with an upper convective plume of gas, fine ash, sulfur 

particles, and aerosols (Head and Wilson, 1989; Sparks 

and others, 1997). At Kīlauea, typical high-fountaining 

episodes emit lava and gas at rates of 105–106 kg/s and 

103 kg/s, respectively. The ratio of gas to pyroclast is on 

the order of 70:1 by volume. Direct observations reveal 

that molten pyroclasts are concentrated in the hot, central 

axis of the fountain. The thermal core grades outward to 

a fiery orange-red region of slightly lower temperature 

and clot concentration, and then to a sparse black halo of 

quenched pyroclasts “wafted high into the air by the hot 

turbulently rising fume cloud” (Richter and others, 1970). The 

temperatures characteristic of fountains can be estimated from 

olivine-glass geothermometry. Helz and Hearn (1998) found 

maximum quenching temperatures of 1,160–1,190 °C across 

a large sample suite comprising Kīlauea Iki and Pu‘u ‘Ō‘ō 

fountain pyroclasts, with multiple samples from within single 

fountaining episodes ranging ± ≤10 °C.

Dynamics of the Ballistic Fountain

Molten clots exit the vent at steep angles and reach a 

height, hf, approximated by the ballistic equation of motion 

 hf= uc

2

/2g, (10)

where uc is the exit velocity of pyroclasts and g is the 

gravitational acceleration. Early models (for example, Walker 

and others, 1971; Wilson and Head, 1981; Head and Wilson, 

1987) assumed conditions of dynamic equilibrium, with a clot 

velocity less than the gas velocity given by the quantity

 uc= ug–U, (11)

where U is the terminal velocity of the clot and ug is 

the velocity of the gas. These equations were combined 

to use fountain height as an indicator of the amount of 

dissolved volatiles originally stored in the magma. It was 

later demonstrated, however, that such calculations can be 

misleading, because substantial kinetic energy is lost by 

reentrainment of previously erupted, higher density clots that 

have fallen from the fountain and pooled around the vent 

(Parfitt and others, 1995; Wilson and others, 1995). Lava 

recycling at a central vent diminishes the mean clot velocity, 

uf, according to

uf=(ucM)/(M+Mi), (12)

where M is the mass flux of fresh lava from the vent, Mi is 

the radial inflow of recycled, lower density lava coalesced 

around the vent, and uc is, as above, the clot exit velocity under 

conditions of no reentrainment. The energy loss due to recycling 

can be substantial; for typical mass eruption rates and dissolved 

H2O contents, a 1-m-thick pond will decrease fountain height by 

as much as 35 percent (Wilson and others, 1995).

Another dynamic complexity influencing fountain height 

was tackled by Parfitt (1998), Wilson (1999), and Parfitt 

and Wilson (1999). Whereas earlier models assumed that 

clots were all of uniform diameter and density, Parfitt and 

Wilson considered the full range of clots across all sizes (dc) 

and densities (ρc). Unlike the finer, more uniform pyroclasts 

ejected in silicic eruptions, Parfitt and Wilson determined 

that, because they are coarser, most basaltic pyroclasts never 

reach dynamic equilibrium (uc≠ug–U). Also, by assuming 

dynamic equilibrium, earlier models apportioned too much of 

the internal energy liberated by gas expansion to the kinetic 

energy of clots, thus leading to an underdetermination of 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 341

Buoyant

plume

Ballistic fountain

Humidity 100%

Humidity 50% 2

4

6

8

10

12

0

2 3 4 5 6 7

Gas exit velocity ~ 500 m/s

-10 -9 -8 -7 -6 -5

50

100

150

200

250

300

0 20 40 60 80 100

0

500

1,000

1,500

400

100

200

300

0

0 0.2 0.4 0.6 0.8 1.0

Fountain

Plume

dc ρc ~690 kg/m2

dc ρc ~30 kg/m2

C

B

D

A

Logarithm of mass eruption rate, in kilometers per second

Mass of lava in eruption column, in percent

Distance from vent, in kilometers

Clast size, in phi units

Clast exit velocity,

in meters per second

Clast size x density,

in kilogram per square meter

Plume height, in kilometers Height above vent, in meters

the gas exit velocity by as much as 300 percent. Accounting 

for the full clast distribution measured in 1959 Kīlauea Iki 

deposits by Parfitt (1998), Wilson (1999) calculated pyroclast 

exit velocities ranging from 80 to 250 m/s and a gas exit 

velocity of ~500 m/s, contrasting, respectively, with the 100 

and 125 m/s he had obtained previously, assuming uniform 

pyroclasts in dynamic equilibrium with the gas (fig. 13A).

The Kīlauea Iki case study also showed that >95 percent 

of the eruptive mass is contained within the ballistic region of 

the eruption column, with clasts having dcρc≥100 kg/m2 falling 

within 200 m of the vent (figs. 13B,C; Parfitt and Wilson, 

1999). Inspection of the data tables of Wilson (1999) suggests 

that only grains with dcρc<0.5 kg/m2 are likely to be coupled 

to the gas and, thus, carried aloft in the buoyant plume rising 

above the fountain (for example, dc= 0.84 mm, ρc= 600 kg/m3; 

Wilson, 1999, table 1).

Dynamics of a Buoyant Plume

The convective plume of ash, gas, sulfate particles, and 

sulfuric acid aerosols that billows above a Hawaiian lava fountain 

offers a pale comparison to the stratospheric phenomena that are 

silicic ash columns. Most plumes from Hawaiian-style eruptions 

do not intrude into the tropopause, regardless of latitude or season. 

Surprisingly, few data exist on the heights of plumes from basaltic 

eruptions. The 1984 Mauna Loa eruption was one of the few 

eruptions for which direct measurements of plume height were 

made. At 7 km above the vent (11 km absolute altitude), the plume 

top was well below the tropopause, which was at 18-km altitude at 

the time of the eruption (Lockwood and others, 1984).

Although basaltic eruptions are certainly hotter than silicic 

ones, much of the thermal energy needed for plume buoyancy is 

locked in the large molten clots that fall rapidly out of the fountain. 

Calculations of the time required for thermal waves to travel from 

a clot interior to the gas flowing rapidly by it indicate that clots 

larger than 1 cm across probably have residence times too short 

for substantial heat loss during their upward trajectory (Sparks and 

Wilson, 1976; Wilson and others, 1978; Woods and Bursik, 1991).

Figure 13. Diagram and graphs illustrating features of basaltic eruption columns. Ballistic region of the fountain contains a dispersion of 

centimeter- to meter-size molten clots rising tens to hundreds of meters at speeds less than that of gas. Above the fountain rises a buoyant plume 

of magmatic gas and aerosols, entrained air, and modest proportions of very fine ash. Unlike silicic plumes, which reach stratospheric heights, 

basaltic plumes are not expected to penetrate the tropopause. A, Clast exit velocity versus clast size (from Wilson, 1999), based on data from the 

1959 Kīlauea Iki eruption. Modeled gas exit velocity of ~500 m/s exceeds that of clasts for all sizes shown. B, Mass percentage of clasts in eruption 

column versus height above vent, based on modeling of the Kīlauea Iki eruption by Parfitt and Wilson (1999). C, Density-modulated clast size (clast 

diameter, dc, multiplied by clast density, ρc) versus lateral distance from the vent, based on modeling of the Kīlauea Iki eruption by Parfitt and Wilson 

(1999). D, Plume height versus logarithm of mass eruption rate at two different humidities (from Woods, 1993, and Sparks and others, 1997).
342 Characteristics of Hawaiian Volcanoes

Basic plume theory gives the height, hp, of a convecting 

plume as

 hp=8.2Q1/4,

 (13)

where Q is the rate of thermal energy release as modulated 

by atmospheric conditions, the specific heats of magma 

and air, initial magma temperature, temperature decrease 

in the fountain due to cooling of pyroclasts, volumetric 

eruption rate, and the mass fractions of volatiles and fine 

lava fragments (see, for example, Sparks and Wilson, 1976; 

Wilson and others, 1978; Stothers and others, 1986; Woods, 

1993; Sparks and others, 1997). 

The flux of hot, juvenile ash adds significant buoyancy 

to silicic plumes but is a comparatively meager heat source 

for plumes above Hawaiian lava fountains, because ash is 

<5 percent of the total mass erupted (for example, Parfitt 

1998). Likewise, the flux of hot magmatic volatiles released 

during fountaining episodes is “too small to have significant 

influence on column [plume] behavior” (Sparks and others, 

1997). Basaltic plume models show that the amount of 

atmospheric water vapor incorporated into the fountain 

actually exceeds that of juvenile H2O (Woods, 1993; Sparks 

and others, 1997). Entrainment of ambient water vapor adds 

significantly to buoyancy because, as the plume ascends 

into cooler atmospheric layers, condensation releases latent 

heat and decreases plume density (fig. 13D). The mass of 

entrained water vapor, ma, is given by

ma = m(Cm∆T)/(Ca(To–∆T)), (14)

where Cm and Ca are the specific heats of magma and air, 

respectively, m is the eruptive mass flux, To is the initial 

eruption temperature, and ∆T is the temperature drop in 

the fountain. In tropical regions, where humidity is high, 

the added buoyancy can be substantial. Sparks and others 

(1997) calculated that the proportion of ambient water 

vapor entrained by basaltic lava fountains can be as much as 

1–2 percent of the total mass in the fountain and an order of 

magnitude greater than the proportion of magmatic H2O in 

the plume.

A “moist air” plume model (75–100 percent humidity) 

constructed by Woods (1993) and Sparks and others (1997) 

successfully reproduced the 7-km-high plume above the 

1984 Mauna Loa vent by assuming ash fractions of 1–10 

percent and using the known mass eruption rates (106 kg/s) 

and eruption temperatures (1,130 °C) of Lockwood and 

others (1984). A similar but “dry air” model (0 percent 

humidity) for the Mauna Loa plume constructed by Stothers 

and others (1986) underpredicted the Mauna Loa plume 

height by ~20 percent. Comparatively, entrained water vapor 

has the greatest effect on plume height at low mass eruption 

rates. At higher mass eruption rates, increasing proportions 

of cooler, less humid air from higher in the atmosphere are 

entrained (fig. 13D).

Dispersal Patterns

Airborne quenching of high-porosity pyroclasts is 

attributed to rapid, wholesale outgassing with simultaneous 

creation of an air-permeable network (for example, 

Wentworth and Macdonald, 1953; Mangan and Cashman, 

1996; Namiki and Manga, 2008; Rust and Cashman, 2011). 

Clots too large for wholesale outgassing will continue to 

evolve upon deposition, their fate dictated by pyroclast 

temperature and local accumulation rates. Head and Wilson 

(1989) diagrammed post-deposition pyroclast evolution in 

the context of vent construction, delineating lava ponds, 

rootless lava flows, welded spatter, and loose scoria on the 

basis of pyroclast landing temperature and accumulation 

rate (fig. 14). Augmenting the diagrammatic approach of 

Head and Wilson with quantitative data on glass transition 

temperature for rapidly quenched basaltic melts (Potuzak and 

others, 2008), pyroclast quenching temperatures (Helz and 

Hearn, 1998), measured accumulation rates (Swanson and 

others, 1979; Heliker and others, 2003), and stratigraphically 

constrained componentry (Heliker and others, 2003) reveal 

the physical and thermal complexity of basaltic cones that 

form during episodic high fountaining from a central vent. 

Cold Warm Hot

Moderate Slow Fast

Pyroclast accumulation rate, in meters per second

Pyroclast landing temperature, in degrees Celsius

10-3

10-5

750 1,125 1,170

Brittle scoria

320 kg/m3

Plastic scoria Fluid “plops”

Partially welded spatter

Completely welded spatter

1,500 kg/m3

Rootless lava flows

Lava lakes 

Figure 14. Plot of pyroclast accumulation rate versus landing 

temperature, illustrating the types of vent deposits produced during lava 

fountaining from a central vent. Values along the axes, which are not 

to scale, show the ranges in pyroclast landing temperature (olivineglass geothermometry by Helz and Hearn, 1998, and basaltic glass 

calorimetry by Potuzak and others, 2008) and pyroclast accumulation 

rate (field measurements by Heliker and others, 2003). Vent deposits 

transition from loose accumulations of brittle scoria (bulk density 

~320 kg/m3) to mounds of completely welded spatter (bulk density 

~1,500 kg/m3) with increasing temperature and accumulation rate. At 

the highest values, falling pyroclasts coalesce upon landing, forming 

fluid lava lakes and (or) rootless lava flows. Plot modified from Head 

and Wilson (1989). Density data from Heliker and others (2003). 
The Dynamics of Hawaiian-Style Eruptions: A Century of Study 343

Pyroclast landing temperatures may range from <750 °C 

(glass-transition temperature) to >1,170°C (eruption 

temperature), and pyroclast-accumulation rates may range 

from <10-5 to >10-3 m/s, depending on mass eruption rate, 

fountain height, and windspeed. Steeper slopes characterize 

the downwind direction of the cone where deposits of 

unconsolidated scoria and welded spatter abound. 

Heliker and others (2003) emphasize that the “cinderand-spatter-cone[s]” created by high lava fountains are 

distinct from the cinder cones that result from eruptions 

of less fluid basalt (for example, during Strombolian- or 

Vulcanian-style eruptions). They described the 255-m-high 

Pu‘u ‘Ō‘ō cone as a “striking landform . . . composed of 

cinder, agglutinated spatter, and lava flows.” After 3 years 

and 44 high-fountain episodes, densely welded spatter and 

rootless lava flows composed more than three-quarters of 

its volume.

Though visually striking, the growing Pu‘u ‘Ō‘ō 

cone represented only about 20 percent of the volume 

of lava emitted during fountaining episodes. Beyond 

the cone base, lava flows dominated, with distal ash fall 

contributing ≤2 percent of the eruptive volume. Typically, 

the sheetlike deposits of frothy pyroclasts and ash abruptly 

thin outward from the cone, generally diminishing to 10 

percent of the maximum near-vent thickness over dispersal 

areas of <10 km2. 

A Look Toward the Future

The predictability of the eruption process will 

depend on how well we can constrain the model 

parameters using the monitoring data. 

 —Keiiti Aki and Valérie Ferrazzini (2000)

Pursuit of Jaggar’s (1917) vision to “protect life and 

property on the basis of sound scientific achievement” in the 

21st century calls for a new generation of eruption models 

with strength in hazard prediction. These models must be 

sufficiently sophisticated, multidisciplinary, and realistic 

to answer the vital questions: When and where will an 

eruption occur? What hazards are expected? How far afield 

will their impact be felt? How long will the hazard persist? 

Models must be adaptable to cover changing conditions and 

assess possible outcomes. Field-based studies must advance 

in tandem with model development, because real-world 

observations form the conceptual framework on which 

models are built.

Monitoring Instrumentation and Methodologies

Possibly most critically needed at present are 

investigations that more precisely constrain shallow conduit 

geometry and mixed-volatile degassing phenomena, because 

these parameters strongly influence the final ascent of 

magma and eruptive style. Research on Kīlauea since 2000 

shows much promise on these fronts. In particular, smallaperture seismic and infrasound networks comprising 

instruments with a wide dynamic range are providing 

high-resolution mapping of the location and geometry of 

Kīlauea’s shallow plumbing and degassing systems. Their 

value in conduit imaging is exemplified in recent studies 

integrating seismic and acoustic long-period, very-longperiod, and tremor signals associated with the vigorous 

degassing that accompanied and followed the opening 

of the Halema‘uma‘u vent in 2008 (Chouet and others, 

2010; Dawson and others, 2010; Fee and others, 2010; 

Matoza and others, 2010; Chouet and Dawson, 2011; 

Patrick and others, 2011a). Seismic and acoustic signals, 

in combination with lidar (light detection and ranging) 

and FLIR (forward looking infrared radar) imaging, reveal 

conduit discontinuities at depths of ~1 km (a constriction) 

and ~ 200 m (flare opening to a cavity), which strongly 

influence eruptive patterns. The amplitude, period, and 

duration of signals are helping to determine flow patterns 

in the conduit and the distribution of gas. Particularly 

promising are recent efforts to correlate seismic and 

infrasound records with continuous infrared-spectroscopic 

monitoring of gas emissions (see Sutton and Elias, this 

volume, chap. 7). Since the start of renewed eruptive 

activity at Halema‘uma‘u in 2008, gas compositions 

obtained by OP-FTIR show variations in CO2/S ratio 

indicative of fluctuating depths of gas accumulation and 

varied mode of release.

Field Geology

Sophisticated and integrated monitoring networks are 

essential, but the role of traditional geologic investigations 

cannot be overlooked if predictive models are to have true 

utility. A prime example is afforded by recent studies of 

highly explosive eruptions through analysis of Kīlauea’s 

tephra deposits. Although early workers, most notably 

Perret (1913d), Powers (1916), Stone (1926), Wentworth 

(1938), and Powers (1948), recognized Kīlauea’s explosive 

past, modern follow-up studies have been limited to a few 

seminal investigations, including a summary of earlier 

observations by Decker and Christiansen (1984) and studies 

on tephra stratigraphy (McPhie and others, 1990; Dzurisin 

and others, 1995) and physical volcanology (Swanson and 

Christiansen, 1973; Dvorak, 1992; Mastin, 1997; Mastin 

and others, 2004). 

In contrast to purely magmatic eruptions, research on 

phreatomagmatic (and phreatic) eruptions in Hawai‘i is 

still in a discovery phase. Most or all of the most powerful 

explosive eruptions in Kilauea’s past 2,500 years involved 

external water, including the eruption in 1790 that killed 

many people near the present location of HVO. Today’s 
344 Characteristics of Hawaiian Volcanoes

heightened interest comes from the realization that 

recurrence of such explosive activity could have severe 

consequences for the growing numbers of island residents 

and visitors—and that events in the past 2,000 years have 

sent ash into the jetstream, well within the flight altitudes of 

commercial aircraft (Swanson and others, 2011).

The hazards are all the more alarming in the context 

of recent findings showing that Kīlauea’s explosive events 

have occurred more frequently and over longer periods 

of time than previously thought (Fiske and others, 2009; 

Swanson and others, 2012a, b). For example, new data from 

extensive 14C dating and deposit analysis reveal that the 

widely known Keanakāko‘i Tephra, once believed to have 

been erupted in 1790, actually records a series of violent 

events spanning the period from ca. 1500 to the early 

1800s (Swanson and others, 2012a, b). Moreover, these 

data suggest that Kīlauea’s explosive activity is clustered 

over time. From 500 to 200 B.C.E., effusive eruptions 

dominated. This relatively benign era was followed by 

~1,200 years of mostly explosive activity, a subsequent 

500-year-long period of effusive activity, and then the 

~300-year-long period of explosive eruptions responsible 

for the Keanakāko‘i Tephra, ending in the early 1800s. 

Effusive eruptions have dominated since, except for the 

small phreatic explosions of 1924. This notable cyclicity 

of different eruption styles is currently under intense study 

(Swanson and others, 2014).

Acknowledgments

We are grateful for many opportunities to share 

ideas and fieldwork with our colleagues at the Hawaiian 

Volcano Observatory, and are honored by their invitation 

to contribute to this collection of centennial papers. 

Many thanks to Tom Sisson and Matt Patrick for their 

constructive reviews of our chapter, and to Mike Poland, 

Jim Kauahikaua, Jane Takahashi, Manny Nathenson, 

George Havach, and Peter Stauffer for their help in editing. 

Our chapter also reaped the benefit of expert illustrations 

by Mae Marcaida and the gracious help of Bob Tilling, Tim 

Orr, Ben Gaddis, and Jane Takahashi in culling through 

HVO’s photo archives. Layout and illustration refinements 

by Jeanne S. DiLeo. 

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A lava fountain emanates from part of the Kamoamoa fissure eruption, Kīlauea’s East Rift Zone, March 5–9, 

2011. USGS photograph by M.P. Poland, March 7, 2011.
Tinted photograph of lava fountaining during an eruption of Mauna Loa in 1919. Photograph by I. Morihiro, October 15, 1919, courtesy of 

Roger and Barbara Meyers. 
156° 155°

20°

19°

1950

1926

1859 1935 1984

1942

1800–01

ca. 1800

1960

1955

1955 1983–present

1969–74

1974

1823

1919

1950

1926

1919–1920

0

0

20 40 KILOMETERS

10 20 MILES

Hualālai

Mauna Loa

EXPLANATION

Kīlauea

Figure 1. Shaded-relief map of Hawai‘i showing locations of all historical 

lava flows, color-coded according to source volcano. Flows mentioned 

specifically in text are labeled by date and shown in darker colors.

Abstract

The Hawaiian Volcano Observatory (HVO) was established 

as a natural laboratory to study volcanic processes. Since the 

most frequent form of volcanic activity in Hawai‘i is effusive, a 

major contribution of the past century of research at HVO has 

been to describe and quantify lava flow emplacement processes. 

Lava flow research has taken many forms; first and foremost it 

has been a collection of basic observational data on active lava 

flows from both Mauna Loa and Kīlauea volcanoes that have 

occurred over the past 100 years. Both the types and quantities 

of observational data have changed with changing technology; 

thus, another important contribution of HVO to lava flow studies 

has been the application of new observational techniques. Also 

important has been a long-term effort to measure the physical 

properties (temperature, viscosity, crystallinity, and so on) of 

flowing lava. Field measurements of these properties have both 

motivated laboratory experiments and presaged the results of 

those experiments, particularly with respect to understanding the 

rheology of complex fluids. Finally, studies of the dynamics of lava 

flow emplacement have combined detailed field measurements 

with theoretical models to build a framework for the interpretation 

of lava flows in numerous other terrestrial, submarine, and 

planetary environments. Here, we attempt to review all these 

aspects of lava flow studies and place them into a coherent 

framework that we hope will motivate future research.

Introduction

Overview

Rivers of lava are an iconic image of Hawaiian volcanism. 

With the frequent eruptions of Mauna Loa throughout the 19th and 

early 20th centuries, and the persistent activity of Kīlauea from 

1969 to 1974 and since 1983 (fig. 1), Hawai‘i has served as one 

of the most important natural laboratories in the world for studies 

of lava flows. Furthermore, the higher flow rates associated with 

most eruptions of Mauna Loa, and the longer duration of eruptions 

from Kīlauea, have provided important insight into a wide range 

of flow emplacement processes. In this chapter, we review observations of lava flow activity over the past century, outline many 

of the techniques developed in Hawai‘i to map and analyze lava 

flows, and discuss the physical conditions of flow emplacement. 

Finally, we examine some of the contributions of Hawaiian lava 

flow studies to the investigation of planetary volcanism.

As demonstrated by the global perspective of The Volcano 

Letter (compiled in Fiske and others, 1987; herein abbreviated 

as “VL” followed by the appropriate issue number) and other 

Hawaiian publications, research in Hawai‘i has not gone 

on in a vacuum; instead, it has clearly been influenced by 

1

University of Bristol, U.K. 

2

University of Oregon.

3

U.S. Geological Survey.

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 9

A Century of Studying Effusive Eruptions in Hawai‘i

By Katharine V. Cashman1,2 and Margaret T. Mangan3
358 Characteristics of Hawaiian Volcanoes

Mounts Vesuvius and Etna in Italy, but also that the frequency 

and accessibility of Hawaiian lava flows “make them a 

peculiarly instructive field for the student of volcanic science, 

as well as an attractive one for the lover of the marvellous” 

(p. v). Dana proceeded to review key observations from 19th

century Hawaiian studies, including (1) the eruption of (dense) 

olivine-phyric flows from high-elevation Mauna Loa vents, (2) 

the simultaneous activity of Mauna Loa and Kīlauea, (3) the 

“mobility” of liquid basalt, (4) the recognition of pit craters 

as characteristic of basaltic volcanoes, and (5) the activity of 

lava lakes. He also listed areas that required additional study, 

including the dynamics of lava streams (channels) and the 

formation of lava tubes, as well as the relation of Kīlauea 

to Mauna Loa and the driving force for eruptive activity. 

These broad topics identified by Dana have, indeed, framed 

volcanologic studies in Hawai‘i over the past 100 years.

In addition, 19th century volcanologists provided a 

foundation for more focused topics of research that continue 

to garner attention. One is the origin of the two primary 

Hawaiian lava flow types, pāhoehoe and ‘a‘ā. These words 

were adopted from the Hawaiian language and introduced 

into the volcanological lexicon to describe lava with smooth 

or broken (clinkery) surfaces, respectively (fig. 2). In his 

geologic summary of Hawai‘i, Dana (1849) noted that 

‘a‘ā and pāhoehoe formed during “different phases in the 

volcanic action of one and the same period” (p. 162), with 

the only difference being “a variation in the rapidity of 

A

D

B

C

Figure 2. Photos of characteristic Hawaiian lava flow morphologies. A, Ropy, degassed pāhoehoe flow surface. B, Active ‘a‘ā flow surface. C, 

Spongy pāhoehoe flow interior (individual vesicles are 2–5 mm across). D, Stacked, thin ‘a‘ā channel overflows, showing characteristic dense 

cores (10–20 cm thick) and clinkery flow boundaries. All photographs by K.V. Cashman except B, which is by S.A. Soule.

academic researchers throughout the United States, as well 

as by scientists studying volcanoes around the world. In fact, 

only the frequent lava flow activity at Mount Etna, Italy, is 

comparable to that in Hawai‘i in terms of its accessibility and 

impact on the study of effusive volcanism. Where appropriate, 

we note these global links, although here we focus primarily on 

the contributions to lava flow studies from research by, and in 

conjunction with, the Hawaiian Volcano Observatory (HVO). 

The View from the 19th Century

The first written accounts of Hawaiian lava flows were by 

Ellis (1825), who traveled throughout Hawai‘i as a missionary 

in 1823. Another missionary—Titus Coan—provided detailed 

observations of most of the eruptions of Mauna Loa and 

Kīlauea between 1835 and 1882; many of these accounts were 

published in the American Journal of Science (see Wright and 

Takahashi, 1989, 1998; Barnard, 1990). In the 19th century, 

scientists, such as J.D. Dana, W.D. Alexander, C.E. Dutton, 

W.T. Brigham, and C.H. Hitchcock, assembled their own 

observations and those of local observers (including Coan), so 

that, by the end of that century, the morphology of Hawaiian 

lava flows, the formation of lava channels and tubes, and 

the importance of cooling on flow evolution were all well 

described, albeit qualitatively. By the end of the century, a 

summary by Dana (1890) argued convincingly, not only for 

Hawai‘i as a natural laboratory equivalent to that provided by 
A Century of Studying Effusive Eruptions in Hawai‘i 359

motion, or a renewal of movement from a cessation” (p. 163). 

Close observations of paired ‘a‘ā and pāhoehoe lavas from 

the 1859 Mauna Loa eruption allowed the surveyor W.D. 

Alexander to extend this observation by inferring that the 

“mode of cooling” controlled the morphology, with thin, 

commonly near-vent pāhoehoe flows being in a “state of 

complete fusion,” while less fluid ‘a‘ā formed “grains like 

sugar” (Alexander, 1859). This metaphor has persisted 

not only through the HVO literature (for example, Jaggar, 

1947; Macdonald, 1953), but also inspired the use of sugar 

crystallization (in fudge) as an analog for educational 

purposes (Rust and others, 2008). Alexander (1886) also 

recognized that the slope (flow rate) over which lava flowed 

was important in ‘a‘ā formation; thus, by the end of the 

19th century, two primary controls on surface texture—

crystallinity and flow rate—had already been identified.

Another long-lived topic of research begun in the 19th 

century was the quest to measure the physical properties 

of flowing lava (particularly temperature and viscosity). 

The spectacle of rapidly moving lava rivers of “irresistible 

impetuosity” captivated the earliest Western visitors to Hawai‘i 

(Ellis, 1825). Descriptions of lava as being “at a white heat 

and apparently as liquid as water” (Haskell, 1859) and of 

the “superior mobility” of Hawaiian lava, the relation of 

this mobility to lava temperature, and the role of mobility 

(viscosity) in processes of both eruption and flow emplacement 

(Dana, 1890) all illustrate the importance that early 

investigators placed on the relation between basic physical 

properties and the dynamics of lava flow emplacement. 

A Summary of Effusive Activity from 

1912 to 2012

The science of effusive eruptions has advanced in 

conjunction with both changes in the eruptive behavior of 

Kīlauea and Mauna Loa volcanoes and the advent of new 

technologies. The first 5 decades of activity during the period 

1912–2012 are recorded primarily in HVO documents, such as 

the VL and its early serial publications (compiled by Fiske and 

others, 1987; Bevens and others, 1988). More recent eruptions 

are thoroughly documented in HVO reports, U.S. Geological 

Survey Professional Papers, journal articles, and the HVO Web 

site (http://hvo.wr.usgs.gov/).

The First Five Decades

When HVO was founded in 1912, Mauna Loa was 

erupting at intervals of 3.5 years at the summit and 6 years 

on its flanks (VL 440), and was considered the “grand theater 

for lava flows” (Hitchcock, 1911). This trend continued 

with important eruptions along Mauna Loa’s Southwest Rift 

Zone in 1919 and 1926 and Northeast Rift Zone in 1935 and 

1942 (fig. 1). These eruptions were used by HVO scientists 

to (1) document the initiation and shutdown of activity at 

different vents and (2) monitor the rates and mechanisms of 

lava flow advance through steep and forested reaches, such 

as those of Hawai‘i’s southwest coast. Employment of U.S. 

military airplanes allowed scientists to obtain both real-time 

observations and aerial photographs of eruptive activity. 

Aerial capabilities prompted the first modern attempt at lava 

flow diversion in 1935, when bombs were dropped to disrupt 

robust lava tubes that were funneling lava flows toward Hilo 

(VL 431–432, 442, 445, 465, 506). Aerial observations in 

1942 also permitted estimates of both early rates of flow 

advance (>6 mi in 5 hours [0.5 m/s]) and final flow volumes 

(2×109 ft3 [5.7×107 m3]; VL 476). 

A summit eruption of Mauna Loa in 1949 preceded a 

massive eruption in June 1950 along the Southwest Rift Zone 

(VL 508, 509; Finch and Macdonald, 1953). Aerial and groundbased observations of the 1950 lava flow showed the source to 

be an en echelon fissure that stretched for nearly 20 km along 

Mauna Loa’s Southwest Rift Zone. The flow fed a wide area of 

anastamosing lava streams that plowed down the volcano’s forested southwest flank and poured into the ocean (fig. 3). Early 

flows covered the 24 km from the vent to the coastal highway 

in less than 3 hours (average flow advance rate, >2.2 m/s). 

HVO scientists were able to provide good estimates of both the 

area (35.6 mi2 [91 km2]) and the volume (514×106 yd3 on land, 

another 100×106 yd3 in the ocean [470×106 m3 total]) of the lava 

and to demonstrate that from one-half to two-thirds of the total 

volume was erupted during the first 36 hours of activity. These 

measurements yielded impressive initial (rather than average) 

effusion rates of ~4,700 yd3/s (~3,600 m3/s).

During the same period, activity at Kīlauea was restricted 

primarily to the caldera and Halema‘uma‘u Crater, which 

contained an active lava lake until 1924 (for example, Bevens 

and others, 1988). Effusive eruptions within the caldera were 

common but generally small in volume. One exception was an 

Figure 3. Photograph showing the Ka‘apuna flow, 1950 Mauna Loa 

eruption, entering the sea. Note modest amount of steam generated 

at ocean entry, as well as absence of explosions caused by lava-water 

interactions. Photograph by Transpacific Airways.
360 Characteristics of Hawaiian Volcanoes

eruption on the Southwest Rift Zone in 1919–20 (fig.1), which 

produced paired ‘a‘ā and pāhoehoe flows and the Maunaiki 

shield (Rowland and Munro, 1993). This eruption is important 

in that it indicated a complex storage system beneath 

Halema‘uma‘u with an intricate connection to Maunaiki that 

evolved over time (see Bevens and others, 1988). Activity 

at Halema‘uma‘u ended with a phreatic eruption in 1924; 

subsequently, Kīlauea Volcano outside the summit was quiet 

until 1955, when a major flank eruption occurred on the East 

Rift Zone (Macdonald and Eaton, 1964; Moore, 1992). 

The location of the 1955 eruption in the eastern Puna 

District (fig. 1), close to a seismic station installed in 1954, 

triggered a rapid response by HVO scientists. The proximity 

to populated areas helped HVO staff to locate the first flows, 

allowing them to be on site from the start of the eruption. 

From a perspective of geologic hazards, the 1955 eruption 

was the first to require extensive evacuations; as a result, it 

prompted the first modern attempts to construct barriers to 

deflect lava flows from critical areas (Macdonald, 1958). From 

a scientific perspective, this eruption provided scientists with 

their first opportunity to observe and photograph, at close 

range, the formation of a volcanic vent system—from the first 

opening of a fissure in the ground, through the appearance 

of lava, to the formation of cones and flows and, finally, to 

the cessation of activity. Other opportunities provided by this 

eruption included observations of pit crater formation and of 

active flow fronts at close range, which allowed study of the 

mechanics of flow movement and temperature measurements 

of both flow fronts and lava fountains (VL 529–530). Finally, 

this eruption provided an opportunity for direct comparisons 

between instrumental and field-based observations. Specifically, HVO scientists were able to correlate the amplitude of 

harmonic tremor with lava extrusion rate and record the progress of an eruption by monitoring changes in summit tilt. 

The 1955 eruption was followed by a large summit eruption in Kilauea Iki pit crater in 1959 (see Helz and others, this 

volume, chap. 6; Mangan and others, this volume, chap. 8) 

and, in 1960, by an eruption at Kapoho just downrift from the 

1955 eruption site (fig. 1). Again, mitigation was attempted, 

with several barriers constructed in an effort to save homes 

and places of historical interest (Macdonald, 1962). Lava ultimately covered an area of 10 km2

 (including all the barriers) 

with an estimated volume of 0.122 km3 of new material.

1969–2012

Although numerous Hawaiian eruptions occurred during the period 1969–2012, those of Mauna Loa in 1984 and 

Kīlauea in 1969–74 (Mauna Ulu) and from 1983 to the present 

(Pu‘u ‘Ō‘ō) are particularly well documented and have 

cemented HVO’s reputation as a laboratory for studying 

basaltic volcanoes. These eruptions have provided new perspectives on old questions of lava flow emplacement, including measurements of the thermal efficiency of lava channels 

and tubes, the rheologic changes that accompany cooling and 

crystallization, the mechanisms of flow advance on both steep 

and shallow slopes, and the development of characteristic flow 

morphologies. Observers of these eruptions have also benefited from increasing access to eruption sites by helicopter, 

from the digital revolution (with its accompanying transformations in the acquisition, storage, and global transfer of a vast 

array of data), and from the application of numerous remotesensing techniques. 

Like many previous Mauna Loa eruptions, the 1984 

eruption started at the summit caldera (Moku‘āweoweo) on 

March 25. Within hours, eruptive activity migrated down 

the Northeast Rift Zone to establish a stable vent at 2,850-m 

elevation that fed the next 3 weeks of eruptive activity. The 

eruption, which produced 0.22 km3 of lava and covered an 

area of 48 km2 (fig. 4), illustrates characteristics that are typical of many “open channel” lava flows in Hawai‘i. As eruptive 

vents migrated from east to west over time, they directed lava 

into different drainage basins, so that the flow direction shifted 

from east to northeast, toward Hilo. Each of the primary flows 

had a complex form, with numerous bifurcations (and some 

confluences), commonly around topographic barriers. Thus, 

Hawaiian lava flows are generally distributary, such that the 

total lava volume is divided between increasing numbers of 

flow lobes with distance from the vent. 

From a geologic-hazards perspective, the advance of 

lava flows toward both the Kūlani prison and the city of Hilo 

caused some concern, although cooling ultimately arrested 

flow advance (Lockwood and others, 1987). From a scientific 

perspective, repeat observations at several places along the 

main lava channel provided unprecedented data on the flow of 

lava within the channel, as well as on the mechanisms of flow 

advance (Lipman and Banks, 1987); this unique dataset stimulated analysis of transport conditions through lava channels 

(Crisp and Baloga, 1994; Crisp and others, 1994) and continues to serve as a benchmark calibration for interpretations of 

older flows (Riker and others, 2009) and construction of flow 

models (for example, Harris and Rowland, 2001).

Two protracted eruptions on Kīlauea’s East Rift Zone 

have provided similarly valuable datasets on the formation 

of compound pāhoehoe flow fields. The Mauna Ulu eruption 

(1969–74) produced 0.34 km3 of lava that covered an area of 

46 km2, with the last 3 years of eruptive activity focused on 

the Mauna Ulu shield. Detailed observations of this eruption provided important new insight into shield formation 

and the characteristics of pāhoehoe lava (Swanson 1973), 

the behavior of shallow submarine lava flows (Moore and 

others, 1973), the pāhoehoe-to-a‘ā transition (Peterson and 

Tilling, 1980), the formation of lava tubes (Peterson and 

others, 1994), and the dynamics of lava lakes (Tilling, 1987). 

The importance of this eruption has been somewhat eclipsed, 

however, by the unusually long Pu‘u ‘Ō‘ō-Kupaianaha eruption farther downrift. This eruption, which began in early 

1983 and is still ongoing as of September 2014, has provided 

a unique opportunity not only to study complex ‘a‘ā and 

pāhoehoe flow fields, but also to connect effusive activity 

to the petrology, geophysics, and geochemistry of Kīlauea’s 
A Century of Studying Effusive Eruptions in Hawai‘i 361

magmatic system (for example, Heliker and others, 2003; 

Poland and others, this volume, chap. 5). Below, we summarize observations from these eruptions on the basic characteristics of Hawaiian lava flows before addressing advances in 

more quantitative aspects of lava flow emplacement.

Basic Lava Flow Characteristics

Hawaiian lava flows are commonly classified as either 

‘a‘ā or pāhoehoe on the basis of their surface morphology 

(fig. 2), although numerous transitional forms also exist. 

Flow type is not a function of composition but is controlled, 

instead, by processes related to the eruption and transport 

of lava (Macdonald, 1953). Here, we review observations 

on both flow morphology and flow geometry because they 

provide the framework for modern studies of lava properties, 

flow emplacement conditions, and hazard assessments.

Lava Morphology 

Field geologists working in Hawai‘i have long been 

fascinated by the morphology of lava flow surfaces. Pāhoehoe, 

with its endless variety of surface forms, has spawned a 

proliferation of colorful descriptors, such as sharkskin, 

toothpaste, rubbly, slabby, festooned, dendritic, shelly, blue 

glassy, pillow-like, and even entrail-like (for example, Stearns 

and Macdonald, 1946; Wentworth and Macdonald, 1953; 

Swanson, 1973; Rowland and Walker, 1987; Hon and others, 

1994; Self and others, 1998; Kauahikaua and others, 2003). 

2,180,000

2,170,000

2,160,000

260,000

Easting, in meters

Northing, in meters

3000

2000

1000

250,000

300 m3/s

100 m3/s

30 m3/s

Cumulative volume, in millions of cubic meters

March April

225

200

175

150

125

100

75

50

25

25 30 1 5 10

0

0

4 8 KILOMETERS

2 4 MILES

A B

Figure 4. Lava flow from the 1984 Mauna Loa eruption (see fig. 1 for flow location). A, Sketch map showing extensive branching that accompanied 

flow emplacement. Heavy black lines denote “lava sheds,” or topographically defined drainage basins. Coordinates in UTM Zone 6, Old Hawaiian datum 

(modified from Kauahikaua and others, 1995). B, Cumulative erupted volume of 1984 flow versus time. On the basis of this plot, eruption can be divided 

into three phases: rapid effusion (300 m3/s) for the first 5 days of activity, a protracted period of moderate effusion rates (100 m3/s), and, finally, a waning 

stage (30 m3/s; redrafted from Lipman and Banks, 1987).

‘A‘ā flows have not inspired a similar proliferation of names 

but also vary according to the size and shape of the surface 

clinkers (for example, Jones, 1943; Soule and others, 2004). 

Each morphologic variation reflects a set of intrinsic and 

extrinsic conditions that includes the specifics of the lava 

properties (particularly rheology) and external factors that 

control emplacement (such as volumetric flow rate, underlying 

slope, and topographic confinement, all of which contribute 

to the deformation conditions of the flowing lava). For this 

reason, investigation of the apparently simple question of crust 

morphology has led to a much deeper understanding of the 

thermal, rheologic, and dynamic evolution of lava flows. 

Jaggar (VL 281) summarized the state of knowledge 

of ‘a‘ā and pāhoehoe derived from HVO studies covering the first 3 decades of the 20th century. By that time, he 

could state that “there is no essential difference chemically 

between aa and pahoehoe,” that “fountaining pahoehoe at the 

source of a flow may turn into aa clinkers within a half mile 

of the vent, and remain aa for the rest of its course down the 

mountain into the sea,” and that “When an observer stands 

on the bank of a golden, liquid torrent of lava flowing so 

rapidly as to make no crusts or skins, he can not tell from the 

appearance of the liquid whether it will solidify as pahoehoe 

or aa.” He confirmed Alexander’s (1859) hypothesis that ‘a‘ā

is more crystalline than pāhoehoe and was able to extend this 

analysis to include the effects of stirring, with reference to 

the experimental data of Emerson (1926). Jaggar also recognized that the characteristics of pāhoehoe surface folds are 

determined by the “thickness of the flexible crust” at the time 

of deformation and that the vesicular crust of pāhoehoe “is 

an excellent heat insulator.” This work was later extended by 
362 Characteristics of Hawaiian Volcanoes

application of folding analysis to the ductile layer (Fink and 

Fletcher, 1978), which can be used to determine the cooling and emplacement history of individual lava flows (for 

example, Gregg and others, 1998).

Jaggar’s (VL 281) summary of the characteristics of 

‘a‘ā and pāhoehoe has stood the test of time. Subsequent 

experiments have reproduced Emerson’s (1926) results, using 

different basaltic compositions and experimental conditions 

(for example, Kouchi and others, 1986; Sato, 1995); together, 

they show that shear (dynamic) crystallization is critical 

for ‘a‘ā formation (see Rust and others, 2008). Macdonald 

(1953) assembled both a comprehensive description of the 

physical attributes of ‘a‘ā and pāhoehoe flow morphologies 

and an extensive list of the conditions under which pāhoehoe 

lava could transform to ‘a‘ā. He noted that flows change 

downslope from pāhoehoe to ‘a‘ā but not the reverse (see 

Jurado-Chichay and Rowland, 1995, and Hon and others, 

2003, for a more nuanced discussion of this point), that 

the distance lava travels from the vent before changing 

from pāhoehoe to ‘a‘ā varies inversely with eruptive vigor 

(volumetric flow rate from the vent), and that pāhoehoe is 

hotter and more gas rich, and contains more quenched glass, 

than ‘a‘ā. 

Peterson and Tilling (1980) formalized these 

observations by defining the pāhoehoe-to-‘a‘ā transition 

as a threshold in the relation between shear-strain rate and 

apparent viscosity (fig. 5). Two views of this threshold 

exist. The perspective from observations of crystalline lava 

flows at Mount Etna is that it represents a failure envelope 

for flow crusts under conditions of continuous deformation 

(Kilburn, 1990, 1993). In Hawai‘i, however, where lava 

emerges from the vent at near-liquidus temperatures, the 

importance of achieving a critical crystallinity is more 

apparent (for example, Rowland and Walker, 1990; Crisp 

and others, 1994; Cashman and others, 1999). Coupling of 

field and laboratory measurements with the results of analog 

experiments shows that the transition from pāhoehoe to ‘a‘ā 

is determined primarily by a threshold value of apparent 

viscosity except at very low strain rates, where pāhoehoe 

morphologies can be maintained to higher viscosities (fig. 5). 

This assessment of the pāhoehoe-to-‘a‘ā transition presages 

recent rheologic studies that show the dependence of critical 

rheologic transitions on particle shape, volume fraction, size 

distribution, and strain rate (for example, Costa and others, 

2009; Castruccio and others, 2010; Mueller and others, 2010, 

2011; see subsection below entitled “Rheology”).

Geometry of Flows and Flow Paths—An 

Observational History

The distribution of flow surface morphologies varies in 

both space and time, is directly linked to changing conditions 

of flow emplacement, and is determined by the type and 

geometry of lava transport systems. Flow surface mapping 

techniques have evolved in conjunction with changes in 

available technology, including first aerial, and then satellite, 

observational platforms that allow flow fields to be viewed, and 

analyzed, in their entirety.

The establishment of aerial monitoring in the 1930s 

permitted detailed observations of flow geometries and 

active flow surfaces. As a result, accounts of the 1935 

eruption of Mauna Loa contain the first detailed descriptions 

of the intricate geometry of active lava channels and spatial 

changes in surface morphology from proximal braided 

pāhoehoe streams (“braided torrents of glowing liquid [that] 

were from 30 to 50 feet [10–15 m] wide, near their sources”) 

to channelized ‘a‘ā within about 1 mi (1.6 km) of the 

vents (VL 439). Observers also noted that the lava streams 

“became narrower farther down the mountain,” where “the 

rapidity of their forward motion became less,” and that the 

flow surface showed a temporal progression as early ‘a‘ā was 

covered by later pāhoehoe (VL 429). 

The 1950 eruption of Mauna Loa afforded new 

opportunities to observe large channelized flows, although 

the steep forested flanks of Mauna Loa’s Southwest Rift 

Zone limited most of these detailed observations to areas 

between the highway and the coast (fig. 3). Documentation 

of the 1950 lava flows included measurements of maximum 

channel flow rates over lava cascades (35 mi/h [>15 m/s]), 

standing waves 12 ft (3.6 m) high below the cascades, 

and surges in flow advance, with flow rates (7–8 mi/h 

[3–3.5m/s]) that exceeded those of normal channel flow 

(4–5mi/h [~2 m/s]; Finch and Macdonald, 1953). Scientists 

noted the abundance of blocks transported through the 

channel, as well as their tendency to obstruct channels 

and create overflows. They also made numerous optical 

3

2

1

0

Transition threshold zone

2 3 4

Shear rate, in inverse seconds

Log of apparent viscosity, in pascal-seconds

Pāhoehoe

‘A‘ā

Figure 5. Plot of shear rate versus apparent viscosity fields for flow 

surface morphologies (pāhoehoe and ‘a‘ā), calibrated from a combination 

of field observations (Cashman and others, 1999; Hon and others, 

2003; Soule and others, 2004; Riker and others, 2009) and laboratory 

experiments (Soule and Cashman, 2005), representing best approximation, 

to date, of pāhoehoe-to-‘a‘ā transition in Hawaiian lava flows. 
A Century of Studying Effusive Eruptions in Hawai‘i 363

pyrometry measurements of flow temperature (see subsection 

below entitled “Temperature”) and reported widespread 

methane generation from burning vegetation. 

During the 1970s and 1980s, increased staffing at HVO 

and improved access to active flow fields (by trails, roads, and 

helicopters) brought new detail to flow field maps, including 

the time evolution of flow emplacement (for example, Lipman 

and Banks, 1987; Tilling and others, 1987; Wolfe and others, 

1988; Heliker and others, 2001). These maps, which are now 

posted routinely on HVO’s Web site, allow analysis of the 

relations between effusion rates, rates of lava-flow advance, and 

topographic confinement provided by earlier emplaced flows 

(for example, Kauahikaua and others, 2003). At the same time, 

helicopter support facilitated repeat observations in places along 

lava channels (Lipman and Banks, 1987) and lava tubes (for 

example, Mangan and others, 1995a; Kauahikaua and others, 

1996; Heliker and others, 1998) that can be used to constrain 

the conditions of lava transport (see section below entitled 

“Dynamics of Lava Flow Emplacement”). 

The prolonged Pu‘u ‘Ō‘ō-Kupaianaha eruption has 

produced a range of new flow mapping techniques. During 

the first decade of activity, maps were constructed directly 

on aerial photographs, using helicopter surveillance and 

ground-based observations, coupled with postemplacement 

aerial photographs (for example, Wolfe and others, 1988; 

Mattox and others, 1993). The advent of hand-held Global 

Positioning System (GPS) units drastically changed mapping 

techniques by providing both accurate locations of flow 

features and digital data appropriate for use in geographicinformation-system (GIS) mapping utilities. Conversion to 

digital mapping has improved the accuracy and efficiency of 

mapping efforts, especially since GPS maps of flow outlines 

can now be generated by helicopter surveys. 

The 1990s also saw the application of remote-sensing 

techniques to near-real-time mapping of lava flows. Frequent 

(every 15 minutes) low-resolution (4 km/pixel) views of the 

lava flow field generated by the Geosynchronous Orbiting 

Environmental Satellite (GOES) allow tracking of hot spots 

related to surface activity (Harris and others, 2001). Advanced 

Very High Resolution Radiometer (AVHRR; 1 km), Landsat 

Thematic Mapper (TM; 30–120 m) and Moderate Resolution 

Imaging Spectroradiometer (MODIS) thermal images can 

be calibrated to yield estimates of the time evolution of lava 

effusion rates, which provide important input to predictive 

models (Flynn and others, 1994; Harris and others, 1998; 

Wright and others, 2001, 2002). Recent advances in flow 

mapping in both Hawai‘i and Italy include the use of 

airborne light detection and ranging (lidar) and satellitebased synthetic-aperture radar (SAR) for the generation of 

digital elevation models (DEMs; Rowland and others, 1999; 

Mazzarini and others, 2005), relative flow age determination 

(Mazzarini and others, 2007), and flow mapping (Zebker and 

others, 1996; Favalli and others, 2010; Dietterich and others, 

2012; Cashman and others, 2013). Importantly, these data also 

provide new insights into flow field evolution by supplying 

detailed views of flow field construction. 

Physical Properties of Flowing Lava

At the same time that HVO scientists were observing and 

mapping lava flows, they were also attempting to measure the 

physical properties of flowing lava and to link these properties 

to flow emplacement conditions. These properties include 

not only lava temperature, but also changes in bubble and 

crystal content and their effect on lava rheology. Here we 

show how field-based observations and measurements of the 

physical properties of Hawaiian lavas have provided important 

data on the structure and rheology of silicate melts; spurred 

laboratory and theoretical research on the relations between 

the temperature, rheological and material properties of mafic 

magma; and provided key information on the rates and types 

of phase changes during lava transport. These data are critical 

to understanding the dynamics of lava flow emplacement.

Temperature

Early observers used the color of lava (for example, 

“white hot” versus “cherry red”) to determine the relative 

temperature of different parts of individual lava flows. Use 

of a color scale to measure temperature was not unique to 

Hawai‘i; for example, Perret (VL 202) used a color-based 

scale to estimate a temperature of 1,200 °C for Etna lava from 

an eruption in 1908. E.S. Shepherd and F.A. Perret made the 

first direct temperature measurements of Hawaiian lava, using 

a platinum-rhodium thermocouple (Shepherd, 1912) to obtain 

a temperature of 1,000 °C for lava in the Halema‘uma‘u lava 

lake. Jaggar (1917, 1921) experimented with the use of Seger 

cones (used in firing pottery) to measure temperature-depth 

profiles within Halema‘uma‘u. The Seger cones, however, 

produced sufficiently confusing results that most workers 

relied on temperature measurements by optical pyrometer, 

which have yielded temperatures of 1,075–1,130 °C for 

Halema‘uma‘u fountains and flank vents, 1,120–1,190 °C for 

the unusually hot fountains accompanying the 1959 eruption of Kīlauea Iki, and 900–1,030 °C for channelized lava. 

Problems with these readings lie primarily in the difficulty in 

obtaining an unobstructed view of the fountain/flow interior 

and of knowing the appropriate correction for emissivity; as a 

result, it has commonly been assumed that optical pyrometry 

readings are 20–30 °C too low (Macdonald, 1963). In fact, in 

situ temperature (thermocouple) measurements at source vents 

of 1,140–1,147 °C (Mauna Loa 1984) and 1,110–1,150 °C 

(Pu‘u ‘Ō‘ō) are slightly higher than most optical pyrometry 

measurements, confirming Macdonald’s (1963) suspicions. 

When conditions are optimal, however, temperature measurements made by thermocouple and two-color infrared pyrometer can agree to within 5 °C (Lipman and Banks, 1987); used 

together, they allow documentation of the thermal history of 

lava fountains and flows in both space and time. 

New experimental and analytical capabilities in the 

1970s fueled a boom in the design and calibration of geothermometers that form the basis of modern petrologic 
364 Characteristics of Hawaiian Volcanoes

investigations of magma storage conditions (for example, 

Blundy and Cashman, 2008; Putirka, 2008). Most useful for 

lava flow studies were glass geothermometers based on an 

observed linear relation between the temperature and MgO 

content of Kīlauea and Mauna Loa melts (Helz and Thornber, 

1987; Montierth and others, 1995). These glass geothermometers have been used to examine spatiotemporal patterns in 

the temperatures of active and solidified flows (Cashman and 

others, 1994, 1999; Mangan and others, 1995b; Clague and 

others, 1999; Soule and others, 2004; Riker and others, 2009) 

and to constrain the thermal efficiency of lava tubes and 

channels (Helz and others, 1995, 2003, and this volume, chap 

6; see subsection below entitled “Lava Tubes”). 

Short-term spatial and temporal variations in the temperature of lava flow surfaces are best captured by ground-based 

thermal imaging systems, such as forward-looking infrared 

(FLIR) (fig. 6). The potential of FLIR data is illustrated by a 

detailed study of small, tube-fed pāhoehoe lobes formed at 

Kīlauea in August 2004 (Ball and others, 2008), where FLIR 

images were used to test models of stationary-flow cooling 

(for example, Keszthelyi, 1995a; Keszthelyi and Denlinger, 

1996; Harris and others, 2005; Ball and Pinkerton, 2006) 

and to examine the effects of flow emplacement dynamics 

on heat loss. FLIR data document rapid initial cooling of 

pāhoehoe flow surfaces by radiative cooling and help explain 

the low apparent temperatures of flowing lava obtained by 

optical pyrometry. The FLIR data can also be used to test 

cooling models (for example, fig. 6B) and to link surface 

temperatures to changes in the material properties of lava. 

For example, Ball and others (2008) document pāhoehoe 

rope formation at Tsurface ~800 °C and development of crust 

strength at Tsurface ~700 °C, as manifested by the onset of flow 

inflation. These threshold temperatures are higher than those 

inferred from glass geospeedometry analysis, which suggest 

that ductile deformation can continue to temperatures as low 

as 627 °C, as long as mechanical perturbations to the crust 

Temperature, in degrees Celsius

Time, in seconds

400

600

800

1,000

0 100 200 300 400

No wind

Wind speed 20 m/s

Measured lobe temperature

Modeled temperature; 

wind speed 15 m/s

A B

occur at sufficiently long time scales (1–10 s; Gottsmann 

and others, 2004). These temperatures are lower than those 

inferred from lava lake drill cores obtained by HVO scientists 

and suggest a transition from ductile to brittle deformation at 

T~800 °C (Wright and Okamura, 1977). 

Rheology

The first estimates of lava viscosity compared the velocity and depth of channelized lava with measurements of water 

flows (Becker, 1897; Palmer, 1927). The results of these early 

studies—that lava was only 10 to 60 times more viscous than 

water (that is, 1×10-2 to 6×10-2 Pa·s)—were orders of magnitude 

too low, because the applied formulation assumed turbulent 

flow, which is appropriate for water but not for lava (Nichols, 

1939; Wentworth and others, 1945). Subsequent application of 

laminar flow models derived apparent viscosity estimates of 

3×103 to 20×103 poise (300–2,000 Pa·s) for Mauna Loa lavas 

and 2×103 to 100×103 poise (200–10,000 Pa·s) for Kīlauea lavas 

(Macdonald, 1963). Observational constraints on viscosity were 

limited, however, by both the accuracy of field measurements 

and assumptions of flow homogeneity. A critical but difficult 

measurement for lava viscosity estimates is that of flow depth, 

which cannot be measured directly (VL 480; Lipman and Banks, 

1987). Also problematic is the complex thermal structure of an 

active lava flow, because assumptions of homogeneity ignore 

the formation of surface crusts. The end result is that field-based 

measurements yield only apparent (integrated) lava viscosities 

that are difficult to correlate with laboratory studies of homogeneous liquids or liquid-particle suspensions. 

Detailed observations of the 1984 Mauna Loa lava flows 

confirmed that the apparent viscosity of vent lavas may be as 

low as 100–200 Pa·s. Apparent viscosity increases exponentially 

along the channel (fig. 7A) because of both internal 

crystallization (fig. 7B) and crust formation (Moore, 1987; Crisp 

and others, 1994). Interestingly, the crystallinity data plotted 

Figure 6. Temperature variation of lava flow surface at Kīlauea. A, Composite of thermal image and color digital photograph. White and yellow areas, 

active pāhoehoe breakouts and red areas, inactive, but still warm, parts of flow surface (USGS photograph by Matt Patrick taken on April 26, 2012). B, 

Modeled and forward-looking-infrared lava flow surface temperatures versus time. Note rapid (radiative) surface cooling over first few minutes after initial 

lobe breakout (from Ball and others, 2008).
A Century of Studying Effusive Eruptions in Hawai‘i 365

in figure 7B show two different slopes (average crystallization 

rates). Earlier (higher) rates of crystallization over the first 

~100 hours of effusive activity were apparently driven by uprift 

degassing (that is, the erupted lava was initially supercooled 

with respect to its temperature at atmospheric pressure because 

of the presence of dissolved volatiles). Degassing-induced 

crystallization continued for the duration of the eruption but at a 

reduced rate as atmospheric equilibrium was approached (Crisp 

and others, 1994).

Experimental investigations of the rheology of 

Hawaiian basalts paralleled field-based measurements. In 

situ measurements by Jaggar monitored the entry rate of lava 

into a metal cylinder (VL 357); in situ experiments were 

later conducted in Kīlauea’s lava lakes, using a rotating shear 

viscometer (Shaw and others, 1968; see Mangan and others, 

this volume, chap. 8). In situ studies were complemented by 

laboratory studies on basaltic magma, although these were 

complicated by the need to employ different measurement 

1

2

3

4

5

6

Log apparent viscosity, in Pascal seconds

0 5 10 15 20

Distance, in kilometers

Calculated viscosity

0

0.16

0.08

0.24

0.32

0.40

Crystal volume fraction

0 100 200 300 400

Time since eruption start, in hours

500

Crystal volume fraction

Log relative viscosity, in Pascal seconds

0 0.2 0.4 0.6 0.8 1.0

0

1

2

3

4

5

6

7

8

9

ɸm = 0.6

ɸm = 0.7

ɸm = 0.8

ɸm varies with crystal size

distributrion, shape, and

orientation (stress)

0.55 0.6

Crystal volume fraction

Log yield strength, in Pascal seconds

0

-1

-2

1

2

3

0.35 0.45 0.4 0.5

Τy(ɸ)=D(ɸ- ɸc)

8

0

0 1

Yield strength

Crystal volume fraction

Solid

f

y fm

Newtonian

fluid

Bingham

fluid

A B

C D

techniques at high-temperature (low crystallinity and low 

viscosity) and low-temperature (high crystallinity and high 

viscosity) ends of the measurement spectrum (fig. 7C; for 

example, Shaw, 1969; Murase and McBirney, 1973). Another 

type of in situ analysis involved settling of olivine phenocrysts 

through lava flows of different types (Rowland and Walker, 

1987, 1988). This approach suggested a progressive increase 

in viscosity as lava morphology changes from smooth 

pāhoehoe (600–1,500 Pa·s) to rough pāhoehoe (6,000 Pa·s) to 

“toothpaste” lava (12,000 Pa·s). These values are about an order 

of magnitude higher than those calculated using measured melt 

temperatures and crystal contents (fig. 7A; Riker and others, 

2009). For this reason, estimates of the apparent viscosity 

controlling the threshold transition between pāhoehoe and ‘a‘ā 

shown in figure 5 are about an order of magnitude lower than 

that estimated on the basis of crystal-settling calculations.

Early workers also recognized that crystal-bearing magmas 

were not simple Newtonian fluids, and Bingham (1922) first 

Figure 7. Lava flow rheology. A, Apparent viscosity (diamonds) versus distance along length of lava flow from 1984 Mauna Loa eruption, calculated using 

Jeffrey’s equation and data from Lipman and Banks (1987). Core lava viscosity (circles), calculated from measured glass temperatures and crystallinity 

(from Riker and others, 2009), is typically at least an order of magnitude lower than bulk apparent viscosity. B, Microphenocryst crystallinity of lava emerging 

from main vent of 1984 Mauna Loa eruption. Increasing crystallinity most likely reflects gas loss from magma during transport downrift (redrafted from 

Crisp and others, 1994). C, Relative viscosity (viscosity ratioed to pure-liquid values) versus critical, or maximum, crystallinity, Φm (redrafted from Costa and 

others, 2009). Note abrupt increase in viscosity close to critical crystallinity (“jamming point”). Absolute value of critical crystallinity depends on specifics of 

crystal population (distribution of crystal sizes, shapes, and orientations). D, Yield strength versus crystal content Φ beyond threshold crystallinity Φc for 

yield-strength onset; curve is a power-law function with a constant D=5×106 Pa (modified from Castruccio and others, 2010). 
366 Characteristics of Hawaiian Volcanoes

suggested that his concept of yield-strength fluids (fig.7D) 

might extend to lava. The idea that lava flows might have 

Bingham rheologies was extended by the burgeoning planetaryvolcanological community in the 1970s. Hulme (1974) used 

analogue experiments to construct a model for the flow of a 

Bingham fluid on a slope. Using this model, he formulated a 

theoretical relation between channel formation and yield strength, 

where yield stress is defined as the minimum stress required 

for a homogeneous crystal-liquid suspension to flow. He tested 

his model using observations on lava flows from Mount Etna 

(Hulme, 1974) before using it to infer lava rheology on the Moon 

and on Mars (Hulme, 1976; Hulme and Fielder, 1977). 

Subsequently, Hulme’s model has been applied to 

Hawaiian lava flows (for example, Fink and Zimbelman, 

1986; Moore, 1987), where measured channel and levee 

dimensions suggest apparent yield strengths ranging from 0 

to 5,000 Pa. Significantly, however, field estimates of yield 

strength in Hawaiian lava flows are complicated by many of 

the same problems that affect viscosity estimates (for example, 

Griffiths, 2000; Kerr and others, 2006). For this reason, 

field measurements based on lava flow properties should be 

considered apparent (or effective) values, and comparison with 

laboratory measurements should be made with caution. 

In the laboratory, the onset of yield strength requires 

development of a “touching framework” of crystals (Kerr and 

Lister, 1991) that can bear stress, such that the crystal-melt 

suspension develops a viscoplastic rheology (Pinkerton and 

Sparks, 1978; Robertson and Kerr, 2012). At this point, the 

lava will cease to flow if the shear stress is sufficiently low or 

may tear rather than deform ductily under the imposed stress 

of continued downslope flow. The crystal volume fraction at 

which this transition occurs strongly depends on crystal shape 

and orientation (for example, Philpotts and others, 1998; 

Philpotts and Dickson, 2000; Hoover and others, 2001; Saar 

and others, 2001). Once the threshold crystallinity is achieved, 

yield strength increases as a power law function of crystal 

volume fraction (fig. 7D ; Castruccio and others, 2010). 

Crystal shape also controls the maximum crystal volume 

fraction at which suspensions can continue to flow (generally 

designated the maximum packing fraction φm; Costa and others, 

2009; Mueller and others, 2010, 2011). The deformation (shear) 

rate also affects suspension rheology by changing the spatial 

arrangement of crystals. For this reason, particle-melt suspensions 

are commonly modeled using a relation between applied stress 

(τ) and strain rate (γ) appropriate for Herschel-Buckley fluids (for 

example, Pinkerton and Norton, 1995). This treatment allows 

three fit parameters: the consistency, K (a measure of viscosity, 

determined by fitting φm for the suspension); the yield strength, 

τy; and the flow index, n, a measure of the extent to which the 

suspension is shear thinning or shear thickening. Also important, 

however, is the size distribution of crystals (Probstein and others, 

1994; Castruccio and others, 2010; Cimarelli and others, 2011), 

which has not yet been fully incorporated into rheologic models 

(see review by Mewis and Wagner, 2009). 

The presence of bubbles also affects the rheology of lava, 

although the magnitude of the effect is much less than that of 

crystals. Field observations suggest that bubble-rich lava can 

behave either more (Lipman and Banks, 1987) or less (Hon and 

others, 1994) fluid than its bubble-poor counterpart, depending 

on flow rate. These field observations are supported by laboratory 

experiments (Rust and Manga, 2002) and models (Pal, 2003; 

Llewellin and Manga, 2005) that demonstrate the relation 

between suspension viscosity and capillary number, Ca, which 

is a measure of the extent to which the bubbles deform during 

flow. Ca is defined as μV/σ, where μ is the melt viscosity, V is the 

characteristic velocity, and σ is the interfacial tension between 

the gas and liquid phases. Thus, bubbles increase viscosity when 

strain rates are low (bubbles are undeformed) and decrease 

viscosity when strain rates are high (bubbles deform). 

No models for magmatic systems account for the effects 

of both bubbles and crystals because of difficulties in modeling 

more than one suspended phase (for example, Tanner, 2009). 

Recent experiments on three-phase materials (both analogue 

and natural) indicate complex rheologies that may include 

both thixotrophic and viscoelastic behavior (for example, 

Bagdassarov and Pinkerton, 2004; James and others, 2004). 

Taken as a whole, the linkage between laboratory- and fieldscale controls on rheology has been advanced by studies in 

Hawai‘i, but further research is needed to fully characterize the 

rheology of active lava flows.

Kinetics of Phase Change

Interest in the bulk properties (temperature and rheology) of 

lava led naturally to an interest in the bubbles and crystals present 

within the melt. Bubbles form in response to depressurization 

and, once formed, can move through fluid lava; for this reason, 

the bubble content of lava can be used to monitor gas loss during 

eruption and emplacement. Crystal formation is sensitive to the 

rate and extent of lava cooling, as well as to stirring (for example, 

Emerson, 1926); thus, changes in the crystal population can be used 

to monitor the thermal and dynamical evolution of flowing lava.

Bubbles

Bubbles (or their frozen equivalents, vesicles) are nearly 

ubiquitous in the products of volcanic eruptions. Since the initial 

volatile composition of Hawaiian magma does not vary substantially, the bubble population in eruptive products is determined 

primarily by the vesiculation history (controlled by the decompression path) and subsequent patterns of bubble escape. In 

general, pyroclasts formed in Hawaiian lava fountains contain 

more and smaller bubbles than lava flows with the same total 

vesicularity generated by purely effusive activity (for example, 

Cashman and others, 1994; Mangan and Cashman, 1996; Stovall 

and others, 2011); this variation is inferred to reflect higher rates 

of decompression during fountaining eruptions (as measured by 

observed variations in mass eruption rate, fig. 8A).

Early workers recognized five different styles of vesiculation, two of which characterized lava fountain eruptions (see 

Mangan and others, this volume, chap. 8) and three possibly 
A Century of Studying Effusive Eruptions in Hawai‘i 367

related to different lava emplacement styles (Hitchcock, 1911); 

these categories include (1) “ordinary lava streams,” with ≤60 

volume percent elongate vesicles; (2) “spherically vesiculated 

lava,” with 30–60 volume percent bubbles in the upper parts of 

the flow; and (3) “the scum of the lava, which is often troublesome because one breaks through it in walking,” with 65–75 

volume percent bubbles. The first vesiculation style, which is 

considered to be diagnostic of ‘a‘ā flows (Macdonald, 1953; 

fig.2D), reflects both the high crystallinity (viscosity) and large 

shear strains applied to lava that has been transported through 

open channels (for example, Cashman and others, 1999; Soule 

and others 2004; Riker and others, 2009). The second style, 

which is also common, is diagnostic of inflated pāhoehoe flows 

(for example, Hon and others, 1994; Katz and Cashman, 2003). 

Lava flows

Lava fountains

8

6

4

2

2 3 4 5 6

Log MER, in kilograms per second

Log vesicle number

per cubic centimeter

Growth Nucleation

Rate

Supersaturation

Distance, in kilometers

Vesicularity, in percent

0

10

20

30

40

50

60

70

80

0 2 4 6 8 10 12

Mean vesicle size,

in millimeters

0

0.2

0.4

0.6

0.8

1.0

}

{ LF

LF

A

B

In these flows, the vesicular upper crust forms while the flow is 

actively inflating. In contrast, thin (rapidly quenched) pāhoehoe 

lobes may show a uniform distribution of spherical vesicles 

throughout (spongy pāhoehoe; Walker, 1989; fig. 2C). The third 

style, lava “scum,” is termed shelly pāhoehoe (Jones, 1943; Wentworth and Macdonald, 1953) and is common near eruptive vents 

(for example, VL 502). Shelly pāhoehoe forms by gas accumulation beneath thin, rapidly formed crust (Swanson, 1973).

Later workers have added to the lexicon of vesicle textures 

by identifying “flow pumice” and “P-type” pāhoehoe. Flow 

pumice is described as a tan, highly vesicular (~75 volume 

percent vesicles) glass skin on short flows adjacent to fissures. 

For example, pumice-surfaced pāhoehoe that formed during 

the 1942 eruption of Mauna Loa was “restricted to gushes 

of lava which traveled less than a quarter of a mile from 

their vents” and inferred to have formed by rapid (“nearly 

explosive”) vesiculation during emplacement (VL 502). Nearvent flow pumice produced during the 1859 eruption of Mauna 

Loa approaches reticulite in both vesicularity and structure, 

supporting this interpretation. P-type pāhoehoe is characterized 

by the presence of pipe vesicles in the lower parts of the flow 

(Wilmoth and Walker, 1993) and commonly occurs as breakouts 

from hummocky tumuli, where lava stagnates and partially 

degasses (Swanson, 1973). Pipe-vesicle-bearing flows tend to 

be dense, with “blue glassy surfaces,” probably a consequence 

of bubble loss during temporary lava residence within 

hummocky flows before final emplacement (Hon and others, 

1994). The pipe vesicles grow inward from the cooling front 

as “cold fingers” (for example, Philpotts and Lewis, 1987), and 

although commonly seen at the flow base, they may also form 

radially around flow margins. This observation suggests that the 

use of pipe vesicles as flow-direction or paleoslope indicators 

(Waters, 1960; Walker, 1987) must be done with caution 

(Peterson and Hawkins, 1972; Swanson, 1972).

To date, field measurements also provide most of the 

constraints on the vesiculation kinetics of basaltic magmas 

(for example, Mangan and others, 1993, this volume, chap. 8; 

Mangan and Cashman, 1996), because the rapidity of bubble 

formation makes laboratory experiments challenging (see 

Murase and McBirney, 1973; Pichavant and others, 2013). Direct 

measurements of bubble populations in flowing lava have been 

made along both the 1984 Mauna Loa channelized flow (Lipman 

and Banks, 1987) and robust lava tubes that formed during 

the Pu‘u ‘Ō‘ō eruption (Mangan and others, 1993; Cashman 

and others, 1994). Both datasets show a downflow decrease in 

bubble content with distance from the vent that is consistent with 

observations of gas escape from flow surfaces (for example, fig. 

8B). In lava tubes, vesicularity decrease is accomplished primarily 

by loss of large bubbles, as observed in lava tubes, described as 

“an orange-hot cavity with a golden river sweeping by underneath, 

little bubbles continually breaking the surface of the glowing 

stream, and adding gas to the evenly brilliant walls” (VL 345). 

Bubbles frozen into the growing upper crust of inflated lava flows 

can also record pressure changes within the tube during flow 

emplacement (Cashman and Kauahikaua, 1997; see subsection 

below entitled “Lava Tube Formation”). 

Figure 8. Vesicularity of Hawaiian lava. A, Plot showing bubble number 

density in lava-flow and lava-fountain samples. Higher bubble number 

densities record higher rates of bubble nucleation, which are associated 

with more rapid magma ascent (higher decompression rates), as inferred 

from estimated mass eruption rate (MER); data from Cashman and 

others (1994), Mangan and Cashman (1996), and Stovall and others 

(2010). B, Plot of lava vesicularity (gray diamonds) and mean bubble size 

(open diamonds) versus distance along two Kīlauea lava tubes. Steady 

decrease in both parameters reflects escape of gas bubbles from lava free 

surface. Redrafted from Cashman and others (1994). Brackets, ranges 

of vesicularity and mean bubble size reported by Mangan and Cashman 

(1996) and Stovall and others (2010) for lava-fountain (LF) samples.
368 Characteristics of Hawaiian Volcanoes

Crystals

As reviewed above, Hawaiian lavas are commonly 

erupted at near-liquidus temperatures; for this reason, they can 

crystallize extensively, particularly during flow through open 

channels. Early 20th century crystallization studies were aimed 

at defining conditions of ‘a‘ā and pāhoehoe formation (for 

example, Emerson, 1926). Then, five decades later, a renewed 

interest in crystallization kinetics accompanied the advent of the 

semiconductor industry (for example, Kirkpatrick, 1981) and the 

collection of lunar samples (for example, Dowty, 1980; Basaltic 

Volcanism Study Project, 1981). Hawaiian lava lake samples, in 

particular, provided a well controlled natural laboratory for these 

studies (see Helz and others, this volume, chap. 6; Mangan and 

others, this volume, chap. 8). 

More recently, protracted lava flow eruptions have allowed 

relatively easy access to lava flows. Analysis of quenched 

samples from these flows can be used to link crystallization 

conditions directly to cooling rates (for example, Crisp and 

others, 1994; Cashman and others, 1999; Soule and others, 

2004; Riker and others, 2009). Typical Hawaiian lava is erupted 

at temperatures of 1,150–1,170 °C, when the melt is saturated 

with both plagioclase and pyroxene. Under these conditions, 

measured cooling rates of ~0.005 °C/s along near-vent open 

channels drive crystallization at a rate of ~0.005 to 0.01 volume 

percent per second (18–36 volume percent per hour). At these 

high cooling rates, crystallization occurs primarily by nucleation 

of new crystals rather than by growth of existing crystals (fig. 

9A). The dominance of crystal nucleation is illustrated by a 

steady increase in the measured number of both plagioclase 

and pyroxene crystals with increasing crystal volume fraction 

(fig.9B), as well as by patterns of crystal-size distributions 

(CSDs). Samples collected along a single lava channel on the 

same day show parallel CSD trends (fig. 9C). Here, the slope 

of the line provides a measure of dominant crystal size, and the 

area under the line is the total number of crystals; therefore, 

parallel trends reflect addition of crystals of the same (small) 

dominant size (for example, Cashman and others, 1999). 

In contrast, slow cooling accompanying lava transport 

through well-insulated lava tubes or solidification of stable 

lava lakes promotes crystal growth over crystal nucleation. 

Growth-dominated crystallization is manifested by the 

maintenance of constant crystal numbers with transport 

distance (fig. 9A) and total crystallinity (fig. 9B). CSDs 

also show patterns characteristic of growth-dominated 

crystallization, such that CSDs pivot around a point (fig. 

9D) rather than showing the parallel trends plotted in figure 

9C. Fanning CSDs record increases in dominant crystal size 

(inversely proportional to slope) with increasing volume 

fraction at either constant, or even decreasing, total crystal 

numbers (Cashman and Marsh, 1988).

Quenched samples can also provide insight into 

the relation between crystallization and flow-surface 

morphologies, which, in turn, place constraints on flow 

rheology (for example, Cashman and others, 1999; Soule 

and others, 2004; Riker and others, 2009). Such studies 

suggest that smooth pāhoehoe surfaces can be maintained 

to groundmass crystallinities of <~20 percent if shear rates 

are sufficiently low. Flow surfaces can start to develop 

transitional (rough) surface characteristics at groundmass 

crystallinities as low as ~15 percent if shear rates are higher; 

groundmass crystallinities >35 percent typically have fully 

formed ‘a‘ā textures regardless of shear rate (fig. 10A). These 

observations can be explained by (1) a shear-rate-dependent 

onset of yield strength at 10–20 volume percent crystals 

and (2) a drastic increase in viscosity at ~35 volume percent 

crystals, suggesting a relatively low value of φmax in Hawaiian 

lavas. This low value of critical crystallinity (compare with 

fig. 7C) is consistent with the high anisotropy of groundmass 

plagioclase crystals. Support for this interpretation comes from 

observations of lava flows that maintain pāhoehoe surfaces 

to moderately high crystal contents (~50 volume percent). 

These flows (such as the picrites common on Kīlauea and 

Mauna Loa’s rift zones and the “cicirara” [chickpea] lavas of 

17th century Mount Etna eruptions) are dominated by nearisotropic phenocrysts, which should have a “jamming” point 

in excess of 40–50 volume percent. 

Together, the studies outlined above show that (1) 

channelized Hawaiian lavas crystallize rapidly during early 

stages of transport; (2) rapid crystallization produces numerous 

small crystals that create an abrupt increase in the effective 

viscosity of the lava; and (3) subsequent deformation of the 

magma occurs by tearing rather than ductile flow, thereby 

creating the rough surface that is characteristic of ‘a‘ā. 

Solidified ‘a‘ā flows are finely crystalline throughout (fig. 10B), 

testifying to the efficiency of stirring in these flows (Griffiths 

and others, 2003; Cashman and others, 2006). In contrast, 

tube-fed surface flows commonly have glassy pāhoehoe 

surfaces, a consequence of minimal syntransport cooling 

and crystallization. Postemplacement crystallization textures 

mirror cooling rates, so that glassy pāhoehoe surfaces rapidly 

transform to, first, finely and then coarsely crystalline interiors 

(fig. 10B; Oze and Winter, 2005) with declining rates of cooling. 

Contrasting crystallization textures preserved in channelized 

and tube-fed lava flows thus provide critical information on both 

syneruption and postemplacement cooling; the resulting textural 

contrasts are particularly useful for interpretation of drill-core 

samples that lack an areal context (Katz and Cashman, 2003).

Dynamics of Lava Flow Emplacement

Understanding where lava flows are likely to go, how far a 

given flow will travel, and how quickly it will advance are questions that are critical to the assessment of lava flow hazard. The 

answers to these questions differ for the short, high-effusionrate channelized lava flows from Mauna Loa and the long-lived 

pāhoehoe flow fields that characterize current (and previous; for 

example, Clague and others, 1999) eruptive activity at Kīlauea. 

In this section, we review the physical controls on both styles of 

flow emplacement and illustrate some of the ways in which flow 

models are being used for hazards assessment.
A Century of Studying Effusive Eruptions in Hawai‘i 369

ML1859 ‘a‘ā

0

1,000

2,000

3,000

4,000

0 10 20 30 40 50

Distance from vent (km)

Plagioclase number

per square millimeter

ML1859 pāhoehoe

0

2,000

4,000

6,000

0 0.1 0.2 0.3

Pāhoehoe ‘A‘ā

Crystal volume fraction

Crystal number

per square millimeter

Pyroxene

Plagioclase

ML pāhoehoe

2σ

ML 1984 ‘a‘ā

12

20

14

10

16

18

0 0.01 0.02 0.03 0.04

Increasing time

0.3 km

4.9 km

1%

16%

0

12

2

14

4

10

6

8

0 0.05 0.1 0.15 0.2 0.25 0.3

13% 26%

39%

45%

28%

Crystal size, in millimeters

Ln crystal number

per quad millimeter

Increasing time

A B

C D

Ln crystal number

per quad millimeter

Crystal size, in millimeters

EXPLANATION

Figure 9. Crystal textures in Hawaiian lava flows. A, Plot of plagioclase microlite number density versus distance along paired ‘a‘ā (blue circles) 

and pāhoehoe (open squares) flows from 1859 Mauna Loa (ML) eruption. Two contrasting trends represent nucleation (‘a‘ā) and growth (pāhoehoe) 

dominated crystallization (redrafted from Riker and others, 2009). B, Plot showing microlite number densities of plagioclase (blue dots and circles, lines) 

and pyroxene (red squares, line) versus changes in crystallinity. Data from small ‘a‘ā flow from the 1997 Kīlauea eruption. Closed symbols, samples 

with pāhoehoe flow surfaces; open symbols, samples with transitional flow surfaces (Cashman and others, 1999). Lines and shaded field are data from 

paired ‘a‘ā and pāhoehoe flows from the 1984 Mauna Loa eruption (from Riker and others, 2009). C, Plot of plagioclase crystal size distributions (CSDs; 

for example, Cashman and Marsh, 1988) from Episode 16 of the Pu‘u ‘Ō‘ō eruption, collected at different sites along main lava channel. Explanation 

indicates distance from vent. Parallel trends are another reflection of nucleation-dominated crystallization in channelized ‘a‘ā flows. D, Plagioclase 

CSDs of samples collected through crystallization interval of Makaopuhi lava lake (reanalyzed samples from Cashman and Marsh, 1988). Plagioclase 

crystallinity is labeled for each line; fanning CSDs show growth-dominated crystallization trends.

Figure 10. Properties of selected lava samples collected along individual open-channel flows during episodes 1–20 of the Pu‘u ‘Ō‘ō eruption (from 

Folley, 1999 and Cashman and others, 1999). A, Plot of total groundmass crystallinity of samples with different surface textures. B, Plot of crystal 

number density versus normalized depth through typical pāhoehoe (circles) and ‘a‘ā (squares) flows; dashed lines show approximate trends of data 

points. In solidified flows, crystal number density varies inversely with crystal size (redrafted from Katz and Cashman, 2003).

0.8

0.6

0.4

0.2

0

Volume fraction crystals

1.0

0.8

0.6

0.4

0.2

0

0 1 2 3 4

Normalized depth in flow

Log crystal number per square millimeter

A B

Pāhoehoe

Transitional

‘A‘ā

Pāhoehoe

‘A‘ā
370 Characteristics of Hawaiian Volcanoes

Open-Channel Flows

When lava effusion rates are moderate to high, 

incandescent “rivers” of lava flow downhill, confined by 

lateral levees of lava rubble. In Hawai‘i, these channelized 

flows are typically emplaced over time scales of hours (for 

example, 1974 Kīlauea; Lockwood and others, 1999) to 

weeks (for example, 1984 Mauna Loa; Lipman and Banks, 

1987). The first detailed measurements on open-channel flows 

(including channel velocities, flow volume, and effusion rate) 

were made during the 1919 and 1926 Mauna Loa eruptions 

(VL 480). As described previously, these eruptions also 

saw the advent of aerial observations, which permitted not 

only descriptions of the plan-form geometry of braided lava 

streams, but also details of flow surfaces. With the increase in 

real-time observations came questions about the construction 

of, and flow through, lava channels, as well as about the 

evolution of lava flux and channel geometries over time; these 

processes must be understood for predictions of flow length 

and aerial coverage. Although many parameters important for 

characterizing channel-fed flow are now routinely measured, 

some key measurements, such as the depth of lava flowing in 

channels, are still poorly constrained. 

Morphology of Channelized Flows

Channelized flows commonly initiate from fissure vents, 

from which they travel as broad lava sheets that focus into 

lava channels after a flow of tens to hundreds of meters. Early 

proximal flows form anastomosing channels of hot, fluid lava 

with thin surface crusts and small marginal levees (fig. 11A). 

Flow away from the source vents causes rapid cooling, crust 

formation, and continued levee construction. Detailed observations of the 1984 Mauna Loa lava flow showed that stable 

channels form from a zone of dispersed flow at the propagating flow front through a transitional zone (Lipman and Banks, 

1987). The stable channel may evolve over time as the flow 

focuses by inward solidification and as channel surges and (or) 

blockages create overflows (fig. 11B). 

Open-channel flows are distributary, and the mass of 

flowing lava decreases along the channel from the vent to the 

flow front. This process is illustrated by measurements made 

along the 1984 Mauna Loa channel on a single day (April 

4, 1987; fig. 11C), which show that bulk (lava plus bubbles) 

volumetric flow rate through the channel dropped by a factor 

of 5 over a distance of 15 to 20 km. This decrease was caused, 

in part, by loss of bubbles along the flow; however, even 

when corrected for changes in vesicularity, the volumetric 

flux decreased by a factor of 3. This loss of volume shows 

the extent to which lava channels are prone to mass loss by 

both overflows and storage within stagnant or near-stagnant 

marginal parts of the channel system (fig. 11B). 

Overflows form when the channel is constricted or 

blocked by surface crusts or rafted accretionary lava balls 

or when temporary increases in flux exceed the channel’s 

carrying capacity. Overflows may develop into new flow 

branches if sufficiently sustained; for this reason, channel 

systems are typically distributary in plan form (fig. 4A). 

Individual channels generally widen as slopes decrease and 

narrow as effusion rates decline (Kerr and others, 2006). 

Posteruption surveys of the 1984 Mauna Loa flow showed 

channel floors to be close to preflow surfaces, suggesting that 

initial flow material is eroded as channels mature, but that 

erosion does not ordinarily extend into older, colder rock. An 

exception may be the ca. 1800 Ka‘ūpūlehu lava flow from 

Hualālai (fig. 1), where channels show evidence of mechanical 

erosion during transport of large dunite xenoliths (Kauahikaua 

and others, 2002). 

Figure 11. ‘A‘ā channel formation. A,B, Typical cross sections through 

early and later flow stages. Early flow (A) is characterized by outer 

levees surrounding inner fluid lava, which is separated from levees by 

a zone of sluggish lava flow. Later channel (B) has evolved to include 

solidified lava separating flowing core lava from levee; an inner channel 

has also formed, separated from original channel by overflow levees 

(redrafted from Lipman and Banks, 1987). C, Volumetric flux versus 

distance along lava channel from Mauna Loa eruption on April 4, 1984. 

Apparent flux is calculated from measured flow width, depth, and 

velocity; dense rock equivalent (DRE) flux is corrected for variations in 

lava vesicularity. Data from Lipman and Banks (1987). 

Early levees

Overflow levees

Rapid flow

Sluggish flow Dense

solidified lava 

A

B

C

0 5 10 15 20

Distance, in kilometers

Mass flow rate,

in cubic meters per second

0

200

400

600

800

1,000

Apparent flux

DRE flux
A Century of Studying Effusive Eruptions in Hawai‘i 371

1

10

100

1 10 100 1,000

Flow length, in kilometers

Effusion rate, in cubic meters per second

Mauna Loa

Mauna Loa (ocean limited)

Kīlauea

Kīlauea (1983–86)

Walker’s upper limit

Walker’s lower limit

200 hours

100 hours

Figure 12. Plot of the lengths of channelized Hawaiian lava flows as a 

function of average volumetric flow (effusion) rate. Lines, original trends 

identified by Walker (1973; dashed lines labeled “Walker’s upper limit” and 

“Walker’s lower limit”) and theoretically defined cooling-rate limits (solid 

lines; from Pinkerton and Wilson, 1994). “Ocean limited” flows reached the 

ocean and therefore are minimum flow distances (modified from Riker and 

others, 2009). Red diamonds represent short-duration (<12 hours) lava 

flows erupted from Kīlauea in July (short) and December (long); the latter 

was confined. Similarly, blue squares represent Mauna Loa eruptions of 

~3 weeks' duration in 1984 (short) and 1859 (long); the latter were confined.

Cooling Along Lava Channels

Radiative heat loss from exposed lava surfaces is responsible for the bulk of the heat loss from most lava channels. 

Application of glass geothermometry shows that initial cooling 

rates may be as high as 0.01–0.03 °C/s where surface crusts are 

very thin and eruptive temperatures high (Cashman and others, 

1999; Riker and others, 2009). Maintenance of exposed fluid 

lava along the channel margins thus requires efficient convection from the flow interior (Griffiths and others, 2003). Cooling 

rates drop to ~0.005 °C/s along medial to distal ‘a‘ā channels 

(Crisp and others, 1994; Cashman and others, 1999), reflecting 

both the formation of an insulating ‘a‘ā crust and buffering of 

lava temperatures by latent heat of crystallization.

Controls on Flow Length and Advance Rate

The length of a simple channelized lava flow will increase 

with lava effusion rate if the maximum flow length is controlled 

by cooling, if cooling rates are constant, and if effusion rate 

controls flow velocity (for example, Walker and others, 1973; 

Pinkerton and Wilson, 1994; Harris and Rowland, 2009). Alternatively, flow length may be limited by eruptive volume (flow 

duration), such that the flow does not reach its cooling-limited 

extent. In general, short duration flows tend to be limited in 

length by lava supply (“volume limited”), whereas the lengths 

of long-lived flows are limited by cooling (“cooling limited”). 

Most Hawaiian lava flows are not simple, in this sense, 

however, as illustrated by the lack of correlation between 

channelized-flow lengths and effusion rates, even for flows of 

similar duration (fig. 12). There are several possible explanations for this. First, the rheology of Hawaiian lavas at the vent 

may vary widely with bulk composition, temperature, crystallinity, and bubble content (Riker and others, 2009). Additionally, 

the high fluidity of Hawaiian lavas makes them susceptible to 

topographic confinement that will promote flow lengthening 

(Soule and others, 2004), whereas channel bifurcations caused 

by topographic obstacles create multiple parallel channels that 

can limit individual flow-lobe lengths (Lockwood and others, 

1987). Finally, flow fields generally widen rather than lengthen 

when magma supply is unsteady and flows are emplaced as 

discrete events (Guest and others, 1987; Wolfe and others, 1988; 

Kilburn and Lopes, 1991; Heliker and others, 1998). 

Effusion rate does exert a fundamental control on the 

initial advance rates of channelized lava flows (fig. 13A; 

Rowland and Walker, 1990), consistent with treatment of the 

initial stages of flow emplacement as Newtonian (or Bingham) 

fluid with constant viscosity (Takagi and Huppert, 2010). 

This relation is nicely illustrated using data from recent wellobserved eruptions, where initial effusion rates for channelized 

flows have ranged from ~25 to 1,000 m3/s and corresponding 

initial rates of flow advance have varied from <0.02 to 3–4m/s 

(fig.13B; Kauahikaua and others, 2003). Importantly, both 

datasets suggest that slope plays only a secondary role in 

controlling initial rates of flow advance.

These initial flow advance velocities form a trend that 

parallels, but is offset toward higher velocities from, the data 

compiled by Rowland and Walker (1990) on older flows 

(fig. 13A). This discrepancy probably reflects the fact that 

the observational data for older flows were obtained mainly 

from distal sites, where flow advance rates are lower because 

of decreasing flux (from losses along the channel), crust 

formation (Kerr and Lyman, 2007), and increases in internal 

lava viscosity because of cooling-induced crystallization. 

The observed correspondence between flow-advance 

rate and volumetric flow rate suggests that advance rates 

inferred from flow features, such as runup heights on tree 

molds (Moore and Kachadoorian, 1980), superelevations on 

channel bends (Heslop and others, 1989), clinker size, and 

lava crystallinity, can be used to estimate effusion rates from 

older eruptions (for example, Soule and others, 2004). These 

data also highlight the importance for hazard assessment of 

obtaining accurate estimates of effusion rate during early 

stages of eruptive activity (see next subsection “Real-Time 

Flow Monitoring and Hazards Assessment”).

Observational data on lava flow advance also show that 

the position of eruptive vents relative to local populations 

affects the relative impact of lava flow eruptions, because lava 

flows typically advance more slowly when farther from their 

source vents (Kilburn, 1996). For this reason, eruptions from 

low-elevation flank vents pose much greater hazards than those 

from vents high on the rift zones of Kīlauea and Mauna Loa, 
372 Characteristics of Hawaiian Volcanoes

1

-1

5

4

3

2

0 1 2 3 4

P A

Log of flow-front velocity in meters per hour

Log of volumetric flow rate,in cubic meters per second

Mauna Loa

Kīlauea

0 20 40 60 80 100

0

20

25

15

10

5

>350 m3/s

Time, in hours

Distance, in kilometers

A B

100–250 m3/s

50–100 m3/s

250–300 m3/s

25–50 m3/s

2,000 m/hr

1,000 m/hr

300 m/hr

125 m/hr

25 m/hr

Figure 13. Effusion-rate controls on flow advance rates. A, Volumetric flow rate versus flow-front velocity estimated for historical lava flows 

(diamonds and squares; data from Rowland and Walker, 1990) and constraints from well-observed recent flows (colored boxes; data from 

Kauahikaua and others, 2003). Offset between two datasets may reflect either underestimates of historical flow-advance rates or overestimates 

of historical flow-front velocities. P=pāhoehoe, A=‘a‘ā. B, Distance versus time for recent lava flows (modified from Kauahikaua and others, 2003). 

Individual fields are labeled for volumetric flow rates; dashed lines, flow-front velocities. Colors in A and B refer to same data.

as illustrated by the attempts at barrier construction during the 

1955 and 1960 Kīlauea eruptions. Additionally, unusually rapid 

advance of lava flows was reported for the eruptions of Hualālai 

in 1801 and Kīlauea in 1823. Both of these lava flows were 

erupted from vents near the coast (fig. 1); thus, these accounts 

may derive more from the proximity of the eruptive vents to 

observers (Kauahikaua and others, 2002; Soule and others, 

2004) than from any unusual properties of the erupted lava 

(Baloga and others, 1995). 

Real-Time Flow Monitoring and Hazard Assessment

A compelling reason for studying the formation and 

evolution of lava channels is to construct predictive models of 

flow paths, advance rates, and areal extents. Existing models 

of lava flow behavior include empirical relations (for example, 

Walker, 1973; Pinkerton and Wilson, 1994), probabilistic 

models of inundation (for example, Kauahikaua and others, 

1995, Kauahikaua and Trusdell, 1999; Crisci and others 2010), 

thermal models (Crisp and Baloga, 1990, 1994), rheologic 

models (Hulme, 1974; Dragoni and Tallarico, 1994; Kilburn, 

2004), parametric models (Griffiths and others, 2003; Lyman 

and others, 2005, Lyman and Kerr, 2006; Cashman and others, 

2006; Kerr and others, 2006; Robertson and Kerr, 2012), and 

(one-dimensional) coupled fluid dynamical and thermal models 

(Harris and Rowland, 2001). The array of approaches reflects 

the complexity of the processes involved. Key data required for 

all models, however, are volumetric lava flux (effusion rate) at 

the vent and high-resolution topography. 

Average lava effusion rates, by definition, can be 

determined only in hindsight, using measured flow volumes and 

known eruptive durations (for example, VL 508, 509; Tilling 

and others, 1987; Wolfe and others, 1988). More useful during 

eruptive crises, though more difficult to obtain, are daily, or even 

hourly, measurements of instantaneous effusion rates. Effusion 

rates can be estimated for flow through individual lava channels 

if flow velocities and channel geometries are well known (for 

example, Lipman and Banks, 1987). These estimates, however, 

are complicated by uncertainties in the bulk density of the 

erupted lava, by difficulties in estimating channel depths, and 

by the need to measure different branches of complex channel 

networks simultaneously. For this reason, interest has been 

increasing in creating remote sensing tools for effusion rate 

measurements (reviewed in Harris and others, 2007).

Digital topography (DEMs) forms the basis of all 

probabilistic models of flow inundation and is required for 

statistical modeling of, and response to, ongoing eruptions. 

In Hawai‘i, GIS analysis of lava sheds, or topographic 

areas within which lava will be confined for specific vent 

positions (Kauahikaua and others, 1995, 1998, 2003), have 

proven effective in both general hazard mapping and guiding 

emergency-response efforts (fig. 4). Hazard maps for a specific 

region can be created by an analysis of “paths of steepest 

descent” within lava sheds (for example, Costa and Macedonio, 

2005; Kauahikaua, 2007). Major challenges to developing 

predictive models using this approach include obtaining the 

DEM resolution required to anticipate the advance of thin 

flows and the need for frequent updating of DEMs to account 
A Century of Studying Effusive Eruptions in Hawai‘i 373

for changes in local topography created by the lava flows 

themselves (for example, Mattox and others, 1993). From this 

perspective, new high-resolution DEMs available from repeat 

lidar surveys provide the optimal resolution (for example, 

Favalli and others, 2010), but such surveys are generally 

impractical from the perspective of either cost or time required 

for data processing (reviewed by Cashman and others, 2013). 

More cost effective are photogrammetric surveys (for example, 

James and others, 2010) that provide near-real-time digital 

elevation data on advancing flow fronts, although these surveys 

are generally limited in spatial coverage.

From a theoretical perspective, predictive models must 

incorporate the balance between thermal and dynamical 

controls on flow advance rate (for example, Griffiths, 2000), 

as well as considerations of the internal organization of lava 

flows, particularly in relation to flow focusing and channel 

formation. Laboratory experiments performed using analogue 

materials have provided useful insights into this problem. Early 

experiments by Hulme (1974) used isothermal viscoplastic 

fluids to explain channel formation; levees were assumed 

to form where lateral flow was inhibited by the yield stress 

of the suspension. Lava flows, however, are not isothermal. 

Experiments that examine the flow of solidifying (cooling) 

Newtonian fluids through preexisting channels define two 

regimes: a “tube” regime at low flow rates and a “mobile crust” 

regime at high flow rates (Griffiths and others, 2003). The 

boundary between these regimes is determined by a critical 

value of the combined parameter θ=ψ(Ra/R0)

1/3, where ψ=U0

t

s

/H0

is the ratio of a surface solidification timescale t

s

 to a shearing 

timescale H0/U0; H0 and U0 are the flow depth and centerline 

surface velocity in the absence of solidification, respectively; Ra

is a Rayleigh number, and R0 is a constant (Griffiths and others, 

2003). This scaling appears applicable to the interpretation of 

basaltic lava flow emplacement in other terrestrial (Ventura and 

Vilardo, 2008) and submarine (Soule and others, 2007) settings. 

In a mobile-crust (open channel) regime, the crust width dc

is always less than the channel width W but increases as dc~W5/3, 

except where the flow accelerates through constrictions, around 

bends, or over slope breaks (Cashman and others, 2006). The 

condition of dc=W defines the tube regime. Recent experiments 

have extended this approach to examine solidifying viscoplastic 

flows (Robertson and Kerr, 2012), which show the same criteria 

as for tube development, but in which a single mobile-crust 

regime is replaced by a progression from a shear- to a plugcontrolled regime with increasing flow rate.

In reality, lava flows rarely occupy predefined channels; 

thus, channel formation must also be considered. The 

laboratory studies presented above have given rise to two 

end-member models of channel and levee formation: (1) 

channel width controlled by cooling and solidification of 

Newtonian fluids (Kerr and others, 2006) and (2) channel 

width controlled by the rheologic properties of an isothermal 

Bingham fluid (Hulme, 1974; Sparks and others, 1976). 

In part, these models derive from differences in the initial 

properties of lava erupted in Hawai‘i and at Mount Etna: 

the dynamics of Hawaiian lava flows, which are erupted at 

near-liquidus temperatures, appear to be governed largely 

by initial conditions that control near-vent cooling and crust 

formation (Griffiths and others, 2003), whereas Etna lavas 

typically are very crystalline and, thus, dominated by the 

constraints of non-Newtonian rheologies (Hulme, 1974). 

These models have yet to be combined to examine channel 

formation in solidifying viscoplastic flows.

Pāhoehoe Flows

Macdonald (1953) provided the first systematic 

description of pāhoehoe flow types. He linked the high 

vesicularity of most pāhoehoe flows to the high fluidity 

of the constituent lava. Styles of pāhoehoe flow advance 

depend on the flow rate and physical properties of the lava 

(Macdonald, 1953; Gregg and Keszthelyi, 2004). Rapidly 

advancing pāhoehoe sheet flows move continuously with a 

rolling motion as the faster upper surface travels over the 

lower part of the flow. In contrast, slowly moving flow lobes 

advance erratically, with times of stagnation accompanied 

by internal inflation. At intermediate flow rates, the growing 

crust is torn or fractured to produce a new flow lobe.

Scientific advances from almost three decades of 

flow mapping at Kīlauea include documentation of (1) the 

importance of flow inflation in pāhoehoe flow fields (Hon 

and others, 1994), (2) the effectiveness of initial flows in 

confining and displacing subsequent flow lobes (Mattox 

and others, 1993), (3) the evolution of lava tube systems 

(Kauahikaua and others, 2003), (4) the relation of feeder 

tubes to flow field development (Mangan and others, 1995a; 

Heliker and others, 1998), and (5) the pressure balance 

between the summit reservoir and feeder dikes (Kauahikaua 

and others, 1996). These processes combine to control the 

morphology of pāhoehoe flow fields.

Formation of Pāhoehoe Flow Fields

The formation of the Kalapana flow field in 1990 

stimulated a major advance in understanding pāhoehoe 

flow emplacement. Here, lava flows advancing into the 

community of Kalapana were mapped in great spatial and 

temporal detail (Mattox and others, 1993). The flows were 

of two dominant types: large primary flows and smaller 

breakouts from the primary flows. The primary flows 

advanced rapidly as a sheet and later developed robust lava 

tubes; as such, they were the main conduit for lava transport 

into the area. Breakouts from the primary flows formed 

secondary flows that advanced slowly and developed tubes 

that were transient; however, the persistence of breakout-fed 

secondary flows meant that their volume eventually exceeded 

that of the primary flows. Additionally, despite erratic advance 

of individual primary and secondary lava flows, the aerial daily 

coverage with lava was approximately constant as long as the 

magma supply remained constant. On the low (2°) slopes, 
374 Characteristics of Hawaiian Volcanoes

where lava flow inflation is the dominant flow emplacement 

mechanism, planimetric area thus provides a better estimate of 

effusion rate than does lava flow length (fig. 14).

Lava Flow Inflation

The detailed examination of primary flow emplacement 

by Hon and others (1994) documented the process by which 

inflated pāhoehoe lava flows form. Lava flows inflate when 

the transport rate of lava from upslope exceeds the advance 

rate at the flow front. Because flows advance more slowly 

on low slopes than on steep slopes, flow inflation is most 

pronounced on low slopes. Lava flow inflation can occur 

either within small lobes or across large sheets. Sheet flow 

inflation causes flows to expand by successive breakouts 

at both the flow front and around sheet flow margins (for 

example, Hon and others, 1994; Hoblitt and others, 2012). 

This pattern of areal expansion explains the relations 

between areal coverage and time in figure 14. 

Actively inflating lava flows have four spatially and 

rheologically distinct layers: a basal crust, a liquid core, and 

a crust comprising a lower, partially solidified viscoelastic 

layer (T=800–1,070 °C) and an upper (cooler) brittle layer 

(fig. 15A). If the lava supply remains constant, both the 

lower and upper crustal layers increase in thickness (Hs) over 

time (t) at a rate controlled by conductive cooling, that is, 

Hs=C√t. Here, C is a constant that Hon and others (1994) 

took to be 0.0779 for the upper crust, on the basis of cooling 

measurements in Makaopuhi lava lake. For application to 

different field situations, this value should be modified, 

depending on both lava vesicularity and local moisture (for 

example, Keszthelyi, 1995b; Cashman and Kauahikaua, 

1997; fig. 15C). The lower crustal layer is assumed to grow 

at 70 percent of the rate of the upper crustal layer because 

of contact with the underlying rock (Hon and others, 1994). 

The relation between the thickness of solidified lava and time 

can be inverted to yield flow duration from the measured 

thickness of the upper crustal layer of actively inflating 

flows (Hon and others, 1994) or solidified flows (Cashman 

and Kauahikaua, 1997). Importantly, the thickness of the 

liquid core (L) appears to stabilize after about 100 hours and 

can therefore be measured if the total flow thickness and 

upper-crustal-layer thickness (Hsu) are known: L=H–1.7Hsu. 

Measured liquid cores in the Kalapana flows did not exceed 

2 m in thickness and were more commonly about 1 m thick 

(Hon and others, 1994; Cashman and Kauahikaua, 1997). 

Small sheet flows commonly maintain fluid-core thicknesses 

of only a few tens of centimeters (fig. 15B; Hoblitt and 

others, 2012).

For a flow to inflate internally, the outer crust must 

either deform or break. Tensile stresses in the brittle crust 

cause it to fracture when the flow of lava contained by the 

crust exceeds the tensile strength of the crust (for example, 

Hoblitt and others, 2012). Successive fractures at either the 

sheet flow margins or central axial crack thus record the 

process of lava flow inflation. Crack formation in inflating 

lava flows has been equated with fracture propagation in 

the formation of columnar joints (for example, Aydin and 

DeGraff, 1988; Grossenbacher and McDuffie, 1995). One 

common feature is internal banding of inflation cracks at a 

centimeter scale, a feature that has been attributed to escape 

of hot gases (Nichols, 1939), lava oozing in from the crack 

(Walker, 1991), or various failure mechanisms in the brittle 

and ductile parts of the flow (Hon and others, 1994). A recent 

study, however, indicates that the banding may instead reflect 

subtle changes in inflation rate, thereby highlighting the 

unsteadiness of many lava emplacement processes (Hoblitt 

and others, 2012). Unsteady inflation is also illustrated by 

patterns of vesicle distributions within the upper crust of 

inflated flows (fig. 15C), where zones of small vesicles 

record times of temporary pressure reduction within the 

tube system because of pauses in lava supply (Cashman and 

Kauahikaua, 1997).

Mapping of structures within inflating pāhoehoe flows 

has also proven important for understanding active lava 

transport systems. Elongate tumuli form over primary 

lava feeder networks, and hummocky tumuli characterize 

areas where tubes have slowed or stalled (Hon and others, 

1994). These inflation features generate a distinctive 

surface morphology that has permitted identification of flow 

inflation in many other terrestrial (Chitwood, 1993, 1994; 

Keszthelyi and Pieri, 1993; Atkinson and Atkinson, 1995; 

Stephenson and others, 1998) and submarine (Applegate 

and Embley, 1992; Gregg and Chadwick, 1996; Chadwick 

and others, 1999, 2001) basaltic lava flow fields, as well 

Time, in days

Area, in tens of thousands of square meters

0 3

0

2

3

1

200

400

600

800

6 9 12 15

Length

Area

Length, in meters

Figure 14. Comparison of variations in length and area versus time for 

1990 Kalapana flow. Curve is approximately linear and provides a measure 

of lava flux (redrafted from Mattox and others, 1993).
A Century of Studying Effusive Eruptions in Hawai‘i 375

Figure 15. Lava flow inflation. A, Measured (dark line) and powerlaw approximated height of inflated pāhoehoe sheet flow versus 

time. Extent of fluid lava core (pink) and lower crust were calculated 

according to constraints of Hon and others (1994). B, Inferred fluid 

lava core versus time (redrafted from Hoblitt and others, 2012). 

C, Long-term flow height of inflated lava flow versus date in 1990. 

Fluctuations in elevation of upper crust are due to temporary pauses in 

lava supply, which caused a reduction in internal pressure of lava tube 

and formation of thin vesicular layers (“small vesicle zones”) within 

solidified crust. General correspondence between vesicular layers 

and loss of lava tube pressurization suggest that textural variations in 

upper crust record dynamics of lava flow emplacement (modified from 

Cashman and Kauahikaua, 1997).

as in some large igneous provinces (for example, Self and 

others, 1996, 1998; Coffin and others, 2000). Recognition 

of inflated flow structures in many flood-basalt lava flows 

has provoked a fundamental re-examination of flood basalt 

emplacement processes.

Lava Tubes

Lava tubes are integral to the formation of pāhoehoe 

flow fields and are a signature feature of Hawaiian 

volcanism. Described by Finch as “one of the most 

interesting and picturesque of volcanic formations” (VL 

82), lava tubes are ubiquitous in pāhoehoe flows of all sizes 

and may occur in ‘a‘ā flows, as well (for example, Dutton, 

1884). Titus Coan first recognized not only the formation 

of lava tubes but also their importance in insulating flows 

and, thus, permitting long-distance transport of lava (Wright 

and Takahashi, 1989). Coan’s observations persuaded Dana 

of the importance of lava tubes, despite his original views 

that lava flows were entirely fissure fed (Dana, 1890). It 

was also recognized that lava tubes could form by surface 

crust formation over a lava channel (river) or from lavafilled cracks (VL 82). Early observers commented on the 

similarities and differences between plan-form-geometry 

tube-fed lavas and rivers, noting that, like rivers, lava follows 

topography but that, unlike rivers, lava tubes tend toward 

distributary systems, such that the discharge diminishes with 

distance from the vent (VL 82). 

Lava Tube Formation

Modern studies of Hawaiian lava tubes began with the 

1969–74 Mauna Ulu eruption, in which tube formation was 

common and easy to observe (for example, Peterson and 

Swanson, 1974; Greeley, 1987; Peterson and others, 1994). 

Observational studies of lava-tube formation during the 

ongoing Pu‘u ‘Ō‘ō eruption have taken advantage of stable 

lava tube systems that have extended several kilometers and 

persisted for many months (for example, Kauahikaua and 

others, 1998, 2003; Orr, 2011). 

Lava tubes are characterized by the presence of a 

stationary solid crust over fluid lava. In active lava channels, 

crust formation may initiate at the vent when effusion rates 

are low, may propagate upchannel from the flow toe when 

effusion rates are higher, or may initiate in the midsection 

of a channel at points of channel constrictions or blockages 

(for example, Greeley, 1987; Peterson and others, 1994; 

Cashman and others, 2006). Lava tubes on steep slopes 

typically have headroom (are unfilled) and so are prone to 

roof collapse. Sites of roof collapse (skylights) permit both 

sampling and direct measurements of flow temperature, 

vesicularity, velocity, and (under the right conditions) depth 

of the lava stream (Cashman and others, 1994; Kauahikaua 

and others, 1998).

0 10 20 30

0.8

0.4

0.3

0

0

Time, in hours

Height, in meters

Upper crust

Fluid lava

Lower crust

Fluid lava

H = 0.257t0.371

A

B

July Aug. Sept. Oct.

Height, in meters

8

6

4

2

0

Small

vesicle zones

C

1990
376 Characteristics of Hawaiian Volcanoes

Lava tubes can also form within inflating pāhoehoe 

flows as sheet flows cool to form enclosing crusts. Inward 

cooling from the sheet flow surfaces focus the flow into 

central lava tubes (fig. 15; Hon and others, 1994). The 

shallow slopes that promote lava flow inflation also inhibit 

lava tube drainage, so that these tubes tend to remain filled 

(Kauahikaua and others, 1998, 2003). If flow through the 

tube is maintained, then continued crustal growth creates 

either elongate or hummocky tumuli, depending on the 

rate and steadiness of lava supply. Elongate tumuli mark 

the sites of steady and persistent lava transport; hummocky 

tumuli form as lava supply rates wane and (or) interior tube 

networks start to break down. Ubiquitous lava squeezeups 

and breakouts from elongate tumuli show that lava within the 

tubes is moderately overpressurized (Kauahikaua and others, 

2003; Hoblitt and others, 2012). In contrast, breakouts 

from hummocky tumuli are degassed, as illustrated by the 

formation of dense “blue glassy” lava flows (for example, 

Hon and others, 1994; Oze and Winter, 2005). 

Thermal Efficiency of Lava Tube Systems

Samples collected along lava tube systems confirm the 

inferences by early observers that lava tubes are efficient 

insulators and, thus, permit the transport of lava over long 

distances with minimal cooling (for example, Helz and others, 2003), as illustrated by the tube-fed structure of many of 

the longest terrestrial lava flows (for example, Cashman and 

others, 1998; Self and others, 1998), and, possibly, many of 

the longest planetary flows (Keszthelyi, 1995b).

Application of glass geothermometry suggests average 

cooling rates of 0.6 °C/km along recent Kīlauea lava tube 

systems (Helz and others, 2003), which translates to cooling 

rates of 6×10-4 to 18×10-4 °C/s for average flow rates of 1–3 

m/s (for example, Cashman and others, 1994; Mangan and 

others, 1995a; Thornber, 2001). Similar distance-referenced 

cooling rates have been obtained by analysis of quenched 

samples from older tube-fed pāhoehoe flows from both 

Kīlauea (Clague and others, 1995, 1999) and Mauna Loa 

(Riker and others, 2009), although these studies suggest that 

thermal efficiency may increase with increasing tube length 

(including the extreme case of flood basalts; for example, Ho 

and Cashman, 1997). Harris and Rowland (2009) classified 

lava tube systems as mature if they have thermal efficiencies 

of ~0.001 °C/s and immature if they have cooling rates of 

0.01 °C/s, similar to the maximum cooling rates estimated 

for proximal high-temperature channels (in other words, 

where cooling is predominantly radiative and not buffered by 

extensive crystallization).

In his analysis of all possible contributions to the heat 

budget of lava tubes, Keszthelyi (1995b) found latent heat of 

crystallization to be the single most important heat source; 

he also concluded that the effects of conduction (in filled 

tubes), convection (in open tubes), and rainfall are important, 

whereas radiative cooling from skylights was negligible. 

More recent FLIR measurements of skylight temperatures 

on the Kīlauea flow field, however, suggest that radiative 

heat loss and forced convection over skylights may also 

contribute to the overall heat budget of lava tubes (Witter 

and Harris, 2007). Because skylight formation appears to be 

controlled by both the underlying slope and the stability of 

lava supply from upslope, steady lava supply and gradual 

slopes should promote thermally efficient lava transport. 

Lava Flux Through Tubes

The flow rate of lava through tube systems can be determined using measurements of very low frequency (VLF) electromagnetic induction, from which the cross-sectional area of 

dense rock equivalent (DRE) lava can be calculated (Zablocki, 

1978; Jackson and others, 1985). Under most conditions, VLFbased lava effusion rates correlate well with rates determined 

by SO2 emissions (Sutton and others, 2003). An unusually well 

constrained example is provided by the VLF measurements of 

diminishing lava flux from Kupaianaha that were used to predict 

the shutdown date of the Kupaianaha vent (to within a few days; 

fig. 16; Kauahikaua and others, 1996). Here, the observed linear 

reduction in volumetric flow rate over time supports a model of 

shutdown driven by gradual pressure loss within the lava tube 

system, rather than inward solidification of a cylindrical conduit.

0

20

25

15

10

5

30

4/1/91 7/1/91 10/1/91 1/1/92 4/1/92

Linear magma pressure reduction

DRE volume , in tens of thousands of cubic meters per day

Conduit radius reduction

Figure 16. Plot of volumetric flow rate (measured by very-low-frequency 

electromagnetic induction) through a lava tube fed from the Kupaianaha 

vent of Kīlauea Volcano versus time. DRE, dense-rock equivalent. Steady 

decrease in flow rate was used to predict shutdown of Kupaianaha vent; 

comparison of conduit-freezing and pressure-reduction models indicates 

that pressure-reduction model provides a more reasonable fit to data 

(redrafted from Kauahikaua and others, 1996).
A Century of Studying Effusive Eruptions in Hawai‘i 377

An interesting observation is that lava levels within some 

open tubes (those with headroom) decrease steadily over time, 

despite maintaining a constant flow width and DRE flux, 

strongly suggesting erosion of the tube floor. On one occasion, 

changes in flow depth could be monitored directly, yielding 

erosion rates of 10 cm/day over several months (Kauahikaua 

and others, 1998). This rate can be modeled as thermal erosion 

by assuming steady, forced convective heat transfer by laminar 

channel flow at a large Peclet number (Kerr, 2001, 2009). The 

predictive capability of a thermal-erosion model does not, 

however, rule out the possibility that some component of the 

erosion is mechanical. 

Tube-Generated Flow Features

Lava tube systems evolve over time and undergo various 

types of morphologic alterations. Collapse of lava tube roofs 

can trigger breakouts that may be violent and characterized by 

“a sudden outburst of lava, throwing great blocks of the cooled 

lava, with the molten lava, into the air” (VL 520). Persistent 

weaknesses in tube roofs can produce surface structures, such 

as hornitos, rootless shields, and shatter rings (Kauahikaua and 

others, 2003). Hornitos form when open tubes fill temporarily 

with lava and contain sufficient gas to forcibly expel lava clots 

through a single opening in the tube roof; rootless shields form 

where perched lava ponds overflow repeatedly over a period 

of weeks and produce an accumulation of thin (and commonly 

shelly) pāhoehoe flows; and shatter rings are approximately circular areas of broken pāhoehoe that form over active lava tubes in 

places characterized by abrupt decreases in flow velocity within 

the tube (Orr, 2011). Because a decrease in flow velocity must 

be accompanied by an increase in lava thickness within the tube 

(by mass conservation), these places are susceptible to overpressurization and repeated rupture of the tube roof. In other parts of 

the world, similar features on older flows have been described as 

“collapsed tumuli” (Guest and others, 1984), “unusual craters” 

(Summerour, 1990), “craters with raised rims” (Greeley and 

Hyde, 1972), and “lava ponds” (Atkinson and Atkinson, 1995). 

The spatial arrangement of hornitos, lava shields, and shatter 

rings all provide important information on former lava transport paths and flow emplacement dynamics (both longevity and 

steadiness; Kauahikaua and others, 2003; Orr, 2011). 

Drained tubes are commonly decorated with lava 

stalactites—“some like grapes, some like walking sticks, 

and some like worms” (VL 345)—whose origin has been 

the subject of speculation since the earliest Western visitors to lava tubes (for example, Barton, 1884; Stearns and 

Clark, 1930; Wentworth and Macdonald, 1953). Jaggar (VL 

345) described the stalactites as “material of the gas-melted 

glaze” with an outer coating of “magnetic oxide of iron” and 

interior vesicles lined with crystals of feldspar and augite. 

Recent investigations of such “soda straw” stalactites confirm that their surfaces are enriched in oxide phases and that 

the interior walls are enriched in titanium (Baird and others, 

1985; Kauahikaua and others, 2003), both features consistent 

with remelting of tube walls. Long after flow emplacement, 

another generation of stalactites may form when rainwater 

gains access and hydrous sulfates precipitate as this water 

drips from the ceiling to the floor (Finch and Emerson, 1924; 

Thornber and others, 1999; Porter, 2000). 

Lava-Seawater Interactions

Interactions between flowing lava and the ocean are 

common in Hawai‘i. These interactions may be peaceful or 

explosive, depending on the flow style and specific conditions 

of ocean entry. Accounts of explosive lava-water interactions 

accompanying the ‘a‘ā-producing eruptions of Kīlauea in 1840 

and of Mauna Loa in 1868 and 1919 record both the violence 

of the explosions and the structure of the resulting fragmental 

cones (summarized by Moore and Ault, 1965). The violence 

of the lava-water interaction is believed to reflect the large 

surface area of rough ‘a‘ā flows. This hypothesis can be tested 

by examining ‘a‘ā flows that do not generate explosions at 

ocean-entry sites, such as three discrete sites created during 

the 1950 Mauna Loa eruption (fig. 3). Here, observers noted 

“a huge column of steam” but no violent explosions or ash 

generation. Finch and others (1950) speculated that the absence 

of explosions reflected both the high temperature and relatively 

smooth surface of the 1950 lava streams. 

Pāhoehoe flows are typically quiescent when entering the 

ocean, with lava quenched by seawater shattering only when it 

reaches the surf zone. Tube-fed pāhoehoe flows can, however, 

exhibit a range of lava-water interaction styles, including the 

unusually explosive lava-water interactions at Kīlauea between 

1992 and 1994 (Mattox and Mangan, 1997). One condition 

that promotes explosive activity at pāhoehoe ocean entries is 

the formation of lava deltas with unstable edges that collapse 

repeatedly, thereby exposing the lava tube to ocean water. 

Under these conditions, four types of explosions may occur 

(Mattox and Mangan, 1997): tephra jets, lithic blasts, bubble 

bursts, and littoral lava fountains (fig. 17). Tephra jets, the 

most common explosions, form when lava from the severed 

tube is exposed to the surf zone. Prolonged tephra-jet activity 

at a single site can produce a semicircular agglutinated littoral 

cone. Lithic blasts form during delta collapse, when seawater 

contacts newly exposed incandescent rock. These two types 

of lava-water interaction can be explained by open mixing 

processes at water/melt ratios of ~0.15 (fig. 17A). Bubble 

bursts, possibly the most spectacular (though relatively mild) 

form of lava-seawater interaction, form when seawater gains 

rapid entry into a confined lava tube. Bubble bursts produce 

fluid bombs, limu-o-Pele (thin glass sheets) and Pele’s hair, all 

of which can accumulate to form circular agglutinated cones. 

Littoral fountains are rare but may accompany bubble bursts; 

they form circular spatter cones inland from the shoreline. 

Both require confined mixing of lava and water within the tube 

(fig.17B), with the intensity of the explosion controlled by 

the rate of water entry, which determines the water/melt ratio. 

Confined mixing of lava and seawater may explain the large, 
378 Characteristics of Hawaiian Volcanoes

circular littoral cone associated with a prehistoric pāhoehoe 

flow from Mauna Loa (Jurado-Chichay and others, 1996).

Contributions of Hawaiian Volcanology 

to Planetary Volcanism

The application of observations of Hawaiian lava flows to 

the interpretation of planetary lava flow features has a long history. 

Early geologists (for example, Scrope, 1872) recognized the 

volcanic origin of lunar surfaces, and Hitchcock (1911) reported 

that in 1905, W.H. Pickering of Harvard University visited Hawai‘i 

to make comparisons with lunar features (particularly craters). 

Since that time, Hawai‘i has served as both a benchmark and a 

testing ground for developing methods of analyzing planetary 

surfaces (for example, Greeley, 1974; Carr and Greeley, 1980; 

Mouginis-Mark and others, 2011; Rowland and others, 2011). 

Analysis of surface features to infer the origin of 

landforms (volcanic geomorphology) is the core of planetary 

volcanism; the calibrations required for feature interpretation 

have come from studies of terrestrial basaltic lava flows. 

Hawaiian lava flows, in particular, have provided an 

important baseline for mapping the surfaces of the Moon, 

Mercury, Venus, and Mars; and recent activity in Hawai‘i 

has provided a context for the interpretation of observed 

eruptions on Jupiter’s moon Io. Hawai‘i has served as a 

critical testing ground for applying different types of remote 

sensing to mapping volcanic regions. Hawaiian lava flows 

are also used to benchmark models of flow emplacement, 

Figure 17. Primary types of lava-water interaction at ocean entries of tube-fed lava flows from Kīlauea Volcano. Lava-water interactions are divided 

into two primary types. A, Open mixing occurs when collapse of lava delta exposes lava tube to seawater. Waves breaking on exposed lava can produce 

either tephra jets or, at time of delta collapse, directed lithic blasts. B, Confined mixing of seawater and lava occurs when water gains entrance to tube 

through cracks within lava delta. Under these circumstances, heated steam confined within lava produces spectacular large bubble bursts or, more rarely, 

energetic littoral lava fountains (redrafted from Mattox and Mangan, 1997).

particularly in order to infer conditions of lava flow 

emplacement from measured lava flow geomorphologies. 

Here, we do not attempt a comprehensive review of the ways 

in which Hawaiian studies have contributed to the study 

of extraterrestrial volcanism, but instead provide a brief 

overview of examples where the lava flow features described 

above have been used to deduce volcanic processes on other 

planets. 

Ground Truthing Remote-Sensing Techniques

Exploration of planetary surfaces has been accomplished 

almost entirely by remote sensing. For this reason, numerous 

experiments have been conducted in Hawai‘i to assess the 

sensitivity of remote-sensing techniques for identifying different 

types of volcanic features. Flow-surface characteristics can 

be analyzed by using either multispectral thermal (Kahle and 

others, 1988) or radar (Gaddis and others, 1989, 1990) data. 

Thermal data can be used to map individual lava flows or 

flow lobes but does not work well for larger scale mapping. 

More useful is radar, which can penetrate clouds (essential for 

mapping the surface of Venus) and uses an active sensor (which 

allows data acquisition in permanently shaded regions; see 

overview by Mouginis-Mark and others, 2011). Radar images 

highlight topographic features, such as faults and caldera walls, 

and can be used to distinguish smooth from rough flow surfaces 

(Gaddis and others, 1989). In this context, “smooth” and 

“rough” are defined with reference to the scale of vertical relief 

(measured as the root mean square of vertical elevation changes) 

when calibrated for the incidence angle of the radar. New 

Waves breaking

on lava

Lithic blasts

Tephra jets

Water entering

through cracks

Littoral

lava

fountains

Bubble

bursts

A B
A Century of Studying Effusive Eruptions in Hawai‘i 379

Synthetic Aperature Radar techniques and both airborne laser 

scanning and terrestrial laser scanning have also been tested at 

Kīlauea to improve surface-roughness analysis for application 

to mapping both Mars and Venus (for example, Campbell and 

Shepard, 1996; Carter and others, 2006; Morris and others, 

2008). Development of these remote-sensing techniques has fed 

back to monitoring of Hawaiian volcanoes, particularly with 

regard to improved techniques of thermal imaging and analysis 

(for example, Wright and others, 2010, 2011).

Mapping Planetary Surfaces 

Identifying lava flow features on the Moon and inner 

planets has been a primary goal in mapping planetary 

surfaces. For example, lunar sinuous rilles, first identified 

by Christian Huygens in 1684, were recognized as bearing 

a striking resemblance to collapsed lava tubes because of 

observations during the 1969–74 Mauna Ulu eruption (for 

example, Greeley 1971; Cruikshank and Wood, 1972; Carr, 

1974). Also important has been development of techniques 

to distinguish ‘a‘ā from pāhoehoe flow surfaces, including 

not only the surface-roughness techniques described above 

(Gaddis and others, 1989), but also fractal analysis of lava 

flow margins (Bruno and others, 1994). 

As imaging techniques have improved, so has the 

quantification of planetary lava-flow features. For example, 

maps showing the distribution and morphology of lava flows 

in the Elysium Planitia region of Mars provide information not 

only on flow aspect ratios, but also on their spatial distribution 

(Mouginis-Mark and Yoshioka, 1998). One important 

observation from this work is that all the flows erupted <200 

km from the summit of Elysium Mons are short (on Martian 

scales; <70 km), whereas the 11 longest flows have vents 

>294 km from summit. This observation is similar to eruptive 

patterns that contribute to the classic so-called “inverted soup 

bowl” form of basaltic volcanoes in the Galápagos Islands, and 

may reflect the unusually steep upper flanks (>7°; Wilson and 

Mouginis-Mark, 2001) of Elysium Mons: flow of lava down 

steep slopes enhances cooling and crystallization and, thus, 

limits flow length (for example, fig. 12). 

Increased resolution of Mars surface images provided by 

the Mars Reconnaissance Orbiter’s High Resolution Imaging 

Science Experiment permits detailed mapping of meter-scale 

volcanic features (Keszthelyi and others, 2008). Also possible 

is analysis of the emplacement conditions of individual lava 

flows. One example uses tumulus spacing on hummocky 

pāhoehoe surfaces to infer volumetric flux (Glaze and others, 

2005; Bruno and others, 2006). Mars Orbiter Laser Altimeter 

(MOLA) digital elevation data allow detailed measurements 

of larger flow features, such as channel and flow dimensions, 

and changes in flow dimensions as a function of distance from 

source vents (fig. 18). In an example provided by Glaze and 

Baloga (2006), both flow width and flow thickness increase as 

a function of distance from the source vent, consistent with a 

distributary model of lava transport. 

Dynamics of Planetary Volcanism

 Some models of terrestrial lava flow behavior have been 

constructed specifically for application to planetary processes. 

For example, Hulme’s (1974) widely used model of levee 

formation in Bingham fluids was constructed specifically with 

the intent of obtaining both rheologic and dynamic information 

on planetary lava flows (Hulme, 1982). This Bingham model 

has been justified because Martian lava flows do not widen at 

the rate expected for Newtonian fluids (for example, Wilson and 

Head, 1994). An alternative explanation for the absence of flow 

widening is levee formation because of marginal flow cooling 

(Kerr and others, 2006). Consideration of cooling-limited flow 

widths is important because the cooling histories of planetary 

flows are not expected to differ significantly from those of 

terrestrial lava flows (for example, Wilson and Head, 1983, 

1994; Head and Wilson, 1986, 1992). More important is the 

effect of variations in gravity on magma ascent and eruption in 

planetary and terrestrial environments. Lower gravity will affect 

both magma ascent and flow across the surface and must be 

accounted for if observed relations between lava flow geometry 

and effusion rate are used to estimate the eruptive conditions of 

planetary flows (for example, Malin, 1980; Crisp and Baloga, 

1990; Head and Wilson, 1992). Curiously, Wilson and Head 

(1994) predicted that cooling-limited flows will be longer by 

a factor of 6 on Mars than on Earth, because the lower gravity 

means that flows will be thicker for a given yield strength and 

slope. Alternatively, Rowland and others (2004) suggested that 

channel-fed flows should be shorter on Mars, because lower 

flow rates will allow more extensive cooling and crystallization 

per unit distance. In any case, Martian flows, in particular, are 

generally much larger than current Hawaiian lava flows, more 

on the scale of flood basalts. Interestingly, however, many 

flow features seem to scale similarly to those on Earth. One 

example is the partitioning of lava between stagnant levees and 

active channels, as determined from analysis of MOLA data 

(fig. 18) and data from the 1984 Mauna Loa eruption (Lipman 

and Banks, 1987; fig. 19). Although volumes differ, both flows 

show increasing lava storage in stagnant levees and overflows 

with increasing distance from the vent, consistent with observed 

decreases in volumetric flux along channel networks (fig. 11).

Recent research on Mars outflow channels has also 

revived discussion of the extent to which planetary surfaces 

can be sculpted by volcanic (rather than fluvial) processes 

(for example, Hulme, 1982; Leverington, 2004; Jaeger and 

others, 2010). Pertinent to this discussion is evidence in 

Hawai‘i for both thermal (Kauahikaua and others, 1998) and 

mechanical (Kauahikaua and others, 2002) erosion in lava 

tubes and channels, consistent with models for the formation 

of sinuous rilles and analogous features (for example, Fagents 

and Greeley, 2001). This work has also stimulated mapping of 

constructional features that may reflect lava-water interaction 

(for example, Hamilton and others, 2010).

Finally, discovery of active volcanism on Jupiter’s moon 

Io has prompted direct comparison to Hawaiian eruption 
380 Characteristics of Hawaiian Volcanoes

2,550

2,530

2,510

2,490

2,470

0 10 20 30 40

Distance, in kilometers

Elevation, in meters

Profile 15

w

h

Profile 15

Flow width (w),

in kilometers

Flow thickness (h),

in meters

70

50

30

10

0 50 100 150 200

Distance, in kilometers

Flow dimension

C

A B

0

0

5 10 KILOMETERS

2.5 5 MILES

Figure 18. Martian lava flow. A, Image of segment of a long, channelized lava flow on northern flank of Pavonis Mons (from Baloga and others, 2003). 

Dashed lines indicate flow margin, black arrows point to central channel, white arrows indicate channel cross section in part B. THEMIS infrared image 

I01739006 from Glaze and Baloga (2006). B, Cross section of lava flow, showing marginal levees and central channel (vertical exaggeration ~400x). C, 

Summary of data along flow showing variations in both flow thickness and flow width as a function of distance from vent. From Baloga and others (2003).

Figure 19. Measurements of lava storage along channels. A, Illustration of assumptions used to infer volume stored within marginal levees. B, 

Cumulative stored volume as a function of relative distance along a lava flow. Xs show lava storage along Martian channels (modified from Glaze and 

Baloga, 2006). Squares show lava storage along the 1984 Mauna Loa channel (calculated from fig. 11B). Increases in stored lava volume with distance 

illustrate extent to which lava transferred through proximal channels is lost to either marginal levee construction or channel overflows. 

w(x)

d(x)

h(x)

Lw

Active

channel

11

10

9

8

7

6

5

0 0.5 1

Normalized distance

Cumulative volume,

in cubic meters

X

X X X X X X X X X X X X X X X

X X

X

X X

X X X X X

X X

X

X

Pavonis

Ascraeus

Mauna Loa 1984

A B

EXPLANATION
A Century of Studying Effusive Eruptions in Hawai‘i 381

mechanisms. Repeat observations of volcanic activity on Io 

allow discrimination of two eruption modes: Promethean and 

Pillanian (Keszthelyi and others, 2001). Promethean volcanic 

activity is long lived but pulsatory and produces compound 

lava flow fields. This eruptive style, which has been equated 

directly with that of Pu‘u ‘Ō‘ō during the 1980s, is interpreted 

to result from shallow magma storage (Davies and others, 

2006). Pillanian volcanic activity, in contrast, is characterized 

by short, energetic eruptions that produce large pyroclastic 

deposits and fissure-fed lava flows; this activity is interpreted to 

reflect rapid transport of magma from deeper storage regions. 

An eruption of this type in 1997 involved a 40-km-long fissure 

and emplacement of two 50-km-long flow lobes (Williams 

and others, 2001). Estimated effusion rates of 103 to 104 m3/s 

for this eruption exceed those of historical eruptions from 

Kīlauea but overlap those of the 1950 Mauna Loa eruption. This 

correspondence, together with the general tendency of eruptions 

from Mauna Loa to initiate at higher effusion rates than those of 

Kīlauea (at least historically), raises interesting questions about 

the general relations between magma-storage depths, volumes, 

eruption rates, and styles of basaltic volcanism.

Summary

Succinctly summarizing a century of Hawaiian lava flow 

research is difficult; however, HVO has clearly made a critical 

contribution to volcanology through careful documentation of 

the numerous effusive eruptions that have occurred at Mauna 

Loa and Kīlauea volcanoes since 1900. The value of complete 

records of flow field emplacement cannot be overstated, because 

these data provide a baseline for testing hypotheses and models. 

Eruptions of Mauna Loa in 1984 and of Pu‘u ‘Ō‘ō from 1983 

to the present are particularly well characterized, thanks to 

helicopter support for channel measurements in 1984 and 

persistent monitoring, coupled with expanding technological 

advances during the past three decades. Detailed maps of these 

eruptions illustrate the complexity of lava flow fields and provide 

continued opportunity for analysis and modeling.

Close observations of flowing lava have led naturally 

to detailed studies of lava properties, particularly the relation 

between the thermal and rheologic evolution of lava flows 

and the consequent changes in flow-surface morphology. 

Early pioneering studies using pyrometers, thermocouples, 

and simple rheologic models have been supplemented by 

new field measurement tools, laboratory calibrations, analog 

experiments, and theoretical analyses. Together, these data 

highlight the complex physics underlying the basaltic flowsurface morphologies that have fascinated scientists for so long. 

Observed links between cooling conditions, crystallization, and 

flow dynamics also underline the importance of linking physical 

and chemical driving forces in many aspects of volcanology.

Finally, careful observations during lava flow episodes have 

catalyzed important studies of flow emplacement dynamics. As 

a result, the basic characteristics of channelized ‘a‘ā and tubefed pāhoehoe flows are now well understood. These flow types 

are distinguished primarily by eruption rate, which controls 

the relative balance between flow advection and formation of 

surface crusts. 

Channelized flows form when volumetric eruption rates are 

high. Initial flow advance rates are controlled by effusion rate, 

although flow advance rates generally decrease with increasing 

flow distance from eruptive vents because of loss of lava to 

the channel margins, as well as cooling, crystallization, and 

consequent increases in lava viscosity. Simple channelized flows 

reach lengths that are proportional to the volumetric flux. Longlived flows are not simple, however, because lava channels evolve 

in space and time as levees form, flows split around topographic 

obstacles, and lava spillovers create new channel branches. These 

modifications to channel networks affect both the flow advance 

rate and the ultimate length achieved by individual flow lobes. 

Although both probabilistic and single-channel models have 

been constructed for lava flow prediction, full characterization of 

channelized lava flows has not yet been achieved. 

Tube-fed pāhoehoe flows form when initial volumetric flow 

rates are sufficiently low that lava crusts form more rapidly than 

they are disrupted by advection. Because lava is an excellent 

insulator, tubes can transport lava over long distances with little 

cooling. Tubes on moderate slopes can erode their base; as a 

result, lava tubes are open (have headroom). Tubes on low slopes 

remain filled and, therefore, cool by conduction through both the 

upper and basal crust. Breakouts from filled tubes are common 

and modulate the spatial coverage of pāhoehoe flow fields. The 

tendency of pāhoehoe flows to inflate from very thin initial to 

thick final flows requires both very high-resolution digital terrain 

models and frequent remapping to anticipate the effect of newly 

inflated topographic barriers on paths of subsequent lava flows. 

A century’s perspective is an interesting lens from which to 

view any field in science. In writing this review, we have been 

impressed by the prescience of early scientists in identifying 

topics that have proved to be of fundamental importance not only 

to volcanology, but also to such fields as materials science, fluid 

mechanics, and planetary science. We have also come to appreciate 

the wisdom of developing, promoting, and maintaining a natural 

laboratory for studies of basaltic volcanism and for routine, 

continuous monitoring of eruptive events. Although the sheer 

number of scientists who have visited, worked with, and learned 

from HVO scientists is difficult to quantify, it is through these 

contacts that HVO has had a global impact on lava flow studies.

Acknowledgments

We sincerely thank Laszlo Keszthelyi and Scott Rowland 

for their prompt and thorough reviews on the manuscript, 

Hannah Dietterich for her help with the figures, and Mike 

Poland for his patience and editorial assistance. Additionally, 

we thank all of our HVO coworkers over the years who have 

provided endless insight and training. This work was supported 

by National Science Foundation grants EAR 0738894 and EAR 

1250554, and the AXA Research Fund (both to the first author). 
382 Characteristics of Hawaiian Volcanoes

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Side view into a skylight on a lava tube issuing from the Pu‘u ‘Ō‘ō eruptive vent. Note lava stalactites hanging from the tube’s ceiling. The 

ends of the longer stalactites were flexible and wafted in the breeze that was created by heat billowing from the skylight. USGS photograph 

by M.R. Patrick, February 19, 2008.
Aerial photograph of lava flow from Mauna Loa’s 1926 Southwest Rift Zone eruption, which destroyed the town of Ho‘ōpūloa. Photograph 

was taken April 18, 1926, by the U.S. Army Air Corps, 11th Photo Section, courtesy of the University of Hawai‘i at Hilo. The flow reached 

the ocean that same day, inundating the town. 
Abstract

Significant progress has been made over the past 

century in understanding, characterizing, and 

communicating the societal risks posed by volcanic, 

earthquake, and tsunami hazards in Hawai‘i. The work of 

the Hawaiian Volcano Observatory (HVO), with a centurylong commitment to serving the public with credible 

hazards information, contributed substantially to this 

global progress. Thomas A. Jaggar, Jr., HVO’s founder, 

advocated that a scientific approach to understanding these 

hazards would result in strategies to mitigate their 

damaging effects. The resultant hazard-reduction methods 

range from prediction of eruptions and tsunamis, thereby 

providing early warnings for timely evacuation (if 

needed), to diversion of lava flows away from high-value 

infrastructure, such as hospitals. In addition to long-term 

volcano monitoring and multifaceted studies to better 

understand eruptive and seismic phenomena, HVO has 

continually and effectively communicated—through its 

publications, Web site, and public education/outreach 

programs—hazards information to emergencymanagement authorities, news media, and the public. 

Although HVO has been an important global player 

in advancing natural hazards studies during the past 

100years, it faces major challenges in the future, among 

which the following command special attention: (1) the 

preparation of an updated volcano hazards assessment 

and map for the Island of Hawai‘i, taking into account not 

only high-probability lava flow hazards, but also hazards 

posed by low-probability, high-risk events (for instance, 

pyroclastic flows, regional ashfalls, volcano flank collapse 

and associated megatsunamis), and (2) the continuation of 

timely and effective communications of hazards information to all stakeholders and the general public, using all 

available means (conventional print media, enhanced Web 

presence, public-education/outreach programs, and socialmedia approaches).

Introduction

Basic studies of volcanoes and their past and present 

behavior provide the solid scientific foundations that underlie 

the mitigation strategies to reduce the risk from volcano 

hazards. Although scientists universally accept this paradigm 

today, that was not the prevailing point of view in the early 

20th century. Thomas A. Jaggar, Jr., founder of the Hawaiian 

Volcano Observatory (HVO), was an early and staunch devotee. 

As emphasized earlier in this volume (Tilling and others, this 

volume, chap. 1), Jaggar was profoundly influenced by several 

natural disasters early in the 20th century (for example, the 

Montagne Pelée eruption in 1902 with 29,000 deaths and 

the Messina earthquake and tsunami in 1908 with 60,000–

120,000 deaths; Tanguy and others, 1998; Risk Management 

Solutions, 2008). He became convinced that the only effective 

way to minimize the death and devastation from eruptions, 

earthquakes, and tsunamis was to study these potentially 

destructive phenomena continuously by means of permanent 

Earth observatories, documenting their processes and impacts 

before, during, and after each event. With the establishment of 

HVO in 1912 and throughout his entire career, Jaggar continued 

to advocate that “The main object of all the work should be 

humanitarian—earthquake prediction and methods of protecting 

life and property on the basis of sound scientific achievement” 

(Jaggar, 1909). The Hawaiian Volcano Research Association, 

a private business organization formed to support Jaggar and 

HVO, adopted for its motto: “Ne plus haustae aut obrutae 

urbes” (“No more shall the cities be destroyed”). 

Jaggar’s commitment to, and approach in, using scientific 

data to reduce the risk from natural hazards have remained a 

hallmark of HVO’s studies since its inception. The Island of 

Hawai‘i2, with its five volcanoes (Kohala, Hualālai, Mauna 

1

U.S. Geological Survey.

2

The differences in our usage of the words “Hawaii” and “Hawaiʻi” in this 

chapter are intentional and specific: “Hawaiʻi” is used to denote the eight 

main islands, while “State of Hawaii,” “Hawaii State,” or “Hawaii” refers to 

anything related to the State government (which includes the northwestern 

Hawaiian Islands, extending northwest to Kure Atoll, as well as Hawaiʻi). The 

use of “Island of Hawaiʻi” applies only to the southeasternmost island in the 

Hawaiian archipelago.

Characteristics of Hawaiian Volcanoes

Editors: Michael P. Poland, Taeko Jane Takahashi, and Claire M. Landowski

U.S. Geological Survey Professional Paper 1801, 2014

Chapter 10

Natural Hazards and Risk Reduction in Hawai‘i

By James P. Kauahikaua1 and Robert I. Tilling1
398 Characteristics of Hawaiian Volcanoes

Kea, Mauna Loa, and Kīlauea, plus the submarine volcano 

Lo‘īhi off the south coast) has shown itself to be a dynamic—

and at times hazardous—landscape over the past century, and 

this is not expected to change in the near future. Between 1912 

and 2012, HVO and affiliated scientists documented 12 Mauna 

Loa eruptions, almost 50 Kīlauea eruptions, one Hualālai intrusion, two Lo‘īhi eruptions, 8 earthquakes of magnitude 6.0 or 

greater, and several tsunamis. Detailed documentation of these 

events has led to many important discoveries. 

Jaggar intrinsically combined hazards (potentially 

destructive physical processes, such as lava flow, earthquake, 

tsunami) and risk (estimation of the potential loss, such as 

life, property, infrastructure, and productive capacity) in his 

early predictions and assessments. In subsequent decades, 

when HVO came under U.S. Geological Survey (USGS) 

administration, the emphasis was primarily on hazards and 

only recently began to again include risk studies (for example, 

Trusdell, 1995; Wood, 2011). 

This chapter presents a brief history of efforts to understand 

and characterize volcano (fig. 1A) and earthquake (fig. 1B) 

hazards in Hawaiʻi and efforts to minimize their adverse effects. 

Although not a direct part of HVO’s current mission, tsunami 

hazards (fig. 1C) are also mentioned because HVO scientists 

have provided pivotal results that allowed forecasts of tsunami 

generated by distant earthquakes and documentation of locally 

generated tsunami. The authors feel that, for this chapter, it is 

more important to describe general progress in each hazard 

in which HVO scientists have played key roles, rather than 

confining the chapter’s scope to what HVO has done specifically 

toward hazard and risk mitigation in its first century. This 

general approach, we believe, may be more useful to hazard 

specialists. The geologic processes and concepts needed for 

understanding the root causes of hazards in Hawaiʻi are treated 

by Clague and Sherrod (this volume, chap. 3). 

Orchid

Prince

Paradise

Ali‘i

Hoku

Plumeria

Paradise

A

B

C

Figure 1. Photographs illustrating volcano-, earthquake-, and 

tsunami-related hazards in Hawai‘i. A, Thermal image overlaid upon 

a visual image showing active lava flows advancing through the 

now-abandoned Royal Gardens subdivision southeast of Pu‘u ‘Ō‘ō on 

February 24, 2012. The subdivision streets are named, and the last 

remaining house, which was destroyed about a week after this photo 

was taken, is circled in red (USGS image compiled by Matt Patrick). 

The Royal Gardens subdivision has been inundated by lava flows 

several times during the ongoing East Rift Zone eruption of Kīlauea 

Volcano. Less than 5 percent of the original property currently remains 

uncovered by lava (Tim Orr, written commun., 2012). B, View of the 

Kalahikiola Congregational Church showing the collapsed portions 

of the wall under the eave, the framing, and the wall above the doors 

resulting from the magnitude 6.7 Kīholo Bay and 6.0 Mahūkona 

earthquakes on October 15, 2006. Church is located at the northern tip 

of the Island of Hawai‘i (USGS photograph by T.J. Takahashi, October 

20, 2006; from Takahashi and others, 2011). C, Photograph of piled-up 

vehicles at Napo‘opo‘o point near Kealakekua Bay, Island of Hawai‘i, 

transported by runup from the March 11, 2011, Tohoku-oki (Japan) 

tsunami (photograph from Trusdell and others, 2012).

A primary aim of this chapter is to provide nontechnical 

background information for use by individuals and agencies 

responsible for developing and implementing programs to 

mitigate risks posed by natural hazards. Nonetheless, we 

believe that it should also be useful to geoscientists with an 

interest in the history and evolution of hazards studies in the 

Hawaiian Islands.
Natural Hazards and Risk Reduction in Hawai‘i 399

Volcano Hazards: Effusive Eruptions

Hazardous processes directly or indirectly associated 

with volcanic activity have posed the greatest threat to the 

residents of, and visitors to, the Island of Hawai‘i during 

the past two centuries. Although explosive eruptions have 

occurred at Kīlauea during the past 250 years and at other 

Hawaiian volcanoes in the recent geologic past, the most 

recent two centuries have been overwhelmingly dominated 

by nonexplosive (effusive) activity. Accordingly, lava flow 

hazards have been the most common and have caused the most 

damage and disruption to daily life in Hawaiʻi, and, thus, will 

receive the most attention in the discussions to follow.

How lava flows form and move must be fully understood 

before the hazards they pose can be characterized. At 

Hawaiian volcanoes, eruptive vents are primarily located 

within the summit areas and along the curvilinear rift zones 

that extend radially away from the summits. Lava flow 

mechanics are relatively well understood because of the 

detailed studies done over the past century (see Cashman and 

Mangan, this volume, chap. 9), but more needs to be done. 

For example, internal lava flow structures, such as lava tubes 

or channels, are critically important to the distribution of lava 

during an eruption, yet only bare essentials are known about 

what initiates formation and failure of these structures.

A comprehensive assessment of volcanic hazards must 

be based, additionally, on geologic mapping of historical 

and prehistoric eruptive products. Lava flow hazards were 

first depicted as a map of zones, with each zone qualitatively 

defined by its proximity to vents and the rate of coverage 

by past lava flows. More quantitative lava flow hazards 

can be estimated as probabilities of coverage based solely 

on the recurrence interval of past lava flows within a given 

area. Finally, some aspects of lava flow hazards can also be 

investigated using computer software that simulates flowing 

lava (see, for example, Rowland and others, 2005). However, 

present computer and numerical models simplify lava flow 

dynamics and cannot yet fully simulate lava flow behavior, 

such as tube and channel development and transitions from 

ʻaʻā to pāhoehoe modes (Harris, 2013), and they must be used 

with caution.

Efforts to mitigate risk can only be implemented after 

the hazards are characterized. For lava flows, some attempts 

have been made on the Island of Hawaiʻi and in Iceland and 

Italy to mitigate the risk of their potential hazards by diverting 

or impeding lava flows using water-cooling, explosives, 

or construction of barriers (see, for instance, Williams and 

Moore, 1973; Lockwood and Torgerson, 1980; Barberi and 

others, 1992; Williams, 1997; Peterson and Tilling, 2000).

Understanding Lava Flow Emplacement 

“Lava” is an all-inclusive term for magma that breaches 

the Earth’s surface to erupt effusively or explosively. Once 

still-molten lava is confined to channels or tubes, its flow 

dynamics and emplacement are controlled primarily by 

viscosity, eruption rate at the vent(s), eruption duration, and 

topographic attributes of the terrain, such as steepness. Lava 

can be erupted as smooth pāhoehoe or rough ʻaʻā flows, with a 

transition between the two that is defined primarily by internal 

shear rates (Peterson and Tilling, 1980; Soule and Cashman, 

2005). ʻAʻā lava flows generally advance faster and, therefore, 

pose the greater hazard. The progression of lava flow types 

during an eruption usually starts with ʻaʻā flows that become 

channelized during the initially high effusion rates and may 

change to pāhoehoe flows that could eventually cover the 

initial ʻaʻā flows. Pāhoehoe flows often change to ʻaʻā as they 

flow over steep slopes, with gravity providing the increase in 

internal shear rates. It stands to reason that when an ʻaʻā flow 

reaches flatter terrain, it may change back to pāhoehoe, as has 

been observed in the field (Hon and others, 2003). On gentle 

slopes, pāhoehoe flows can be emplaced endogenously (from 

within; see Hon and others, 1994; Walker, 1991), forming 

extensive inflated flow fields that include tumuli (small 

mounds) and lava tubes (Kauahikaua and others, 2003).

Geologic mapping has been able to define the source 

regions from which past lava flows were erupted, and current 

digital elevation models (DEMs) can be used to forecast 

the paths that lava flows will follow in advancing from the 

sources. Less well known are the factors that control the rate 

of advance of lava flows and the ultimate width of the flow 

field produced by prolonged eruptions. Our understanding of 

lava flow mechanics has advanced significantly over the past 

century (Cashman and Mangan, this volume, chap. 9), and this 

improved understanding has allowed us to isolate the critical 

parameters that we do not yet measure routinely but that are 

necessary to accurately estimate lava flow advance rates. 

Of the parameters influencing lava flow emplacement, 

probably the most important is eruption rate. Several methods 

for computing eruption rate have been studied—differencing 

DEMs (for instance, Rowland and others, 1999), geophysical 

measurements of lava flux through tubes (for instance, 

Kauahikaua and others, 1998a), proxy measurements of 

sulfur dioxide emission rates (Sutton and others, 2003), and 

satellite measurements of thermal radiance from advancing 

active flows (for instance, Harris and others, 1998)—but, to 

date, none of these has provided a reliable, routine monitor of 

eruption rate for Hawaiian volcanoes.

The true measure of how well we understand lava flows 

will be our ability to model their advance. Several different 

approaches have been used over the past many years, and they 

have all necessarily involved simplifications (Harris, 2013). 

Each works well for a specific set of circumstances, but there 

is still a need to develop a more robust, generalized approach 

that incorporates the full physics of the emplacement process 

and rheology of lava. It should also be emphasized that 

numerical or computer simulations are critically sensitive to 

uncertainties in parameters, such as eruption rate, the changing 

rheology of molten lava as it cools, degasses, and advances 

away from the vent, and topography; propagation of error 

through the resulting modeled flow may have a significant 
400 Characteristics of Hawaiian Volcanoes

effect on the accuracy and reliability of any prediction based 

on the model. While these efforts have greatly improved over 

the past several years, they may not yet be ready for diagnostic 

use during emergency situations.

Characterizing Lava Flow Hazards

For characterization of lava flow hazards, arguably the 

most important consideration for any volcano is where, and 

how often, lava inundation has occurred in the past. The 

basic data for hazards determination are a detailed geologic 

map and comprehensive age dating of the surface lavas. In 

general, eruptions at shield volcanoes that are in their most 

vigorous stage of development, such as Kīlauea and Mauna 

Loa, originate within the summit and rift zone areas. On 

the other hand, volcanoes that may be entering into, or are 

already within, the less vigorous postshield or the more mature 

rejuvenated stage of development have had eruptive vents that 

are not so confined spatially. In either case, detailed geologic 

mapping is key in identifying the eruptive vents and pathways 

of past lava flows and areas of possible future lava inundation.

Mullineaux and others (1987) made the first 

comprehensive assessments of the volcanic hazards affecting 

the Hawaiian Islands. Because of the geologic mapping and 

dating studies made since that work, we now have improved 

assessments for the short-term, most frequently occurring lava 

flow hazards associated with the dominantly effusive eruptions 

of Kīlauea (mostly) and Mauna Loa during the past two-plus 

centuries. With additional geologic and radiometric data, 

we can expect to obtain even more refined assessments and 

precise zonation maps for earthquake, lava flow, and volcanic 

gas hazards.

Lava Flow Hazard Zone Maps

The first lava flow hazards map for Hawaiʻi was recently 

rediscovered in HVO files and may date to the 1940s or 1950s. 

The unknown author mapped linear zones that included the 

summits and rift zone vents as having the highest lava flow 

hazard; other zones were rated by their proximity to the 

highest hazard zone (vents) and by the recurrence interval 

of historical lava flows within the zone (estimated without 

any radiometric dates for the lava flows prior to 1800). The 

recurrence rate estimates ranged from 0 (Mauna Kea and 

Kohala) to 1,000 (Kīlauea summit) flows “per 10,000 years 

per square mile” (fig. 2). With the benefit of hindsight, it is 

apparent that the estimates of volcano productivity portrayed 

on this map were low compared to our current estimates based 

on more abundant and much improved knowledge.

Beginning in the 1970s, following additional mapping 

and dating studies, a number of USGS assessments and lava 

flow hazards maps were prepared, some supported in part 

by the Department of Housing and Urban Development 

(Crandell, 1975, 1983; Mullineaux and Peterson, 1974). A 

few of these studies also included other volcanic hazards in 

addition to lava flows (for instance, tephra fall, pyroclastic 

surges, subsidence, ground fractures); Mullineaux and others 

(1987) summarized and updated such work through the 

mid-1980s. 

The first published lava flow hazards map for the Island 

of Hawaiʻi was that of Mullineaux and Peterson (1974), which 

portrayed hazards zones rated according to severity from A 

through F, with zone F the most hazardous. This map was later 

modified by Mullineaux and others (1987), using additional 

data collected since the 1970s, and it depicted lava flow hazard 

in terms of nine zones. These lava flow hazard zones were 

qualitative and based on volcano structure and coverage rates. 

Hazard Zone 1, the most hazardous, was linear, because it 

included vents in the summits and linear rift zones of the most 

active volcanoes—Kīlauea and Mauna Loa. Hazard Zones 2 

and 3 reflected areas downslope of vent areas on those same 

volcanoes. Hazard Zone 4 included all of Hualālai volcano, the 

third most active volcano on the island. Hazard Zones 5 and 

6 were areas on Mauna Loa and Kīlauea that were protected 

from lava inundation by topography (for instance, caldera 

walls or locally high relief). Hazard Zone 7 included the most 

recently active vents on Mauna Kea (eruptions between 5,400 

and 4,600 years before present; Wolfe and others, 1997), Zone 

8 included the rest of Mauna Kea, and Zone 9 included Kohala 

Volcano, which has had no active volcanism in the past 10,000 

years. 

Additional dating and review of the 1974 and 1987 maps 

for Hawaiʻi resulted in an updated, slightly modified version 

of the map (fig. 3), published by Wright and others (1992), 

that is still in use today. Although the 1992 map (fig. 3), which 

included coauthors from the County of Hawaiʻi, was intended 

to guide development planning within the county, it has yet 

to be used for that purpose. It has, however, been used by 

property insurance and mortgage companies to set increased 

rates in hazard zones 1 and 2. 

Qualitative hazard zonation, the basis for the existing 

lava flow hazard maps, requires prioritization of frequency 

and magnitude of the hazard. For example, the map of Wright 

and others (1992) emphasizes the rate of lava flow coverage, 

which probably best reflects the combination of these two 

factors (frequency and magnitude) and is appropriate because 

the range of intensities for effusive eruptions is limited. Smallvolume effusive eruptions, which are most common, have 

a more restricted spatial impact than the high-volume, but 

more infrequent, effusive activity, such as during the ‘Ailā‘au 

eruption of Kīlauea about 600–550 years ago (Clague and 

others, 1999; Swanson and others, 2012a). Areas near vents 

are inundated by more lava flows than areas that are more 

distant, reachable only by eruptions of the longest duration 

sustaining far-traveling tube-fed flows. 

The Island of Hawaiʻi is not the only one for which lava 

flow hazards have been mapped. Crandell (1983) constructed a 

lava flow hazards map for Haleakalā volcano (Maui), based on 

ages of recent flows and vent distribution, using only five zones; 
Natural Hazards and Risk Reduction in Hawai‘i 401

Figure 2. Scan of unpublished Hawaiian Volcano Observatory map (drafted in the 1940s or 1950s) that estimated lava flow incidence per unit 

area per 10,000 years, before dating of precontact (before 1778) lava flows became possible.
402 Characteristics of Hawaiian Volcanoes

historical frequency and volume of lava flows 

mapped within the Hilo city limits (U.S. 

Army Corps of Engineers, 1980). This study 

also estimated the potential socioeconomic 

impact of lava flow inundation in Hilo, in 

part to ascertain whether the construction 

of permanent lava flow barriers upslope 

of Hilo (discussed later) was warranted. 

This study found the logarithmic values 

of the probability of occurrence (lava flow 

inundation) to be linearly related to the 

logarithmic values of the areal coverage of 

lava—in other words, the larger the target, 

the more likely a future lava flow will enter 

the target area.

Interest in geothermal energy in 

the 1980s and 1990s prompted an HVO 

assessment of the volcanic hazards in 

areas around likely geothermal resources 

along Kīlauea’s East Rift Zone (Moore and 

Kauahikaua, 1993; Moore and others, 1993; 

Kauahikaua and others, 1994). Estimated 

probabilities of lava inundation in these 

areas ranged as high as 60 percent over a 

50-year period. 

A later study estimated lava inundation 

probabilities for several specific sites 

on Mauna Loa’s Northeast Rift Zone, 

where a new prison facility was planned 

(Kauahikaua and others, 1998b). This 

study was, in part, prompted by a concern 

during the 1984 eruption of Mauna Loa 

that the existing Kulani Prison might 

be overrun by lava, and the authorities 

worried about possible evacuation of the 

prisoners (Stapleton, 1984). Probabilities 

this map was later revised by Mullineaux and others (1987, 

fig. 22.14), with zone 1 being the most hazardous of the five 

zones but with slightly different criteria than those defining the 

zones for the Island of Hawaiʻi. In considering the differences, 

Mullineaux and others (1987, table 22.1) equated Maui zone 

1 with Hawaiʻi zone 3, Maui zone 2 with Hawaiʻi zone 4, and 

Maui zone 3 with Hawaiʻi zone 6. Sherrod and others (2006) 

revised the lava flow hazard zones for Haleakalā using new 

mapping and dating but did not address the equivalency of the 

new zones with those of the Island of Hawaiʻi. 

To reconcile the different schemes used, to date, we sought 

to redefine lava flow hazard zones for Maui and the other, older 

volcanoes to produce a lava hazards map for the entire State of 

Hawaii. Based on the newest mapping, Mullineaux’s equivalency 

may overestimate the lava flow hazards on Maui (D.R. Sherrod, 

written commun., 2010). Despite the minor differences between 

the various existing lava flow hazards maps, they are all based 

on the general premise that the summit and rift zone vents pose 

the greatest potential hazards and that the hazard decreases with 

distance from the vents. Because known eruptive vents of lava 

flows on the older Hawaiian Islands do not fall neatly into the 

spatial pattern for Haleakalā and the volcanoes on the Island of 

Hawaiʻi, no lava flow hazards maps have been published for 

volcanoes northwest of Haleakalā along the Hawaiian volcanic 

chain. After considering the data now available and redefining the 

lowest hazard category (zone 9) to include all areas not inundated 

by lava in more than 10,000 years, we have compiled a preliminary 

lava flow hazard map for the eight main islands in the State of 

Hawaii (fig. 4).

Probabilistic Estimation

Probability estimation can provide a solid numerical basis 

for comparing lava flow hazards to, for example, tsunami or 

hurricane hazards and can be used directly to estimate risk. 

Like the lava flow hazard maps, probabilities are based on a 

geologic map, complete with lava flow ages. 

The first effort to estimate probabilities of lava inundation 

for Hilo, the second largest city in the State, was based on the 

Mauna Kea

Kohala

9

8

7

4

6

6

5

3

3

3

2

2

Hualālai

Kīlauea

Mauna Loa

3

2

2

1

1

1

1

1

156° 155°

20°

19°

Lava flows

since 1800

1

2

3

4

5

6

7

8

9

Increasing severity of hazard

Hilo

KailuaKona

MLO

Kapoho

Kealakekua

Bay

Puʻu

ʻŌʻō 

0

0

20 40 KILOMETERS

10 20 MILES

Figure 3. Generalized version of the most recent lava flow hazards map of the Island 

of Hawai‘i, updated and revised by Wright and others (1992). The severity of the hazard 

increases with decreasing zone number. Thick black lines separate the five individual 

volcanoes on the island. Boxes denote communities discussed in the text. Triangle 

shows location of Pu‘u ‘Ō‘ō eruptive vent. “X” gives location of the National Oceanic and 

Atmospheric Administration’s Mauna Loa Observatory (MLO). 
Natural Hazards and Risk Reduction in Hawai‘i 403

of lava inundation, 12 percent or less over a 50-year period, 

were again determined to be proportional to the area of 

interest. In that assessment, steepest descent lines and 

lavasheds (calculated as watersheds) were determined from 

the best available DEMs and plotted to forecast lava paths 

and maximum warning times after the start of a threatening 

eruption (no more than 72 hours). 

Flow-Path Forecasting

Qualitative and probabilistic lava flow hazard maps 

both provide a long-term perspective of the threat posed by 

lava flows, although the probabilities may be expressed over 

different time frames. These long-term hazard assessments 

are useful for planning on a time scale of decades to 

centuries but not very useful for shorter time frames. For the 

estimation of more immediate lava flow threats, it would be 

more helpful to have the ability to simulate, and therefore 

predict, the path and advance rate of lava flows while they 

are active during an ongoing eruption.

Although lava flow simulation software may not yet 

be accurate enough to use in all situations, even a simple 

forecast of their paths based only on digital topography 

can be valuable. In this regard, it is noteworthy that, in 

1912, Jaggar convinced Governor Frear of the Territory of 

Hawaiʻi that a topographic map of the Island of Hawaiʻi was 

necessary for the purpose of forecasting where lava flows 

might go once they were erupted.

Figure 4. Preliminary integrated lava flow hazards map for the State of Hawaii. Zones 1 through 8 are 

taken from Wright and others (1992) for the Island of Hawai‘i, and zone 9 is defined as areas that have 

not experienced eruption or unrest in the past 10,000 years.

160° 159° 158° 157° 156° 155°

20°

22°

19°

21°

1

2

3

4

5

6

7

8

9

Increasing severity of hazard

0

0

50 100 KILOMETERS

25 50 MILES

“It will have to be done anyway, some day,” said 

the governor yesterday, and we quite agree with 

Professor Jaggar that it is better to do it now, for 

a flow may occur at any time. The new survey 

will indicate what directions the flows will take, 

and for this reason the work cannot help but be of 

intense interest to all our people. The householders 

and others on the slopes are, of course, more 

immediately interested than any others.

 —Pacific Commercial Advertiser (1912) 

There is no evidence, however, that Jaggar ever used 

topographic maps to forecast lava flow paths.

During the past two decades, improved maps and 

DEMs have allowed quantitative forecasting of lava flow 

paths. Trusdell and others (2002) defined lava-inundation 

zones for Mauna Loa, based on the paths of mapped flows 

modified for current topography, and Kauahikaua and others 

(1998a, 2003) and Kauahikaua (2007) calculated steepest 

descent paths and lavasheds to define future flow paths. 

Limitations of this approach include unknown sensitivity 

to DEM accuracy and relevance to only the first flows from 

the vent—subsequently the topography changes as a result 

of lava flow emplacement and the initial steepest descent 

forecasts become out-of-date and inapplicable. 

Nonetheless, even with this simplest of approaches, the 

ability to assess sensitivity of the paths to DEM inaccuracies 

is very useful. Favalli and others (2005) developed a method 

to demonstrate flow-path 

sensitivity to possible DEM 

inaccuracies by rerunning 

the calculations with random 

additions to the DEM each 

time. A compilation of the 

results of several iterations of 

this approach make it clear 

which DEM elements have 

critical control on the direction 

of steepest descent. Favalli’s 

algorithm is now incorporated 

in many lava-simulation 

programs (for instance, 

FLOWGO; Harris and 

Rowland, 2001). Rowland and 

others (2005) used FLOWGO 

with the Favalli algorithm 

to forecast the eruption-ratecontrolled extent of lava flows 

from Mauna Loa. Advantages 

of lava flow simulation include 

the ability to forecast multiple 

or compound lava flows; 

however, errors in the early 

forecasts will propagate to 

affect forecasts of later flows.
404 Characteristics of Hawaiian Volcanoes

Mitigation of Risk from Lava Flow Hazards

Because there is no known way to stop an eruption, the main 

method of mitigating the risk posed by active lava flows is to divert 

the flow away from populated areas or to slow its advance to allow 

other mitigation efforts, such as evacuation (see, for instance, 

Peterson and Tilling, 2000). Diverting or delaying lava flows 

have been attempted by the use of explosives, water, and physical 

barriers. Any decision to attempt to control lava flows, of course, 

must be made by officials charged with land-use planning and (or) 

emergency management. 

Many forms of lava flow diversion have been tried on the 

Island of Hawaiʻi, with varying degrees of success. The first known 

efforts occurred in the summer of 1881 as a broad pāhoehoe flow 

from Mauna Loa slowly advanced toward the town of Hilo. Low 

earthen and rock walls were set up, but the advancing flow stalled 

on its own without evidence of being rerouted (Hawaiian Gazette, 

1881). It is interesting to note that explosives also were authorized 

by the Hawaiian Kingdom for possible diversion of this flow, but 

the materials arrived too late to be used. 

During an intense earthquake swarm beneath Hualālai 

volcano in late 1929, a seemingly imminent eruption of lava 

prompted Lorrin A. Thurston to suggest the use of explosives again 

(Honolulu Advertiser, 1929a). Jaggar agreed that if the explosives 

were placed along a feeder tube, the blast could disrupt the flow 

and cause lava to run over lands that had already been covered 

(Honolulu Advertiser, 1929b), thereby minimizing damage. The 

military went so far as to assess that the effort was feasible, but the 

anticipated eruption never took place.

Jaggar had his first opportunity to use explosives in 1935, 

when a Mauna Loa lava flow threatened the water supply above 

Hilo. He worked with a group of Army aviators who planned 

and executed a bombing mission, targeting sites along the upper 

channels of the advancing flow. The bombing was very precise 

and was completed days before the eruption ended. Jaggar claimed 

success (for instance, Jaggar, 1945a, p. 12–16), but most other 

volcanologists believe that the flow stopped because the eruption 

shut down (Lockwood and Torgerson, 1980). In 1942, under 

very similar circumstances, bombs were again used on Mauna 

Loa to blast a spatter rampart along a fissure vent feeding a flow 

advancing on Hilo. The rampart was successfully breached and 

lava diverted, but only for a short distance before it rejoined its 

original channel, effectively negating the diversion attempt. 

The failure of the 1942 diversion suggests that the 

results of future bombing efforts might be improved by taking 

advantage of the steepest descent path maps. If the channel 

had been bombed at a point where the new outflow could be 

directed into a steepest descent path distinct from the original 

one, the bulk of the lava flow might have been diverted away 

from the original channel. Any remaining flow in the original 

channel would then be greatly diminished and pose less threat 

to areas downchannel.

After his retirement in 1940, Jaggar published a detailed 

proposal to divert Mauna Loa lava flows away from Hilo 

(Jaggar, 1945b), but the next opportunities to erect barriers 

were presented on Kīlauea Volcano. During the 1955 and 

1960 Kīlauea eruptions in the Puna District, several barriers 

were built to divert or dam lava flows (fig. 5) in attempts to 

protect downstream homes and farms. Neither effort was 

successful, although it can be argued in both cases that lava 

inundation may have been delayed. Gordon Macdonald, 

the HVO Scientist-in-Charge during the 1955 eruption and 

a representative of the Governor during the 1960 eruption 

(which occurred after he left the USGS), was an advocate for 

lava diversion through the use of permanent barriers to protect 

Hilo (see, for instance, Macdonald, 1958). Unfortunately, 

most of the HVO staff concluded that the plan was “expensive 

beyond prudent economic justification” (Wentworth and 

others, 1961), and a very public debate over diversion ensued 

in the local press in 1960. Barriers have not been used since 

1960, but during the 1980s, much effort went into studying 

the feasibility of building permanent diversion barriers above 

Figure 5. Photograph showing bulldozers constructing a lava flow barrier in 

Kapoho in January 1960. Note the advancing front of an ‘a‘ā flow between the lava 

fountain and the trees at left. Date of photograph and photographer unknown.

Hilo, as both Jaggar and Macdonald had advocated 

(U.S. Army Corps of Engineers, 1980). That 

study recommended construction of an emergency 

barrier when needed rather than construction of a 

permanent barrier.

The use of water to cool and solidify the 

advancing lava front—causing the lava to form 

its own barrier—has been attempted in Hawai‘i 

and was later also tried on a much larger scale in 

Iceland (see, for instance, Williams and Moore, 

1973; Williams, 1997). Water, pumped from a 

nearby lake, was used during the 1960 Kapoho 

eruption to delay consumption of houses by fire 

upon lava contact. In 1989, when lavas from 

Kīlauea’s East Rift Zone were slowly engulfing 

the Waha‘ula Visitor Center within Hawaiʻi 

Volcanoes National Park, water was again used in 

an experiment to delay consumption of the wooden 

structure. Although some delay was achieved, the 

structures ultimately burned to the ground (fig. 6).
Natural Hazards and Risk Reduction in Hawai‘i 405

The current State Lava Flow Hazard Mitigation Plan, 

written by a committee that included HVO and University of 

Hawaiʻi scientists, found that lava diversion barriers were not 

appropriate in most situations; however, some critical facilities 

may be situated in areas where barriers could be a reasonable 

option (Hawaii State Civil Defense, 2002). For example, 

lava diversion barriers (5–7 m high, 700 m long, as designed 

by HVO scientists) were completed in 1986 (Moore, 1982; 

Mims, 2011) to protect the National Oceanic and Atmospheric 

Administration’s (NOAA) Mauna Loa Observatory, which is 

located on the north flank of Mauna Loa volcano (fig. 3). The 

facility is a premier atmospheric research facility that maintains 

the continuous record of atmospheric change since the 1950s 

and is the site of the measurements forming the well-known 

Keeling curve of CO2 concentrations in the atmosphere. 

The 2002 State plan discussed above considered, but 

was not based on, cultural objections, but concerns have been 

raised about some diversion strategies. Gregg and others 

(2008) mined interview data obtained from Kapoho residents 

shortly after their town was destroyed by a Kīlauea eruption in 

1960 and found that ethnic Hawaiians favored the construction 

of earthen barriers but did not support the use of bombs for 

the diversion of lava flows. This finding echoes sentiments 

expressed by the Hawaiian community in response to the 

bombing of Mauna Loa lava flows in 1935 (Jaggar, 1936), as 

well as later interviews for the Hilo barrier feasibility study 

(U.S. Army Corps of Engineers, 1980). Only earthen barriers 

or dams were built to impede the advance of flows during the 

inundation of Puna in 1955 and Kapoho in 1960. Bombing 

was discussed but dismissed when it was clear that such 

measures would be of no use (Wilhelm, 1960). 

Current Status

Potential lava flow sources and paths have been mapped, 

and general lava flow advance rates were estimated, as a 

function of effusion rates, from historical data (for instance, 

Kauahikaua and others, 2003). Methods for estimating 

the probabilities of lava flow inundation using geographic 

information system (GIS) software on an islandwide basis 

are in development (Trusdell, 2010). Although the diversion 

of lava flows around critical facilities may be technically 

feasible, it is clearly a social and political issue whose solution 

is beyond the scope of this review. For the present, we are 

able to characterize the potential paths of lava flows once their 

eruptive sources are identified; however, only broad guidelines 

can be provided for their rates of advance until accurate and 

reliable lava flow simulators are developed.

Volcano Hazards: Explosive Eruptions

Although lava flows constitute the most frequent (and, 

therefore, most probable) volcanic threat for the islands of 

Maui and Hawaiʻi, the less frequent occurrences of explosive 

A

B

C

Figure 6. Photographs of the National Park Service Waha‘ula Visitor 

Center near the coast on Kīlauea Volcano in 1989 and of efforts 

to protect it from approaching lava. A, Fire hose delivers water to 

cool approaching pāhoehoe flows (June 22, 1989). B, Visitor Center 

engulfed in flames later that day after being ignited by hot lava. C, 

Steel girders, twisted by inflating lava flows and later buried, are all that 

remain of the Visitor Center after it burned (October 12, 1989). U.S. 

Geological Survey photographs by Jim Griggs.
406 Characteristics of Hawaiian Volcanoes

eruptive activity can pose significant, more widespread 

hazards. Indeed, the explosive blast and gases produced 

during the 1790 eruption at Kīlauea’s summit (discussed 

below) resulted in the most lethal volcanic disaster in recorded 

U.S. history, and there is growing evidence of even stronger 

Hawaiian explosive eruptions in the geologic record.

In contrast to explosive composite volcanoes related to 

subduction zones (for example, Mount St. Helens, United 

States; Mount Fuji, Japan; Pinatubo, Philippines), intraplate 

shield volcanoes, such as those of Hawaiʻi, are thought to 

typically erupt nonexplosively. Deposits of tephra (ash fall) 

and other pyroclastic debris in Hawaiʻi, however, attest to 

the occurrences of powerful explosive eruptions of Hawaiian 

volcanoes in the geologic past. Recent studies show that 

the frequency of explosive eruptions at Kīlauea is about the 

same as that for Mount St. Helens—several explosive periods 

every millennium (Swanson and others, 2011). Two important 

differences, however, should be noted: (1) Kīlauea explosive 

eruptions are much smaller in volume and in area affected than 

those at Mount St. Helens and (2) during the intervals between 

explosive periods, Mount St. Helens is mostly inactive, whereas 

effusive eruptions occur frequently, sometimes essentially 

nonstop (as since 1983), at Kīlauea. Not surprisingly, therefore, 

the Hawaiian eruptive style is usually considered to be gentle 

or benign; the term “Hawaiian” is used to describe any 

nonexplosive or weakly explosive eruption in the world with a 

Volcanic Explosivity Index (VEI) of <1 (for instance, Siebert 

and others, 2011). Because of their infrequency, relative to 

current effusive activity, explosive Hawaiian eruptions and their 

associated hazards have been less studied until recently. 

Understanding Basaltic Explosive Eruptions and 

Their Products

Kilauea has not always been a quiet volcano, as it is 

today. . . . 

 —Sidney Powers (1916, p. 227)

Kīlauea has long been recognized by Hawaiians as a 

volcano with a dual personality. In Hawaiian mythology, 

Pele, the goddess of Hawaiian volcanoes, whose home is 

in Halemaʻumaʻu Crater within Kīlauea Caldera, is always 

described with two personas: a young, beautiful woman and 

an old, cruel hag. Hawaiian traditions and oral histories also 

support that Kīlauea has erupted explosively since the island 

was settled (see, for example, Swanson, 2008; Kanahele, 2011). 

There have been three historical (postcontact) explosive 

eruptions of Kīlauea Volcano. In November 1790, in the 

most energetic of these three, an estimated 80 warriors, 

and as many as 400 people in all, were reported to have 

been killed near the volcano’s summit (Ellis, 1825; Dibble, 

1843; Kamakau, 1992). The cause of these deaths has been 

variously attributed to directed blasts of ash and hot gases 

(Swanson and others, 2012b), but the explosive eruption 

mechanism remains unknown.

In May 1924, weeks of explosive activity enlarged the 

diameter of Halemaʻumaʻu Crater, spread sizeable debris a 

few kilometers from the crater, and killed one person (Hilo 

Tribune-Herald, 1924). While much less energetic than the 

1790 eruption, this activity ejected ballistic blocks—some 

more than 1.5 m in average diameter—to distances greater 

than 1 km. The 1924 explosive activity was thought to be 

the result of groundwater entering the magma conduit weeks 

after the draining of a lava lake in the summit (Dvorak, 1992), 

resulting in phreatic explosions. 

Most recently, in March 2008, a small crater opened 

explosively along the east wall of Halemaʻumaʻu Crater, 

spreading sizeable (~0.25–1 m in diameter) debris within 

a few hundred meters of the vent. The explosion was 

hypothesized to be the result of high gas pressure blasting 

out rock debris that had blocked a previously open gas vent 

(Houghton and others, 2011). This eruption at Halemaʻumaʻu 

persists as of September 2014, but mostly as a continuously 

active lava lake that also emits gas and small amounts of 

glassy and lithic tephra (Wooten and others, 2009).

Powerful explosive eruptions have played a large part in 

the evolution of Kīlauea’s summit caldera. The general history 

of Kīlauea’s summit was told by native Hawaiians to the first 

westerner to record his visit (Ellis, 1825, p. 137–138): 

From their account, and that of others with whom 

we conversed, we learned that it had been burning 

from time immemorial, or, to use their own words, 

“mai ka po mai,” (from chaos till now,) and had 

inundated some part of the country during the reign 

of every king that had governed Hawaii. That, in 

earlier ages, it used to boil up, overflow its banks, 

and inundate the adjacent country; but that, for many 

kingʼs reigns past, it had kept below the level of the 

surrounding plain, continually extending its surface, 

and increasing its depth, and occassionally throwing 

up, with violent explosion, huge rocks, or red hot 

stones. These eruptions, they said, were always 

accompanied by dreadful earthquakes, loud claps of 

thunder, vivid and quick-succeeding lightning.

Perret (1913) correctly interpreted the Hawaiians’ 

remarks and combined them with his geologic observations 

to propose the following sequence of events in the geological 

history of Kīlauea:

• formation of a shield by overflows from a central vent,

• subsidence forming a great pit or caldera,

• explosive eruptions producing the observed ash,

• lateral subterranean outflows of lava to the sea.

More recently, HVO scientists and colleagues have filled 

in the previous gaps with new work, including extensive 

mapping and assignment of dates to many of the key events, 

resulting in a more detailed interpretation of the evolution of 

Kīlauea’s summit caldera: 
Natural Hazards and Risk Reduction in Hawai‘i 407

• More than 500 years ago, Kīlauea existed as a shield with 

a central vent that produced repeated overflows (Clague 

and others, 1999).

• Caldera formed by collapse about 500 years ago, followed 

by several centuries of explosive eruptions ending in 1790 

(Swanson and others, 2011; Swanson and others, 2012a).

• The 1790 fatalities were a result of a base surge (Swanson 

and Christiansen, 1973).

• Kīlauea has alternated between periods of dominantly 

explosive eruptions and periods of effusive eruptions 

(Swanson and others, 2011). 

Older explosive eruptions produced more widespread 

deposits of volcanic ash that were recognized by a number 

of studies completed in the late 19th and early 20th century. 

Bishop (1887) traced a yellow ash layer around the South Point 

of the Island of Hawaiʻi and deduced that it 

. . . was formed by an explosive eruption of yellow 

cinder, which covered at least 100square miles...

with yellow ashes several feet in thickness. It must 

have belonged to the larger class of explosive 

eruptions. I hereby file my caveat for this discovery, 

in case no one has recorded a patent of prior date.

Emerson (1902), Powers (1916), and Stone (1926) 

described the Pāhala Ash, which is widely distributed across 

the southeastern part of the island and may have been Bishop’s 

“yellow ash.” A very useful compilation and detailed description 

of ash deposits on the Island of Hawaiʻi was published by 

Wentworth (1938), who identified 10 mappable ash units on 

Hualālai volcano. Water wells south of Kona have been drilled 

through ash layers that may have originated from Mauna Kea 

eruptions (Bauer, 2003). Drilling also penetrated ash layers at 

shallow depths at locations in Volcano Village, near Kīlauea’s 

summit, and in the area between the town of Mountain View 

(halfway between Hilo and Kīlauea’s summit) and Mauna 

Loa’s Northeast Rift Zone (Stearns and Macdonald, 1946), 

expanding the known fallout area for past explosive deposits. 

The Hawaii Scientific Drilling Project (HSDP) in Hilo also 

found multiple ash layers deposited by explosive eruptions 

from Kīlauea and Mauna Kea; most were thin, with only two 

major explosive eruptions depositing thick ash at the drill site 

in the past 400,000 years (Beeson and others, 1996).

Characterizing Tephra and Other Hazards of 

Explosive Eruptions

The occurrence of widespread ash deposits (fig. 7) 

is clear evidence that powerful explosive eruptions have 

occurred in Hawaiʻi’s geologic past, albeit infrequently, 

compared to effusive eruptions. Moreover, the studies of 

the lethal 1790 eruption at Kīlauea suggest that the deaths 

in that event resulted from the combined hazards of ash fall 

(tephra) and ground-hugging, high-velocity pyroclastic surges 

(see, for instance, Swanson and Christiansen, 1973), and 

possibly lethal gases. These hazards and their significance 

were evident to Jaggar (1918, p. 16), who emphasized that 

“Hawaii has had Class 1, explosion and volcanic blast

[bold in original], in 1790, from Kilauea crater. . . .” He 

states further that “A repetition of that event would probably 

Kona Tuff Formation

Hualālai pyroclastic materials

Waawaa pyroclastic materials

Keanakāko‘i Formation

Waimea “agglomerate”

Waiau Formation

Kohala pyroclastic materials

Kapoho pyroclastic materials

No significant ash

Nīnole Tuff

Pāhala Tuff

156° 155°

19°

20° 1000 2000

1000

3000

3000 2000

0

0

20 40 KILOMETERS

10 20 MILES

Ash deposits

Figure 7. Map showing 

generalized distribution of 

pyroclastic deposits on the 

Island of Hawai‘i (colorized 

from Wentworth, 1938, figure 

6). The units in Wentworth’s 

legend for the deposits (above) 

are not necessarily ordered 

according to relative age. 

Elevation contours in meters.

the island (fig. 7). Although 

some work has been done on the 

ash deposits since Wentworth’s 

compilation, reliable ages on the 

pre-1790 deposits are still few. 

The Pāhala Ash was inferred by 

Decker and Christiansen (1984) to 

represent the cumulative deposits 

of many explosive eruptions from 

about 25,000 to 10,000 years ago. 

More recent work suggests that the 

youngest unit in the Pāhala Ash 

is about 13,000 years old (F.A. 

Trusdell, oral commun., 2012). 

Ash layers have also been 

discovered below the surface in 

archaeological excavations and 

water-well drill holes. While 

excavating archaeological sites 

along a highway track through 

Kailua-Kona, Schilt (1984, 

p.275) encountered a shallow 

ash layer that was dated by the 

carbon-14 method at 1260–1485 

C.E. and was probably from 
408 Characteristics of Hawaiian Volcanoes

wreck [everything] for a radius approximately six miles from 

Halemaumau in all directions.” Jaggar also emphasized that, 

with strong trade winds, “heavy ash falls would be in the 

[downwind] Kau direction [to the southwest from Kīlauea’s 

summit].” The Hawaiian Volcano Observatory, at its current 

location atop Uēkahuna Bluff, would be destroyed. 

Using only the record of the past 250 years of volcanic 

activity, Mullineaux and others (1987) applied the criteria of 

proximity to summit and rift-zone vents, eruption frequency, 

and wind direction to produce a tephra hazards map (fig.8). 

Separately, they noted that the potential for pyroclastic surges 

existed only near Kīlauea’s summit (within 10 km of the center 

of the caldera—virtually identical to Jaggar’s Class 1 hazard). 

A comprehensive modern assessment of tephra hazards 

for Hawaiʻi would require full consideration of the older 

tephra deposits throughout the island, together with a longterm record of dominant wind trajectories for Hawaiʻi. Though 

highly approximate and dated, Wentworth’s (1938) map (fig. 7) 

nonetheless suggests that the significant tephra deposits from 

larger, much less frequent, explosive events can be distributed 

much farther than the tephra hazard zone 1 boundaries (fig. 8) 

or the Kīlauea pyroclastic surge zone of Mullineaux and others 

(1987), and that the most voluminous deposition will occur 

downwind along dominant wind trajectories. For airborne 

ejecta from Kīlauea eruptive vents, which are all at elevations 

below the thermal inversion layer that starts at ~2 km above 

sea level, direction of deposition would generally be controlled 

by the northeasterly trade winds. 

Above the thermal inversion layer, however, the 

dominant wind patterns are different. Swanson and others 

(2011) attributed tephra distribution southeast of Kīlauea 

Volcano to explosions ejecting tephra above the thermal 

inversion and into the more westerly jet stream wind 

currents. Any future comprehensive long-term assessment 

of tephra hazards and other pyroclastic hazards will require 

the incorporation of new mapping and radiometric age data 

on regional ash deposits acquired since the early 1990s (for 

instance, Buchanan-Banks, 1993; Beeson and others, 1996; 

Wolfe and Morris, 1996; Sherrod and others, 2007). 

Current Status

The details of Kīlauea and Mauna Loa’s explosive past 

are beginning to emerge, and hazards assessments must include 

the possibility of a large, damaging explosive eruption as a 

maximum credible event, along with more frequent effusive 

eruptions, which represent a more probable next event. 

Arguably, the biggest scientific and monitoring challenge 

associated with hazards due to explosive eruptions is the 

recognition of precursory signs of explosive activity. Would 

such precursory indicators differ diagnostically from those that 

precede effusive eruptions or magma intrusions? With just three 

historical examples, only the smallest of which (in March 2008 

at Halema‘uma‘u) was well documented geophysically, HVO 

must be constantly aware of the possibility—however remote—

of explosive activity. To date, no mitigation measures or plans 

have been established for hazardous explosive processes.

Volcano Hazards: Gas Emissions

The past century of frequent Kīlauea and Mauna 

Loa eruptions has made Hawaiʻi one of the premier world 

laboratories for studying volcanic gases (Sutton and Elias, this 

volume, chap. 7). Long-term datasets on gas emissions are 

used to assess the global contribution of volcanoes to climate 

change (for instance, Gerlach, 2011), and also provide insights 

into volcanic behavior that can lead to more accurate forecasts 

of eruptive activity. Over the past decades, however, gas 

studies have taken on a new urgency as the island and State 

are affected by volcanic smog (vog) produced by Kīlauea’s 

persistent eruptions. Vog poses a health hazard by aggravating 

preexisting respiratory ailments, and acid rain resulting from 

vog damages crops and can leach lead from metal roofs into 

household water supplies (Sutton and others, 2000).

156° 155°

20°

19°

1

3A

2

2

2

1

1

1

2A

3

3

2

0

0

20 40 KILOMETERS

10 20 MILES

Figure 8. Shaded relief map showing tephra hazard zones on the 

Island of Hawai‘i (after Mullineaux and others, 1987, figure 22.11); red 

areas represent historical lava flows. Zone 1 includes areas of highest 

eruption frequency (the summit and rift zones of Kīlauea and Mauna Loa). 

Zone 2 includes areas where tephra falls from lava fountains should be 

frequent but thin, with subzone 2A depicting locally thick tephra falls due 

to infrequent eruptions of Hualālai. Zone 3 covers areas where very thin 

tephra falls may occur from eruptions on Kīlauea, Mauna Loa, and Hualālai 

under unusual northerly wind conditions; subzone 3A may be subject to 

very infrequent, but possibly severe, eruptions from Mauna Kea. 
Natural Hazards and Risk Reduction in Hawai‘i 409

HVO was a pioneer in studies of volcanic gas and, since 

the late 1970s, has systematically monitored gas emissions at 

both Kīlauea and, more recently, Mauna Loa. Comprehensive 

discussions of the results of volcanic-gas studies in Hawaiʻi and 

their bearing on eruptive processes are given elsewhere (for 

instance, Gerlach and Graeber, 1985; Greenland, 1987a, 1987b; 

Sutton and Elias, 1993, this volume, chap. 7). Here, we briefly 

summarize the hazards to humans, animals, and vegetation 

that are posed by exposure to volcanic gases from frequent 

and sustained eruptive activity in Hawaiʻi. The International 

Volcanic Health Hazard Network (http://www.ivhhn.org/), a 

Commission of the International Association of Volcanology 

and Chemistry of the Earth, is a useful clearinghouse for 

information about health issues related to volcanic gases. 

Volcanic Gas Composition, Emission-Rate 

Variations, and Dispersion 

Volcanic gas emissions precede and accompany eruptions 

and are composed mainly of H2O (typically 70–90 percent), 

CO2, and SO2, but they also contain varying trace amounts 

of other gaseous compounds (for example, HCl, H2S, H2SO4, 

CO, Ar, HF, F). Typical gas-emission rates measured for 

Kīlauea range from 100s to more than 10,000 (metric) tons per 

day of SO2 and 8,000 to more than 20,000 tons per day of CO2

(Gerlach and others, 2002; Poland and others, 2012). Even 

though fluorine (F) is a minor component, it still can have 

significant harmful impacts to crops and farm animals in small 

amounts; HF emissions have been estimated to be 7–12 tons 

per day for Kīlauea (Sutton and Elias, 1993). 

Gas emissions vary in composition, depending on 

whether a vent is at the summit or within the rift zone of a 

volcano. Carbon dioxide is one of the first gases to exsolve 

from ascending magma and, therefore, dominates nonwater 

emissions at the summits of Hawaiian volcanoes, which are 

generally thought to overlie the deeper magma conduits and 

storage chambers. Sulfur dioxide is one of the last gases to 

exsolve and is associated with shallow magma storage and 

eruptions, both at summits and along rift zones (Gerlach and 

Graeber, 1985). As an example, the opening of a vent at the 

summit of Kīlauea Volcano in 2008 dramatically increased the 

summit emissions of SO2 (Wilson and others, 2008; Sutton 

and Elias, this volume, chap. 7).

Once a volcanic plume rises into the atmosphere, it drifts 

and becomes widely dispersed, depending on wind patterns 

and other atmospheric conditions (for instance, Sutton and 

others, 2000). Trade winds (from the northeast) are dominant 

during summer months and weaken during the winter. As 

noted earlier, the thermal inversion that forms at an altitude 

of ~2 km during trade-wind conditions often separates wind 

directions that can be substantially different above and below 

(see, for example, Schroeder, 1993). Gas and ash from plumes 

rising above 2km (or from vents at altitudes above 2 km) will 

generally be dispersed to the east or northeast by the jet stream, 

whereas gas and ash from plumes or vents lower than 2 km 

altitude will usually be dispersed to the southwest by the trade 

winds (Swanson and others, 2011). All of Kīlauea’s vents are 

below 2km, but most of Mauna Loa’s are above this elevation.

Volcanic Gas Hazard

Vog (volcanic smog) and acid rain are the most common 

hazards related to volcanic gases in Hawaiʻi. These are 

produced when volcanic gases (primarily SO2) and sulfate 

aerosols (tiny particles and droplets) react with atmospheric 

moisture, oxygen, and sunlight (Sutton and others, 1997). 

Close to vents, the emissions are mostly SO2 gas; farther 

from the vents, the emissions become dominantly aerosols 

with particulate matter that is less than 2.5 μm in diameter 

(referred to in air-quality discussions as PM2.5). Vog can 

pose a health hazard by aggravating preexisting respiratory 

ailments, and acid rain can damage crops and leach lead from 

metal roofs into household rain-catchment water supplies.

The word “vog” was coined in the middle to late 20th 

century (Watanabe, 2011), reflecting the relatively recent 

recognition of volcanic gas from Hawaiian volcanoes as 

a hazard. Before the start of the ongoing East Rift Zone 

eruption of Kīlauea in 1983, written observations of volcanic 

smoke, haze, or vog were uncommon, and public complaints 

were rare. Through the many notes of travelers passing the 

summit of Kīlauea since 1823, we know that the volcano was 

obviously emitting gases; however, reports of volcanic-gas 

effects away from Kīlauea were infrequent. Reports of gases 

during Mauna Loa activity, however, were common across 

the entire island chain, possibly because of the great volumes 

of gas that were released over short periods of time by 

discrete Mauna Loa eruptions from vents above the inversion 

layer. Gas emitted from such high vents may be blown to the 

northeast until falling below the inversion layer and being 

blown back to the islands by the trade winds.

Early Descriptions

Many visitors to the active lava lake at Kīlauea’s summit 

throughout the 1800s and until 1924 described the choking, 

suffocating effects of the volcanic gas immediately downwind 

of the lava lakes, but there were very few mentions of gas 

beyond the summit area until the mid-20th century. In the early 

1800s, missionary John Whitman commented on what may 

be the earliest mention of vog from Kīlauea Volcano (Holt, 

1979): “. . . on the N.W. part [of the island] a thin blue smoke 

may be observed coming from a volcano which is described 

by the natives as a lake of burning lava . . .”; this may be the 

earliest mention of vog from Kīlauea Volcano (Holt, 1979). 

On Oʻahu, more than 300 km northwest of the Island of 

Hawaiʻi, southerly, or “kona,” winds were commonly described 

as “volcano weather,” bringing rain and, sometimes, hazy 

conditions (Lyons, 1899). 

Mauna Loa apparently produced large amounts of gas 

during the weeks-to-months-long eruptions in the 19th and 
410 Characteristics of Hawaiian Volcanoes

20th centuries, and the emissions travelled far from their 

sources. The Hawaiian Gazette (1868) reported that Oʻahu 

was “wrapped in smoke” after the Mauna Loa 1868 eruption. 

Lyons (1899) reported that a haze covered the entire island 

group after the 1877 Mauna Loa eruption, and a newspaper 

account (Hawaiian Gazette, 1899) described sulphurous smells 

and minor amounts of ash enveloping the city of Honolulu as 

a result of the 1899 Mauna Loa eruption. The 1950 eruption 

of Mauna Loa produced a visible haze that extended more 

than 3,000 km to Midway Islands atoll (Free Lance-Star, 

1950). Initially, the source of the haze was not known, but its 

significant sulfate content and particulate content, 600 times 

the normal amount, pointed to a volcanic origin. In each of 

these cases, the eruptions occurred at high-elevation vents, 

and the emissions were dispersed away from Hawaiʻi by 

high-altitude winds; when those winds were to the northeast, 

the emissions would eventually descend into the troposphere 

and return to Hawai‘i with the trade winds. Kīlauea eruptions 

produce significant volcanic emissions from vents at lower 

tropospheric elevations than those from Mauna Loa. During 

the Kapoho eruption (from Kīlauea’s East Rift Zone; fig. 3) 

in 1960, SO2 concentrations were so high in Hilo, 35 km to 

the northwest, that the Governor of Hawaii briefly considered 

evacuating the city (Hilo Tribune-Herald, 1960). 

The increase in public concern about vog over the 

past 30 years is puzzling in light of the relative absence of 

concern during earlier years of sustained activity at Kīlauea. 

The lava lake at Kīlauea’s summit, which was active nearly 

continuously from its first documented visit in 1823 through 

1924, must have produced copious amounts of volcanic gas, as 

does the current summit lava lake. The south and west (Kona) 

coast should therefore have been just as impacted by vog as 

they currently are, due to eruptions at both Halemaʻumaʻu 

and Puʻu ʻŌʻō; however, this may not have been the case. 

The Kona coast in the early 1900s reputedly had a very clean 

and dry atmosphere, such that Kona was often mentioned 

as a place of refuge for those with tuberculosis and other 

respiratory problems (Goodhue, 1908). The Kona air at that 

time was described as free from fog, dust, and mud and was so 

clear that it provided unlimited views to the north and south. 

The Kona Hospital was established in 1909 near Kealakekua, 

in part because of the excellent air quality. These conditions 

are in marked contrast to those during the present ongoing 

eruptions of Kīlauea Volcano. Perhaps a careful study of the 

climatological and sociological records of the Kona region 

might yield some clues as to the source of this discrepancy. 

Gas Hazard Maps

The earliest known published assessment of gas hazards 

on the Island of Hawaiʻi was that of Mullineaux and others 

(1987). In their analysis, they commented that, during the 

frequent, nearly continuous eruptive activity from 1967 to 

1974, “trade winds carried gases from Kilauea’s summit and 

east rift zone southwestward into the Kau District, reportedly 

causing a decline in sugar yields. Fumes then drifted around 

to the Kona District on the west coast and were blamed for 

the decline of other crops” (Mullineaux and others, 1987, 

p. 611). They also noted that, during the 1977 East Rift Zone 

eruption, drifting volcanic gases killed vegetation as far as 

30 km from the eruptive vent.

Mullineaux and others (1987) stated that a gas hazard 

map would be essentially identical to their tephra hazard 

map (see fig.8), rationalizing that the severity of gas hazards 

would similarly decrease with increasing distance from the 

expected vents and be commensurate with the short-term 

frequency of expected eruptions. SO2 and CO2 emissions 

can reach fatal levels within their gas hazard zone 1 during 

eruptions, judging from the gas-monitoring data for the longlived current eruption(s) at Kīlauea. 

Nonfatal, but damaging, effects have been demonstrated 

at even greater distances downwind from eruptive vents. The 

continuing persistence of vog from the ongoing Puʻu ʻŌʻō 

eruption—reducing visibility and, sometimes, affecting people 

with respiratory conditions—has been a chronic complaint 

at many locations on the Island of Hawaiʻi during both Kona 

(southerly) and trade-wind (northeasterly) conditions. During 

Kona wind conditions, neighboring islands are also affected. 

Although vog was already a health and agriculture nuisance 

during the first 25 years of the Puʻu ʻŌʻō eruption, the 

volcanic-gas problem worsened when the summit vent opened 

in early 2008 and began emitting larger amounts of gases. The 

total emissions from the volcano were augmented for several 

months before the emission rate settled down to a fairly 

steady-state rate that is still as much as three times higher than 

the total pre-2008 emissions from Kīlauea.

A map of nonfatal, but potentially damaging, longerterm exposure to volcanic gas can be estimated by 

determining the distribution of gas constituents within lichen 

around the island (Notcutt and Davies, 1993). Lichens are 

an ideal medium, because their fluoride uptake is from the 

atmosphere and not from the substrate. Typical life span of 

the sampled plants was estimated to exceed 30 years, thereby 

providing good long-term averages. The pattern of fluoride 

abundances clearly shows the general dispersion pattern of 

gases from Puʻu ʻŌʻō; as expected, the highest concentrations 

are found within Kīlauea’s summit region (fig. 9). Airquality summaries based on monitoring by the Hawaii State 

Department of Health’s Air Quality Branch tell a similar story 

about SO2 dispersion, but also include data on particulate 

matter (PM2.5) concentrations (Fuddy, 2011). Environmental 

Protection Agency (EPA) SO2 standards were exceeded most 

frequently in the towns of Pāhala, directly southwest of 

Halemaʻumaʻu, and Mountain View, to the north of Puʻu ʻŌʻō 

(both locations are within 30 km of active vents), whereas 

EPA PM2.5 exceedances are most frequent in Kona, the most 

distant monitoring site from Kīlauea on the Island of Hawaiʻi. 

The U.S. Department of Agriculture issued a Secretarial 

Disaster Designation for vog damage to agriculture (plants, 

animals, and infrastructure) on the Island of Hawaiʻi in July 

2008 (Sur, 2012), but it was discontinued in 2012 because 

of lack of public interest. Damage from the gas emissions 
Natural Hazards and Risk Reduction in Hawai‘i 411

has ranged from acid burns on commercial flower crops to 

possible fluorosis in cattle and goats, as well as accelerated 

degradation of metal fences and gates.

Mitigation of Risks from Exposure to Volcanic 

Gases 

Based on the gas-emission studies pioneered by HVO 

scientists, monitoring of volcanic-gas dispersion is now done 

by several agencies in Hawaiʻi. HVO continues to monitor 

SO2 and CO2 emission rates at the sources (for instance, Elias 

and others, 1998; Elias and Sutton, 2002, 2007, 2012; Sutton 

and Elias, this volume, chap. 7). The Hawaii State Department 

of Health and the National Park Service monitor air quality in 

the State and Hawaiʻi Volcanoes National Park, respectively. 

The air-quality data support mitigation decisions, such as 

evacuation of areas of high gas concentration until the threat 

diminishes. Finally, satellites on daily passes quantify the 

mass of SO2 in the air around Hawaiʻi. Although the results 

are not available to use for timely mitigation, the satellite 

data are important for documenting the SO2 mass distribution 

throughout the State for providing context for ground-based 

measurements (Carn and others, 2012). 

Using all available data, the next step is to forecast vog 

conditions. To this end, a feasibility study was funded by 

HVO at the University of Hawaiʻi at Mānoa Meteorology 

Department, using the computer model HYSPLIT (http://

weather.hawaii.edu/vmap/) with a dense wind-field model and 

the current SO2 emission rate data measured by HVO. The 

results, to date, are promising and popular with the public, 

but the cost to implement such forecasting on a regular basis 

may be prohibitive. Most people are aware of either Kona or 

trade-wind conditions and expect SO2 concentrations to be 

high or low during one or both situations, depending on their 

location. For television media in Hawaiʻi, vog forecasts have 

increasingly become part of the weather forecasts, especially 

if the conditions are expected to be severe.

Public education plays a critical role in providing 

information sufficient for people to make individual 

decisions for coping with vog. The health effects of longterm, chronic exposure to volcanic gases are not known. It 

is known, however, that short-term exposure to vog does 

not cause respiratory illness, although it can exacerbate 

existing problems like asthma. The best health advice has 

been titled “shelter in place,” meaning that, during high 

SO2 concentrations, people should stay indoors in a closed 

or air-conditioned room, if possible (for example, http://

hawaii.gov/health/environmental/air/cab/cab_precautions.

html). Following that policy, Hawaiʻi County and the State of 

Hawaii have made efforts to provide each school and many 

other public facilities with one clean-air room in which those 

with respiratory conditions can find refuge on bad vog days.

Earthquake and Tsunami Hazards 

The frequent volcanic eruptions and earthquakes were 

Jaggar’s main reasons for locating a volcano observatory 

in Hawaiʻi, and the Kīlauea Volcano site afforded abundant 

opportunity to study both (Tilling and others, this volume, 

chap. 1). The first geophysical instruments installed at 

HVO by Jaggar in 1912 were seismometers to record 

both local and distant earthquakes (Okubo and others, this 

volume, chap. 2). HVO staff quickly got a sense that local 

earthquakes occurred frequently and were often related to 

volcanic activity. 

Jaggar wanted to understand earthquakes better but 

concluded early on that the key to minimizing earthquake 

damage lay in the construction of stronger, earthquakeresistant buildings (see, for instance, Jaggar, 1913), in 

addition to understanding the mechanics of earthquakes. 

A major scientific challenge therefore was then—and 

remains today—quantifying the ranges of expected shaking 

forces on which to base design standards for safe building 

construction. Throughout its history, HVO has characterized 

earthquake occurrences and rates with constantly improving 

technology and, in more recent times, has also participated 

in mitigation planning, zoning, and education efforts with 

the Hawaii State Civil Defense agency and the USGS 

Earthquake Hazards Program.

155°

20°

19°

156°

0

0

20 40 KILOMETERS

10 20 MILES

120–332

90–120

60–90

30–60

0–30 

Flouride

concentration,

in μg/g

Figure 9. Shaded relief map showing fluoride concentrations measured 

in lichen sampled around the Island of Hawai‘i expressed in micrograms 

per gram (from Notcutt and Davies, 1993, based on their figure 2).
412 Characteristics of Hawaiian Volcanoes

Hawai‘i Earthquake Patterns and Frequencies

Hawaiʻi, Alaska, California, and Nevada are the four 

most seismically active U.S. states in terms of M>3.5 

earthquakes (Anderson and Miyata, 2006, table 1). The 

County of Hawaiʻi has the third highest annual earthquake 

losses of any county in the United States (Chock and 

Sgambelluri, 2005). If only seismicity related to magma 

accumulation, transport, and eruption is considered (mostly 

M<3.5), Hawaiʻi almost certainly would rank as the first or 

second most seismically active state in the U.S. 

Earthquakes in Hawaiʻi do not occur randomly in space, 

but rather, are concentrated along pathways of the volcanic 

plumbing system (see Tilling and others, this volume, chap. 1; 

fig. 11), fault structures, and also deep within the mantle (Klein 

and others, 1987). Most of the earthquakes in Hawaiʻi are small 

(M<3) and are directly related to eruptions or magma intrusions, 

generally causing minimal damage. Relatively infrequent large 

(M>6) earthquakes (fig. 10), which can be highly destructive, are 

typically produced by island-scale tectonic processes (Heliker, 

1997; Okubo and Nakata, 2011). 

Hawaiʻi residents observed long ago that earthquake 

activity was closely associated with volcanic eruptions. Jaggar 

and HVO seismologists elaborated on this relation and, with 

only a few seismometers, were able to detect the start of 

increased seismicity before earthquakes were large enough 

to feel. They also were able to determine crude locations of 

earthquakes by estimating their distance and azimuth, based on 

seismic traces recorded at HVO, and so be alerted to likely sites 

of impending eruptions. It was not until the 1960s that the HVO 

seismic network had enough instruments to allow triangulation 

of the locations and depths of local earthquake hypocenters. 

160° 158° 156° 154°

20°

22°

M6

M7

Earthquakes 

before 1925

EXPLANATION Water depth, in meters

0

1,000

2,000

3,000

4,000

5,000

6,000

0

0

50 100 KILOMETERS

25 50 MILES

Figure 10. Map showing locations of M6 or stronger earthquakes in the Hawaiian Islands and adjacent ocean areas from 1823 to 2012 (from 

Okubo and Nakata, 2011); color-coded bathymetry is from Eakins and others (2003). The locations of hypocenters for earthquakes that occurred 

before 1925 are not well constrained. 
Natural Hazards and Risk Reduction in Hawai‘i 413

Earthquake Hazards

Destruction during earthquakes is caused by energetic 

shaking of the ground as seismic waves pass through the Earth. 

To understand the impact of future earthquakes, we need to 

know how past earthquakes have affected the Hawaiian Islands. 

Wyss and Koyanagi (1992) produced the first comprehensive 

compilation of shaking effects from large Hawaiian earthquakes 

dating back to 1823. Their map shows maximum Modified 

Mercalli intensities from eyewitness accounts and is an excellent 

first cut at a seismic hazard map (fig. 11). They also estimated the 

probabilities of future damaging earthquakes and of maximum 

Mercalli intensities (table 1). Although published specifically for 

10-, 20-, and 50-year intervals starting in 1990, these probabilities 

are valid for any time intervals of similar length.

VII

VIII VIII

IX

IX

X

XI

156° 155°

20°

19°

A

160° 159° 158° 157° 156° 155°

20°

19°

21°

22°

B

VIII

VII

V

VI

VII

VII

VII

VII

VI

IX

X

0 XI

0

50 100 KILOMETERS

25 50 MILES

0

0

20 40 KILOMETERS

10 20 MILES

Figure 11. Maps showing maximum Mercalli intensities recorded for all earthquakes in the Hawaiian Islands since 1823 (from Wyss 

and Koyanagi, 1992). A, Island of Hawai‘i, where the highest intensities reflect large earthquakes of 1868, 1929, 1973, and 1975. B, All of 

the Hawaiian Islands, where the highest intensities outside the Island of Hawai‘i relate to the earthquakes of 1871 and 1938.

Table 1. Conditional Poisson probability estimates for earthquakes 

in the Hawaiian archipelago and (in parentheses) in the southern 

part of the Island of Hawai‘i for 10-, 20-, and 50-year time periods 

starting in 1990. 

[From Wyss and Koyanagi, 1992, table 9; M, magnitude; I, intensity 

(Modified Mercalli scale); these probabilities can be used for any 10-, 

20-, and 50-year periods]

1990–2000 1990–2010 1990–2040

M≥6 0.84 (0.71) 0.97 (0.92) 0.999 (0.998)

M≥6.5 0.50 (0.39) 0.75 (0.63) 0.97 (0.92)

M≥7 0.17 (0.17) 0.31 (0.31) 0.61 (0.61)

Imax≥VII 0.67 (0.63) 0.89 (0.86) 0.997 (0.99)

Imax≥VIII 0.50 (0.39) 0.75 (0.63) 0.97 (0.92)

Klein and others (2000, 2001) used these data and 

measurements of ground-shaking forces provided by strongmotion instruments (a type of seismometer that measures 

acceleration rather than velocity) to prepare probabilistic 

seismic hazard maps for Hawaiʻi. Figure 12 estimates the peak 

horizontal ground acceleration (in percent g, the acceleration of 

gravity) that could be exceeded with a probability of 10 percent 

within a 50-year interval. The maximum force on a building 

due to earthquake shaking can be estimated as the product of 

the peak ground acceleration times the building mass.

The 2001 seismic hazard maps were based on a single 

Earth model response for strong motion (Klein and others, 

2001); it was assumed that the Earth responds to earthquake 

waves in the same way everywhere on the island. Recently, 

the shear-wave velocity structure beneath each strong-motion 

sensor on the Island of Hawaiʻi was quantified to depths of 

30m or more (Wong and others, 2011), but these new data 

have not yet been incorporated into improved hazard maps.

Mitigation of Seismic Risk: Building Codes and 

Public Education 

The effects of damaging earthquakes can be mitigated 

by requiring all new buildings to be designed to withstand 

expected shaking forces. The State of Hawaii has been 

proactive in passing laws unifying county building codes and 

requiring regular updates. The State adopted the 2006 edition 

of the International Building Code, which categorizes areas 

as zones 1 through 4, with the strongest shaking expected in 

zone4 (including all of Hawaiʻi County). 

In addition to conducting long-term studies of 

earthquakes in Hawaiʻi, HVO has worked with County 
414 Characteristics of Hawaiian Volcanoes

and State agencies in an advisory capacity with regard to 

adopting appropriate building codes and other measures to 

minimize seismic risk. HVO also works to increase public 

awareness of seismic hazards by presenting public lectures 

and meeting with school and civic groups (see Tilling and 

others, this volume, chap. 1). Finally, the articles in Volcano 

Watch—a column written by HVO staff and published 

in Hawaiʻi newspapers and on HVOʼs Web site—offer 

information about seismic hazards and risks (see, for 

example, Hawaiian Volcano Observatory Staff, 2006). 

HVO continues to monitor Hawaiʻi seismicity for the 

USGS and provides a Web site map display of hypocenters 

located in the State in nearly real time (http://hvo.wr.usgs.gov/

seismic/volcweb/earthquakes/), as well as inputs to the USGS 

Earthquake Notification Service (https://sslearthquake.usgs.

gov/ens/register.php). For any Hawaiʻi earthquake of M4.0 or 

greater, HVO quickly posts a press release with earthquake 

parameters and the seismic history of the source area.

The Connection Between Earthquakes and 

Tsunamis

In 1912 many scientists understood that tsunamis were 

caused by large earthquakes, but transoceanic communications 

at that time were insufficient to warn distant communities of 

an advancing wave. On September 29, 1912, HVO recorded its 

first “teleseismic disturbance” related to an earthquake “of great 

power” that occurred at a distance “no less than 5000 miles 

from Hawaii” (Jaggar, 1947, p. 42). By 1922, HVO reports 

tied the recording of a teleseism at about 6 p.m. with the arrival 

of “a succession of pronounced tidal waves on the beach at 

Hilo the next morning.” Knowing that the teleseismic waves 

from large, distant earthquakes travelled much faster than the 

tsunami waves, HVO scientists came up with a way to forecast 

the damaging wave’s arrival: “. . . the transit time of the first 

preliminary [seismic] waves through the earth in minutes and 

seconds is very nearly equal to the transit time of the seismic 

sea waves in hours and minutes. . . .” (Finch, 1924, p. 148). 

Tsunami forecasting based on teleseismic arrivals was born.

The forecasting of tsunamis was not yet fully operational, 

however. An HVO forecast in 1923 was ignored, and the 

tsunami caused $1.5 million in damages and one death. 

Another HVO forecast in 1927 was well heeded, but the 

tsunami failed to materialize. The next HVO tsunami forecast 

was after a Japan earthquake in 1933. This warning was 

taken seriously, and official actions taken in response were 

successful. No lives were lost, and damage was minimal, 

despite a maximum trough-to-crest wave height of 5.3 m at 

some west-facing shores of the Island of Hawaiʻi. 

Yet, on Monday, April 1, 1946, a devastating tsunami 

hit Hilo Bay just before 7 a.m., killing 96 people. A total of 

159 people died throughout the State in the greatest tsunami 

disaster Hawaiʻi has known, to date. Why was no warning 

given in 1946? The answer is surprisingly simple. The distant 

earthquake that heralded the April 1 tsunami was recorded at 

2:06 a.m. Hawaiʻi time but was not noticed until 7:30 a.m., 

when HVO staff reported for work—a half hour after the first 

waves hit Hilo (Hawaiian Volcano Observatory Staff, 2007).

In direct response to the 1946 tsunami disaster, Congress 

funded the creation and operation of a new agency within 

the U.S. Coast and Geodetic Survey called the Seismic Sea 

4.7

13.0

19.2

19.5

24.7

50

113

126

67

36

0 1 2 3 4 5 6 7 8 9 10 15 20 25 30 40 60 80 100

Horizontal acceleration, in percent of gravity

4.7 Point value of peak acceleration, in percent of gravity

19°

20°

21°

22°

160° 159° 158° 157° 156° 155°

0

0

50 100 KILOMETERS

25 50 MILES

Figure 12. Map of the eight 

main Hawaiian Islands, showing 

probability of future earthquake 

ground shaking. Values are peak 

horizontal ground acceleration 

(proportional to force) that has a 

10-percent probability of being 

exceeded in 50 years, expressed 

as a percentage of g, the 

acceleration due to gravity (from 

Okubo and Nakata, 2011, after 

Klein and others, 2000). This map 

can be accessed at the interactive 

USGS Web site http://gldims.

cr.usgs.gov/hishmp/viewer.htm. 
Natural Hazards and Risk Reduction in Hawai‘i 415

Wave Warning System (SSWWS), now called the Richard H. 

Hagemeyer Pacific Tsunami Warning Center (PTWC), located 

on Oʻahu. HVO continued to provide tsunami warnings for 

Hawaiʻi until SSWWS became fully operational in 1949 and 

also reported the arrival parameters of distant earthquakes well 

into the 1960s. After the 1946 disaster, HVO technicians solved 

the problem of how to provide round-the-clock monitoring 

with limited staff. In December 1946, they installed a buzzer 

that went off at the Observatory and in two staff residences 

whenever seismometers detected a distant earthquake large 

enough to generate a tsunami; thus, even in the middle of the 

night, the buzzer system would roust some HVO staff member 

out of bed when a large, distant earthquake occurred. 

The sensor networks now available to the PTWC are 

more spatially extensive than HVO’s and include tide gauges, 

deep-ocean tsunami-detection buoys, and the combined seismic 

instruments of the Global Seismic Network and the USGS 

National Earthquake Information Center. The PTWC area 

of responsibility includes the Pacific Basin. HVO supplied 

earthquake arrival determinations to PTWC during the early 

years, and starting in 1999, selected channels of real-time HVO 

seismic data have been supplied directly to PTWC to aid with 

tsunami forecasting. Since 2011, all real-time seismic data 

acquired by HVO are supplied to PTWC, and vice versa, via 

Internet connections. Efforts are underway to form a single 

entity—the Hawaiʻi Integrated Seismic Network—composed 

of HVO, PTWC, University of Hawaiʻi at Mānoa, and State 

Civil Defense for the purpose of identifying and monitoring 

earthquake and associated tsunami hazards. HVO and the 

USGS will still be responsible for earthquake monitoring, 

while PTWC will be responsible for tsunami warnings.

Other important contributions made by HVO to tsunami 

science includes the rapid documentation and mapping of wave 

effects on Hawaiʻi’s shores. Areas vulnerable to tsunami hazards 

are largely defined by the runup maps from previous tsunamis 

and, more recently, by modeling potential tsunamis. These 

hazards are largest along shallow-sloping coastlines directly 

exposed to the open ocean. HVO and affiliated scientists 

published runup data for the 1960 tsunami (Eaton and others, 

1961), the locally generated tsunami in 1975 (Loomis, 1975; 

Tilling and others, 1976), and the 2011 tsunami generated by the 

Tohoku, Japan, earthquake (Trusdell and others, 2012).

Indirect Volcano Hazards

Eruptive activity in Hawaiʻi is preceded and (or) 

accompanied by earthquakes and ground deformation. In 

general, such phenomena are detectable only instrumentally 

but are sometimes sufficiently energetic to produce effects 

readily discernible by humans (for example, shaking, ground 

subsidence, and opening or widening of ground fractures). 

Mullineaux and others (1987) termed such eruption-related 

processes “indirect volcano hazards,” which are less severe 

and rarely damaging compared to the “direct” or “primary” 

volcano hazards, such as lava flows, tephra fall, pyroclastic 

surges, and volcanic gases. In this section, we discuss these 

indirect processes, as well as the low-probability, highimpact hazard of flank collapse.

Ground Subsidence and Fractures

Ground subsidence poses an indirect hazard on active 

Hawaiian volcanoes in two structural settings: (1) summit 

regions and rift zones, which actively expand and contract as a 

result of subsurface magmatic activity, and (2) volcano flanks, 

which can slip along the contact surface between the volcano 

and the underlying oceanic crust (the décollement fault). 

Such ground motion has occurred frequently on the Island of 

Hawaiʻi during the past two centuries. Moreover, the gradual 

subsidence of the entire Island of Hawaiʻi from volcanic 

loading has been documented to be on the order of several 

millimeters per year (Moore, 1987).

While gradual subsidence is documented primarily by 

geodetic measurements, geological and archaeological evidence 

also exists. For example, in the Kapoho area, where rapid 

subsidence associated with earthquake and eruptive activity 

was reported in 1868, 1924 (Finch, 1925), and 1960 along a rift 

zone graben structure, evidence for long-term subsidence can be 

seen at a popular coastal swimming and snorkeling spot named 

Wai ʻOpae. Here, snorkelers can see a subaerially erupted 

pāhoehoe flow that is now submerged by several meters, with 

only the tops of tumuli remaining above sea level at high tide. 

In addition, snorkelers within Kapoho Bay can observe ancient 

Hawaiian fishpond walls that are now totally submerged. 

Gradual subsidence related to magmatic activity is 

sometimes, but not always, expressed by formation of new 

ground fractures or widening of existing fractures. Such 

fractures commonly ring depressions and craters; the system 

of ring fractures around Kīlauea Caldera, for example, is well 

expressed in the map by de Saint Ours (1982). Subsidence 

associated with magmatic activity can also be abrupt. For 

example, graben-like structures developed (fig. 13), and then 

rapidly subsided, in the Kapoho area a few weeks before 

the 1960 eruption. Likewise, increases in the rate of gradual 

widening of fractures rimming Kīlauea pit craters (for example, 

Halemaʻumaʻu Crater and Mauna Ulu’s summit crater) are well 

correlated with piecemeal collapse of their walls to trigger rock 

falls (see, for instance, Jaggar, 1930a,b; Tilling, 1976). 

Kauahikaua and others (1994) estimated the impact of 

known hazards to proposed geothermal development in the 

lower East Rift Zone and found that subsidence at a rate between 

1.2 and 1.9 cm/yr could be expected. Delaney and others (1998) 

further quantified the net subsidence over the 14-year period 

1976–89 at 60–180 cm within the summit area and ~20 cm along 

the rift zones—a rate for the East Rift Zone that is about the 

same as that of Kauahikaua and others (1994; fig. 14).

Not only can subsidence be an engineering problem for 

structures built within active volcanic structures, it can be 

additionally damaging in areas where rift zones run across 
416 Characteristics of Hawaiian Volcanoes

the coastline. Where the land surface slopes into the ocean, 

subsidence produces an inland incursion of the sea and 

possible submergence of developed areas along the coast. 

The Kapoho coastal area is the only developed location on 

the Island of Hawaiʻi where these problems currently exist 

due to rift-zone subsidence. Brooks and others (2007) found 

subsidence rates of 0.8 to 1.7 cm/yr for this region.

155 ̊ 30' 155 ̊ 00'

19 ̊ 30'

155 ̊ 20' 155 ̊ 10' 154 ̊ 50'

19 ̊ 20'

19 ̊ 10'

Southwest

Rift Zone

East Rift

Zone

Summit

caldera

Island of Hawai‘i

Mauna

Loa Kīlauea

-20

-20

0

20

60

40

-60

-40

-100

-180

?

-40 Contour of vertical

displacement, in centimeters

Kapoho

0

0

5 10 KILOMETERS

3 6 MILES

Faulting and Displacement of Volcano Flanks

Between the December 1965 East Rift Zone eruption and 

the 1975 Kalapana earthquake, Swanson and others (1976) 

documented slow, continuous, seaward movement of the 

whole of Kīlauea’s south flank. This work, using then-modern 

geodetic methods, quantified earlier, less precise measurements 

that suggested such movement of the mobile flank. Swanson 

and others (1976) suggested that this continuous movement 

could result in large, damaging earthquakes by abrupt slippage 

along the décollement fault. In fact, the 1975 Kalapana 

earthquake (Tilling and others, 1976) occurred while the 

Swanson and others (1976) report was in press; a subsequent 

strong décollement earthquake occurred in 1989 (Árnadóttir 

and others, 1991). The large earthquakes in 1823 and 1868 

also almost certainly involved décollement displacements (see, 

for instance, Wyss, 1988). Reoccupation of the trilateration 

network after the 1975 earthquake showed substantial 

horizontal displacements of Kīlauea’s south flank, in places 

as large as 8 m, and several meters of subsidence along the 

coastline (Lipman and others, 1985).

Volcano Flank Collapse

Mobility of volcano flanks can have cataclysmic 

consequences. Since the 1960s, studies have found abundant 

evidence of huge deposits of debris scattered on the seafloor 

surrounding the Hawaiian Islands. The occurrences of such 

massive submarine slumps, landslides, and distal turbiditycurrent flows in the geologic past provide evidence that the 

flanks of Hawaiian volcanoes occasionally become unstable 

Figure 13. Photograph showing earthquake offset of Old Railroad 

Bed Road along the Koa‘e Fault north of the village of Kapoho, prior to 

a Kīlauea eruption. The road is vertically offset about 1.5 m. The Koa‘e 

and nearby Kapoho Faults are old structures, reactivated in April 1924 

during major collapse of the Kapoho graben and again in 1960. These 

faults still exist and will doubtless move again (USGS photograph taken 

January 13, 1960, by J.P. Eaton).

Figure 14. Map 

showing contours of 

vertical displacement on 

Kīlauea Volcano from 

1976 to 1989 based on 

data from leveling, as 

well as from a tide gauge, 

water wells, and surface 

tilt measurements (from 

Delaney and others, 1998, 

fig. 3C). Positive numbers 

mean uplift, negative 

numbers subsidence.
Natural Hazards and Risk Reduction in Hawai‘i 417

and fail catastrophically (for instance, Lipman and others, 1988; 

Moore and Moore, 1988; Moore and others, 1989; Holcomb and 

Robinson, 2004). The most recent of these flank collapses has 

been dated at about 100,000 years ago off the west coast of the 

Island of Hawai‘i. Present-day deposits of broken coral, rock, 

and other sediment near sea level—an elevation that would have 

been several hundred meters higher at the time of the most recent 

collapse—suggest that the collapse generated a “megatsunami” 

(McMurtry and others, 2004). Such powerful tsunami from 

collapse of Hawiian volcano flanks have been hypothesized 

to deposit rocks and sediments as high as several hundreds of 

meters above sea level in other locations in the Hawaiian Island 

chain and even in Australia (see, for instance, Moore and Moore, 

1984; Young and Bryant, 1992). 

Characterizing the Hazard

The hazard zones for ground subsidence and volcano-flank 

faulting were combined in a single map (fig. 15) by Mullineaux 

and others (1987, fig. 22.12). This map serves only to depict the 

areas of the most common, and relatively small-scale, short-term 

hazards posed by displacements associated with frequent effusive 

eruptions and intrusions. While acknowledging that abrupt 

“large-scale” subsidence and faulting would pose more serious 

hazards, such as those accompanying major tectonic earthquakes 

and volcano-flank collapse, Mullineaux and others (1987) made 

no attempt to assess them because of their very low frequency. 

Volcano flank collapse and associated tsunami are an example of 

the classic high-impact, low-probability hazard.

Mitigative Measures

The only effective mitigation strategies available for 

relatively small-scale indirect hazards are to avoid using the areas 

susceptible to such hazards for high-density development or 

critical facilities through land-use zoning and public education. 

There are no practical mitigative measures for large-scale, truly 

catastrophic hazards associated with volcano flank collapse; such 

hazards would impact areas far beyond the maximum hazard zone 

in figure 15. We anticipate, however, that the close, continuous 

volcano monitoring conducted by HVO—particularly of Kīlauea’s 

mobile south flank—should be able to detect any unusual 

acceleration in ground displacements that might suggest departure 

from “normal” volcano behavior. Armed with such information, 

we can provide early warning to emergency-management 

authorities that the flank is becoming highly unstable and that 

immediate work should begin on contingency plans. 

Future Challenges in Reducing Risk

HVO founder Thomas Jaggar, Jr., advocated using the best 

possible scientific information to reduce risks posed by natural 

hazards—effectively using good science to enhance public safety. 

During the past century, HVO, through its long-term monitoring 

studies and topical research, has adhered to Jaggar’s guiding 

principle in responding effectively to hazards that most frequently 

affect Hawaiʻi. While much has been accomplished, much also 

remains to be done. In future decades, HVO and the USGS not 

only must continue to deal with Hawaiʻi’s most common hazards 

(lava flows, volcanic gases, earthquakes, and tsunami), but will 

have an additional, perhaps more demanding, task in addressing 

low-frequency, high-impact hazards that are known to have 

occurred in the geologic past: (1) voluminous explosive eruptions 

capable of producing tephra deposits on a regional scale (fig. 10) 

and (2) collapses of volcano flanks to produce huge submarine 

landslides and attendant local megatsunami. Below we highlight 

some major challenges that HVO, its partners, and collaborating 

scientists will face in augmenting and improving the scientific 

basis to reduce risk from natural hazards in Hawaiʻi.

Improved Characterization and Understanding of 

the Hazards

In perhaps the first-ever, albeit qualitative, publication to 

assess or rank natural hazards in Hawaiʻi, Jaggar (1918, p. 16–17) 

considered four broad categories under what he called “The Index 

of Danger from Volcanoes”: “. . . Class 1, explosion and volcanic 

blast, . . . Class 2, lava flow, . . . Class 3, earthquakes, . . . [and] 

Figure 15. Map depicting the numbered hazard zones 1 to 4 for ground 

fractures and “small-scale” subsidence for the Island of Hawai‘i (from 

Mullineaux and others, 1987, figure 22.12). Also shown are historical lava 

flows (red areas). Subsidence and fracturing events are frequent in zone 

1, which covers the summit and rift zones of Kīlauea and Mauna Loa, but 

are somewhat less frequent in zone 2, on the south flank of Kīlauea. Zone 

3 includes areas around the Kealakekua and Ka‘ōiki Fault Systems on 

Mauna Loa. Zone 4 includes the remainder of the island. 

156° 155°

20°

19°

1

1

1

4

3

3

3

4

4 4

2

2

1

1

0

0

20 40 KILOMETERS

10 20 MILES
418 Characteristics of Hawaiian Volcanoes

Class 4, tidal waves. . . .” As previously discussed, Wentworth 

(1938) compiled the distribution of regional ash deposits from large 

explosive eruptions and an unknown HVO staff scientist produced 

the first, though unpublished, lava flow hazards map for the Island 

of Hawaiʻi, possibly in the 1940s or 1950s. Hazards studies then 

languished until the comprehensive work of Donal Mullineaux 

and his associates (Mullineaux and Peterson, 1974; Mullineaux 

and others, 1987), who used all data available through the 1970s 

to characterize hazards across the Island of Hawai‘i. With much 

geologic (including mapping), geochemical, and geophysical data 

acquired during the past quarter century, we now have improved 

assessments for lava flow, volcanic gas, and earthquake hazards 

(see earlier discussion). The basic procedures for identifying 

and forecasting lava flow paths, long-term gas dispersion, and 

earthquake- and tsunami-prone areas are also now well understood 

and are becoming more quantitative. These are based both on 

documented past occurrences (historical and precontact) and on 

computer models that simulate the hazard mechanisms and impacts.

With continued data collection and scientific/

technological advances in the next 100 years, we can expect 

to obtain refined assessments and precise zonation for 

Hawai‘i’s most frequently occurring volcano hazards (lava 

flows, gas emissions, and earthquakes) within short-tointermediate time scales. At present, however, we lack the 

necessary data to significantly improve the assessments of 

hazards associated with volcano flank collapses or powerful 

explosive eruptions that produce voluminous tephra, as have 

occurred in the geologic past. Essential first steps to refine our 

understanding of both these low-probability but high-impact 

hazards are detailed mapping and dating studies on land and 

offshore. Such studies are necessarily long-term and will 

require considerable commitment of scientific and economic 

resources. For regional tephra falls, dating of the deposits, as 

currently delimited by the Island of Hawaiʻi geologic map 

(Wolfe and Morris, 1996), is inherently difficult, because the 

ash deposits are generally highly altered and (or) reworked. 

For potential volcano flank collapses, detailed studies, 

including denser geodetic monitoring networks, should focus 

on the likely breakaway zones.

In the future, methods to characterize and assess both 

hazards and risks will improve, affording greater precision and 

accuracy or, at the least, a better understanding of the forecast 

uncertainty. It is clear that increasing use will be made of 

probabilistic hazard and risk assessments for long-term effects 

and of event (or logic or decision) trees for active responses 

(see, for example, Newhall and Hoblitt, 2002; Martí and others, 

2008). Seismic hazard maps are already probabilistic (for 

instance, Hanks and Cornell, 1994; http://redirect.conservation.

ca.gov/cgs/rghm/pshamap/), with lava flow hazard maps not far 

behind. An added advantage of probabilistic hazard estimates 

is that one can compare across all hazards and determine 

whether a specific area is more threatened by, for example, 

floods, hurricanes, earthquakes, or tsunamis compared to lava 

flows. Probabilistic methodologies have already become a 

useful cross-hazards language but, to date, have limited use in 

education or emergency-management efforts (Nathan Wood, 

written commun., 2012). Moreover, in the broad discipline of 

risk assessment, there is considerable debate among specialists 

over the merits and disadvantages of probabilistic versus 

possibilistic assessments. For purposes of assessment of natural 

hazards and risk, probabilistic estimates generally refer to 

the most likely events during a given time interval, whereas 

possibilistic estimates apply to worst-case scenarios. Both 

should probably be considered, suggesting that more studies 

are required of worst-case events with high impact but low 

probability (Clark, 2006; Brunsma and Picou, 2008).

We know what the worst-case scenarios are for Hawaiʻi. 

The largest lava flows (for instance, the Panaʻewa flow south of 

Hilo erupted from Mauna Loa) are now mapped and dated but not 

sufficiently studied to understand their eruption dynamics (were 

they rapidly emplaced at a high eruption rate or were they the 

result of prolonged activity at a lower eruption rate?). Widespread 

deposits from explosive eruptions are also mapped but are difficult 

to date or to identify with a source volcano. Earthquakes as 

strong as M8 are possible within the south flank of Mauna Loa 

(Chock and Sgambelluri, 2005), and such earthquakes will almost 

certainly generate local, destructive tsunamis. Much more work 

needs to be done, however, to add to our knowledge of the worst 

possible outcomes of Hawaiian volcano and earthquake activity.

We are well aware that a fundamental, and probably 

unavoidable, deterrent to assigning high priorities to studies of 

low-probability but high-impact hazardous processes involves 

basic human nature. Human beings, including decisionmakers 

and even scientists, quite naturally tend to pay much more 

attention to the most immediate threats posed by natural hazards 

(in other words, those likely to recur within their lifetime) rather 

than to those that have extremely low probability but remain 

possible in the remote future (in other words, those not likely to 

occur within hundreds or even thousands of years).

Updated Volcano Hazards Assessment and Map for 

the Island of Hawai‘i 

The most recent, and still used, volcano hazards 

assessment and map for the Island of Hawai‘i (Mullineaux and 

others, 1987; Wright and others, 1992) do not reflect the data 

and knowledge gained since the early 1980s (especially for lava 

flows). With present geologic and geophysical information, 

along with better tools to convey the hazards posed, improved 

zonation maps for hazards due to lava flows, volcanic-gas 

emissions, and ground fractures and subsidence doubtless 

can be produced. However, any significant advancement in 

the assessment and zonation maps of hazards associated with 

voluminous explosive eruptions (for instance, pyroclastic 

flows and surges, regional ashfalls) will require new data from 

integrated islandwide mapping and laboratory investigations. 

In any case, preparation in the near future of an interim updated 

assessment and map (based on presently available data and 

methodology, even without waiting for the generation of 

additional data) would be of use to all stakeholders, including 

the USGS, the State of Hawaii, and the County of Hawaiʻi. 
Natural Hazards and Risk Reduction in Hawai‘i 419

Enhancements in Communication of Hazard and 

Risk Information

Traditionally, available hazards information has been 

compiled into a map on which the highest hazard zones reflect 

where the hazard is most frequent, most intense, or both, 

while the lowest hazard zones are the least affected. With 

improvements in computer technology, it is now technically 

easier to update hazards assessments and zonation maps of 

various time scales (long-, intermediate-, and short-range). 

Although it lacks the quantitative parameters—geologic as 

well as socioeconomic—needed to evaluate and compute risk, 

this mode of conveying hazards information remains effective 

and will continue to be used into the foreseeable future. 

HVO currently uses a variety of means to communicate 

hazards information but has yet to take full advantage 

of modern media characteristics. The Internet has made 

information available at any time, and effective hazard 

communications can make use of that. Short Message 

Service (SMS) texting and e-mail already provide 

addressable communication methods that are available 

continuously between scientists and emergency managers. 

Social networking modes, such as YouTube, Facebook, 

and Twitter offer many new options once any potential 

security threats have been cleared by the USGS Information 

Technology security policy. Such established social media 

forums could significantly augment the audience to which 

HVO hazard messages are sent. Also, the wide availability 

of Google Maps and Google Earth provide an almost ideal 

medium by which to distribute spatial information. For 

example, the lava flow hazard map for the Island of Hawai‘i 

and information on HVO-located earthquakes are now 

available in Google Earth format. 

However, it should be emphasized that in any effective 

hazard communication, there is also a need for better ways 

to communicate the uncertainties of hazard forecasts and 

assessments. This is a problem across studies of all hazards, 

whether natural or manmade, and much effort in recent decades 

has gone into improving the messages and warnings regarding 

hazards to explicitly incorporate information on uncertainties.

Over the past 20 years, HVO has developed and actively 

used a Web site to disseminate up-to-date information, 

maps, photographs, and videos of volcanic activity, as well 

as publications to deliver more in-depth studies and hazards 

assessments. In addition, HVO submits a weekly article 

on volcanoes and hazards to local newspapers in Hawaii; 

we have now compiled more than 1,000 of these articles, 

which are available to the public on the HVO Web site. HVO 

scientists have published and kept up-to-date several USGS 

Fact Sheets on timely issues like vog, earthquakes, and 

lava flow hazards. We have become even more proactive in 

the past decade, establishing Volcano Awareness Month in 

2009, in which HVO scientists give public talks on current 

volcano and earthquake hazard issues at locations around the 

Island of Hawaiʻi. Starting in 2012, HVO now participates 

in the local Disaster Preparedness Fairs in both Maui and 

Hawaiʻi Counties to distribute information about volcano 

and earthquake hazards. The USGS Volcano Science Center 

also distributes regular hazard assessments using the e-mail 

Volcano Notification Service (VNS), making it possible 

for anyone to keep up with one or all U.S. volcanoes, even 

using just a smart phone.

Concluding Remarks

Volcanoes are always big news in Hawaiʻi, and 

their activities and impacts are followed avidly by island 

residents, government officials, and visitors. Because of 

the frequent—and sometimes continuous (as from 1983 

to present)—eruptive activity, HVO scientists are in 

constant communication with National Park Service and 

Civil Defense officials with regard to ongoing or potential 

hazards. Effective communication is thus conducted on a 

regular basis via the well-established channels described 

by Tilling and others (this volume, chap. 1). Of course, 

there is sometimes also a need for special or urgent 

communications under crisis conditions, as is the case at 

volcanoes elsewhere in the world that erupt infrequently or 

unexpectedly.

In its first century monitoring Hawaiian volcanoes, 

HVO has communicated hazards information by making 

it accessible in an efficient and timely manner, using the 

best available methods and technology. There is every 

reason to believe that this century-long legacy will carry 

into the next 100 years, with new methods focused on 

pushing hazards information to individuals as simply yet 

comprehensively as possible.

Acknowledgments

As past Scientists-in-Charge of HVO, we are well 

aware and deeply appreciative of the dedication and 

hard work of the HVO staff in conducting the long-term 

monitoring and research needed to understand Hawaii’s 

natural hazards and the risks they pose. To the present and 

past staff, we owe a huge debt of thanks. Our colleagues 

Nathan Wood and Cynthia Gardner (both at the Cascades 

Volcano Observatory) critically reviewed an earlier draft 

of this chapter. Their reviews were thorough, provocative, 

and incisive, prompting us to sharpen the presentation and 

interpretation of the discussion. Nate and Cynthia, mahalo 

nui loa for your thoughtful and helpful reviews! We also 

thank Mike Poland, editor of this collection of papers, for 

his persistent and gentle support throughout the writing, 

reviewing, and preliminary layout phases. Finally, we 

thank our wives, Jeri and Susan, for their patience and 

support throughout this project.
420 Characteristics of Hawaiian Volcanoes

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Árnadóttir, T., Segall, P., and Delaney, P., 1991, A fault model 

for the 1989 Kilauea south flank earthquake from leveling 

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CWEOTS>2.3.CO;2.
Visual (top) and thermal (bottom) images of a lava flow as it entered Pāhoa Village in the Puna District of the 

Island of Hawai‘i. In the thermal images, warmer colors indicate higher temperatures. USGS photograph by T.R. 

Orr and thermal image by M.R. Patrick, November 14, 2014.
Produced in the Menlo Park Publishing Service Center, California
 “The writer, from his experience of a decade of recording on an active volcano, is convinced that all that he learned in seven 

volcano expeditions to distant lands was as nothing compared with a few years in a fixed experiment station. He never had 

the slightest suspicion, from the travellers’ accounts, of the number of changes which occur in a short time. . . . New problems 

open out endlessly and new experiments are suggested. The old experiments get new meanings and need repetition with more 

precise method. Every day of observation opens up a new query.” 

Thomas A. Jaggar, Jr., founder of the Hawaiian Volcano Observatory, looks out on the summit caldera of Kīlauea Volcano from the front porch of the first 

site of HVO, on the west rim of the crater. USGS photograph, 1916 (photographer unknown). 

Poland and others—Characteristics of Hawaiian Volcanoes—Professional Paper 1801

ISSN 1044-9612 (print) 

ISSN 2330-7102 (online)

http://dx.doi.org/10.3133/pp1801

—Thomas A. Jaggar, Jr., in “A plea for geophysical 

and geochemical observatories,” published in the 

Journal of the Washington Academy of Sciences, 1922