Preliminary Assessment of Volcanic and Hydrothermal

Hazards in Yellowstone National Park and Vicinity

By Robert L. Christiansen, Jacob B. Lowenstern, Robert B. Smith, Henry Heasler, Lisa A. Morgan, Manuel

Nathenson, Larry G. Mastin, L. J. Patrick Muffler, and Joel E. Robinson

Open-file Report 2007–1071

U.S. Department of the Interior

U.S. Geological Survey
ii

U.S. Department of the Interior

DIRK KEMPTHORNE, Secretary

U.S. Geological Survey

Mark D. Myers, Director

U.S. Geological Survey, Reston, Virginia 2007

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Suggested citation:

Christiansen, R. L., Lowenstern, J. B., Smith, R. B., Heasler, H., Morgan, L. A., Nathenson, M., Mastin, L. G.,

Muffler, L. J. P., and Robinson, J. E., 2007, Preliminary assessment of volcanic and hydrothermal hazards in

Yellowstone National Park and vicinity: U.S. Geological Survey Open-file Report 2007-1071, 94 p.

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

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iii

Contents

Contents................................................................................................................................................................iii

Brief Summary.......................................................................................................................................................1

Introduction ...........................................................................................................................................................4

Factors considered in this assessment................................................................................................................5

Current State of the Yellowstone system.............................................................................................................5

Geologic background ........................................................................................................................................5

Contemporary activity........................................................................................................................................8

Seismicity .......................................................................................................................................................8

Crustal deformation...................................................................................................................................... 11

Hydrothermal and gas activity..................................................................................................................... 12

Pressures, temperatures, and fluids in geothermal systems ................................................................. 13

Mechanisms of hydrothermal explosion ................................................................................................. 14

Hydrothermal explosions in Yellowstone ................................................................................................ 15

Factors contributing to hydrothermal explosions ................................................................................... 16

Toxic gases ............................................................................................................................................... 17

The Hazards......................................................................................................................................................... 17

Volcanic-eruption hazards .............................................................................................................................. 18

Basaltic eruptions ........................................................................................................................................ 18

Large rhyolitic eruptions .............................................................................................................................. 20

Small rhyolitic eruptions .............................................................................................................................. 25

Large caldera-forming eruption................................................................................................................... 26

Hydrothermal-explosion hazards.................................................................................................................... 29

How often do they occur?............................................................................................................................ 29

Potential effects ........................................................................................................................................... 30

Precursory signals ....................................................................................................................................... 31

Where are hydrothermal explosions most likely to occur?........................................................................ 31

Seasonal and long-term effects on hydrothermal explosions ................................................................... 33

Hazard mitigation ......................................................................................................................................... 33

Gas-emission hazards ..................................................................................................................................... 34

Relevant examples of toxic volcanic or hydrothermal gas hazards .......................................................... 35

Hazard mitigation ......................................................................................................................................... 36

Conclusions ......................................................................................................................................................... 36

Acknowledgments............................................................................................................................................... 37

References Cited................................................................................................................................................. 37

Appendix 1. Description of representative historic hydrothermal explosions ................................................. 86

Porkchop Spring/Geyser ................................................................................................................................. 86

Excelsior Geyser.............................................................................................................................................. 86

West Nymph Creek Thermal Area .................................................................................................................. 87

Black Opal/Wall Pool and Sapphire Pool ....................................................................................................... 87

Historic hydrothermal explosions elsewhere................................................................................................. 87

Appendix 2. Description of large prehistoric hydrothermal eruption sites at Yellowstone ............................. 88

Pocket Basin .................................................................................................................................................... 88

Mary Bay.......................................................................................................................................................... 88

Elliott’s crater................................................................................................................................................... 89

Evil Twin explosion crater ............................................................................................................................... 89

Frank Island explosion crater.......................................................................................................................... 89

Indian Pond ...................................................................................................................................................... 89

Turbid Lake....................................................................................................................................................... 90
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Appendix 3. Probabilities of episodic volcanic eruptions and application to the young intracaldera volcanic

history of Yellowstone......................................................................................................................................... 91
1

Preliminary Assessment of Volcanic and

Hydrothermal Hazards in Yellowstone National Park

and Vicinity

By Robert L. Christiansen1, Jacob B. Lowenstern1, Robert B. Smith2, Henry Heasler3, Lisa A.

Morgan1, Manuel Nathenson1, Larry G. Mastin1, L. J. Patrick Muffler1, and Joel E. Robinson1

Brief Summary

Possible future violent events in the active hydrothermal, magmatic, and tectonic system of

Yellowstone National Park pose potential hazards to park visitors and infrastructure. Most of the

national park and vicinity are sparsely populated, but significant numbers of people as well as park

resources could nevertheless be at risk from these hazards. Depending on the nature and magnitude

of a particular hazardous event and the particular time and season when it might occur, 70,000 to

more than 100,000 persons could be affected; the most violent events could affect a broader region

or even continent-wide areas. This assessment of such hazards is presented both as a guide for

future activities of the Yellowstone Volcano Observatory (YVO) and to aid appropriate response

planning by the National Park Service and surrounding agencies and communities. Although the

assessment is presented here in some technical detail, this summary is intended to be

understandable to non-scientists. The principal conclusions also will be made available in other

forms, more accessible to general readers.

The Yellowstone Plateau was built by one of Earth's largest young volcanic systems, having

episodically erupted great volumes of both lava and explosively ejected pumiceous ash for more

than 2 million years. These eruptive materials are products of two compositional types of

subsurface magma: basaltic magma is relatively fluid and, in this setting, generally produces small

to moderate volumes of lava in relatively brief eruptions; rhyolitic magma is more viscous and

either can erupt effusively to produce small to large volumes of lava or explosively to produce

coarse pumice and finer ash. The three largest Yellowstone eruptions produced blanketing deposits

of rhyolitic ash so hot that they welded into sheets of dense rock covering large areas, extending

beyond the national park. Each of them also produced a rain of ash that spread over much of

western and central North America and beyond; these ash deposits are greater than 2 m thick near

their eruptive sources and as much as a meter thick in surrounding areas. Each of these three

eruptions produced a caldera, or deep crater-like depression, tens of kilometers wide, formed by

collapse of the ground surface into a partly emptied subterranean magma chamber. The latest of

these three great eruptions formed the Yellowstone caldera. Renewed rhyolitic magma influx

beneath the Yellowstone caldera in central Yellowstone National Park uplifted parts of the caldera

floor and produced voluminous intracaldera lavas, the youngest of which extruded in a series of

eruptive episodes about 164,000, 152,000, 114,000, 102,000, and 72,000 years ago. During the

same span of time, generally smaller flows of both basalt and rhyolite have erupted in several areas

outside the Yellowstone caldera: (1) northwest of the caldera, (2) near the southern boundary of the

 

1

 U.S. Geological Survey

2

 University of Utah

3

 National Park Service
2

park, (3) in the basin of Island Park, west of Yellowstone National Park, and (4) especially in the

southern part of a corridor between Norris Geyser Basin and Mammoth Hot Springs.

Disruption of the Earth’s surface by faulting and regional uplift characterize the geologic

framework of Yellowstone Plateau volcanism. Some of the regional faults that bound the mountain

ranges around the Yellowstone Plateau are capable of producing large-magnitude (M>7)

earthquakes. In contrast, faults within the caldera are mainly small, produce smaller-magnitude

(M6.8), relatively shallow earthquakes, and reflect strains in the Earth’s crust associated with

magmatic intrusion and hydrothermal activity. Swarms of generally small earthquakes occurring

within localized areas over restricted periods of time characterize much of the earthquake activity

within and adjacent to the Yellowstone caldera. This seismicity is monitored by a network of

seismographs within and adjacent to the park and is recorded and processed nearly in real time at

the University of Utah as part of the YVO program and archived as a contribution to the U.S.

Geological Survey (USGS) Advanced National Seismic System database.

Leveling surveys, satellite-based measurements, and geologic studies of former shorelines

of Yellowstone Lake all show that the entire area of the Yellowstone caldera and a seismically

active zone to the northwest undergo episodes of ground uplift and subsidence, sometimes

encompassing the entire caldera and sometimes in more local and complex patterns of both uplift

and subsidence. Such deformation in the Yellowstone region is monitored by YVO mainly through

a network of continuously recorded Global Positioning System (GPS) receivers recorded at the

University of Utah. The GPS data are incorporated as part of the National Science Foundation’s

EarthScope Plate Boundary Observatory, archived at UNAVCO and available through the YVO

website (http://volcanoes.usgs.gov/yvo/index.html).

The active hydrothermal system of Yellowstone National Park is one of the largest on

Earth. Although accidents involving hot water injure Yellowstone visitors from time to time,

conformance with normal Park Service procedures and regulations would ordinarily be sufficient to

prevent most of them. By contrast, a commonly recurring, more acute hazard at Yellowstone is the

explosive ejection of steam, water, and rock with no associated volcanism. These hydrothermal

explosions are caused by hot subsurface waters that flash to steam, breaking the overlying rocks

that confine them and ejecting the debris to form a crater. It is generally not clear just what triggers

these events, but possible triggers include strong local earthquakes, seasonal or long-term declines

in ground water levels, and changes in the underground distribution of heat. Many hydrothermal

explosions have few if any premonitory indications.

At least 26 hydrothermal explosions have been documented in the 126-year historic record

of the national park, and others undoubtedly escaped observation. Since the Yellowstone Plateau

was last glaciated, ending about 16,000 years ago, at least 18 large hydrothermal explosions have

formed craters wider than 100 m. Conservatively, at least one rock-hurling explosion every two

years is estimated to occur at Yellowstone, but because most of these events are small and usually

occur when few visitors are present, the likelihood of harm to park visitors is relatively small. The

average recurrence of an explosion that could cause personal injury is probably between 10 and 100

years. Average recurrence time of an explosion large enough to produce a 100-m-diameter crater is

probably about 200 years, but such an event could expel rocks and other hot debris more than 2 km

from the explosion site. Most hydrothermal-explosion craters at Yellowstone are in the Firehole

River geyser basins, beneath or around Yellowstone Lake, and in the southern part of the NorrisMammoth Corridor.

In addition to hydrothermal explosions, toxic gases—mainly carbon dioxide and hydrogen

sulfide—pose hazards. Concentrations of these gases in the atmosphere are generally low at

Yellowstone, but they can build up in confined areas such as valleys, caves, and tunnels, especially

during windless conditions. Most areas with toxic-gas hazards can be kept off-limits to people, but

gas emissions should continue to be monitored.
3

No volcanic eruption has occurred in Yellowstone National Park or vicinity in the last

70,000 years or more. Nevertheless, several types of volcanic eruption hazards are possible in the

future.

Basaltic lavas have erupted around the margins of the Yellowstone Plateau volcanic field

throughout its evolution. These relatively low-viscosity lavas generally erupt rapidly, most

eruptions lasting no more than a few weeks to a few months though the largest flow fields may

accumulate in multiple eruptions lasting months to years. The average period between basaltic

eruptions in the Yellowstone region since formation of the Yellowstone caldera has been about

16,000 years. The most likely location of a future basaltic eruption is within the basin of Island

Park, west of Yellowstone, but basalts could erupt anywhere in a 40-km-wide band around the

caldera. Future basaltic eruptions could cover several square kilometers with lava up to tens of

meters thick. Basaltic ash and cinders also might blanket hundreds of square kilometers to depths

of a few meters to a few centimeters, and if a vent emerged beneath water or saturated ground,

more explosive eruptions could cause significant destruction, such as blasting down trees or

structures.

Large rhyolitic lava flows, many having volumes greater than 10 km3, have erupted within

the Yellowstone caldera during the past 170,000 years. Initially these larger eruptions were

preceded by explosively ejected pumice and ash. In a similar future eruption, ejecta could bury

broad areas, locally to many meters. Subsequent lava extrusion could last for years, covering areas

as great as 350-400 km2

 to thicknesses of tens or hundreds of meters and volumes of 5 to more than

50 km3. Because such voluminous rhyolitic lava eruptions have not been observed anywhere in

historical time, it is uncertain how long such an event might continue; extrusion might be orders of

magnitude faster than for smaller flows. The probability of a future large intracaldera rhyolitic

eruption is difficult to estimate. Available data suggest a highly episodic behavior of past eruptions

of this sort, periods of a few thousand years characterized by numerous eruptions being separated

by longer intervals of about 12,000 to 38,000 years without eruption. One statistical measure of

eruption probabilities based on this episodic behavior suggests an average recurrence of 20,000

years. The fact that no such eruption has occurred for more than 70,000 years may mean that

insufficient eruptible magma remains beneath the Yellowstone caldera to produce another largevolume lava flow.

Small rhyolitic lava flows postdating the Yellowstone caldera have erupted mainly north of

the caldera, but one such flow also lies near the South Entrance to the park. Two distinct types of

primary hazards might be associated with small rhyolitic eruptions at Yellowstone. Just as for

larger rhyolitic lava eruptions, initial venting almost certainly would explosively eject rhyolitic

pumice; the coarser fragments would fall back close to the vent, but finer pumiceous ash would

enter the atmosphere and fall downwind for many kilometers. Structures, power lines, etc. could be

damaged by ash loading, especially if eruption were accompanied by heavy rain. The initial

explosive eruptions could last a few hours to several weeks and be followed by viscous extrusion of

rhyolitic lava, covering several square kilometers to tens of meters thick; lava could continue to

extrude for many months or even years. Viscous rhyolitic lava would advance much more slowly

than a basaltic flow; most affected facilities could be safely evacuated and perhaps relocated. The

average recurrence period of small extracaldera rhyolitic eruptions in the Yellowstone Plateau

volcanic field is about 50,000 years.

In addition to the primary hazards posed by any future eruption of basalt or a small or large

rhyolitic lava eruption, important possible secondary consequences include wildland fires, debris

flows, and floods triggered by the displacement of surface drainages by lava.

Systematic seismic, deformation, and hydrothermal monitoring by YVO is likely to provide

indicators of any impending volcanic eruptive activity in Yellowstone National Park. Premonitory

events detected by such monitoring might include multiple shallow earthquake swarms of
4

increasing frequency and intensity, the ground vibrations called volcanic tremor, localized uplift of

the surface, ground cracks, and anomalous gas emissions.

Of all the possible hazards from a future volcanic eruption in the Yellowstone region, by far

the least likely would be another explosive caldera-forming eruption of great volumes of rhyolitic

ash. Abundant evidence indicates that hot magma continues to exist beneath Yellowstone, but it is

uncertain how much of it remains liquid, how well the liquid is interconnected, and thus how much

remains eruptible. Any eruption of sufficient volume to form a new caldera probably would occur

only from within the present Yellowstone caldera, and the history of postcaldera rhyolitic eruptions

strongly suggests that the subcaldera magma chamber is now a largely crystallized mush. The

probability of another major caldera-forming Yellowstone eruption, in the absence of strong

premonitory indications of major magmatic intrusion and degassing beneath a large area of the

caldera, can be considered to be below the threshold of useful calculation.

Introduction

Yellowstone National Park, justly famous for its unmatched geysers, diverse wildlife, and

uniquely preserved ecologic communities, also encompasses one of Earth’s largest systems of

volcanic, seismic, and hydrothermal activity. In recognition of the importance of this active Earth

system, officials of the U.S. Geological Survey (USGS), Yellowstone National Park, and the

University of Utah in May of 2001 jointly established the Yellowstone Volcano Observatory

(YVO). The stated objectives of this new observatory are: (1) to provide monitoring that enables

reliable and timely warnings of possible renewed volcanism and related hazards in the Yellowstone

region, (2) to notify National Park and other local officials and the public of any significant local

seismic or volcanic events, (3) to improve scientific understanding of the fundamental tectonic and

magmatic processes that create the Park’s ongoing seismicity, surface deformation, and

hydrothermal activity, (4) to assess the long-term potential hazards that volcanism, seismicity, and

explosive or other convulsive hydrothermal activity might pose to the park and its surroundings, (5)

to communicate as effectively as possible to responsible authorities and the public the results of

scientific monitoring and hazard-assessment activities, and (6) to improve coordination and

cooperation among the three institutions responsible for YVO. The observatory is built upon a

substantial base of previous cooperative work among these institutions and seeks to assure a solid

long-term basis for the continuity and improvement of scientific monitoring of the Yellowstone

magmatic-tectonic-hydrothermal system.

The hazard assessment presented in this report is intended to help guide future activities of

the observatory as well as to provide a basis for appropriate management actions by the National

Park Service and other agencies in the Yellowstone area in the event of any future hazardous events

that might result from activity of the Yellowstone system. The assessment is part of an ongoing

three-part process that was set in motion at the outset of YVO’s work. The first part of the process

was a comprehensive review of basic scientific knowledge of Yellowstone’s magmatic-tectonichydrothermal system. Such a review, of course, is never final and must be continually reexamined

and revised, particularly in the light of current monitoring data. The second step of the process is

assessment of the relevant hazards. The current hazards assessment is intended to be fully

documented and scientifically defensible; because this necessarily entails a degree of scientific

rigor and technical documentation not readily understandable by general readers or to all concerned

individuals, additional reports more suitable for non-scientist readers also are part of the ongoing

assessment activity. One general-interest publication on Yellowstone’s volcano, seismic, and

hydrothermal hazards has already been published (Lowenstern and others, 2005). The final step of

the process will be a response plan, based upon the conclusions of this assessment, by the

responsible Yellowstone National Park officials, in cooperation with other appropriate agencies and

with the assistance of observatory scientists.
5

It is worth noting here that, despite the fact that most of Yellowstone National Park and its

nearby surroundings are sparsely populated, many thousands of people as well as National Park

resources could be at risk from the types of hazards considered in this assessment. About 3 million

people visit Yellowstone each year, principally during the three summer months. A similar number

of people visit adjacent Grand Teton National Park (although some of these are the same visitors).

Including residents of the surrounding communities of Wyoming, Idaho, and Montana, the daily

population exposure to these hazards during the summer months could average between 70,000 and

more than 100,000 persons. This exposure produces a largely unappreciated level of risk,

comparable to that of other areas having considerably larger resident populations. Furthermore, the

largest of these hazards, although they have the lowest probabilities of occurrence, could affect

much of western and central North America. Indirect effects, especially on climate, could be

global.

Factors considered in this assessment

In order to clarify the intended scope of this report, it is important to state at the outset just

what factors are considered explicitly in this assessment. Hazards that might result from ongoing

or future activity of the Yellowstone magmatic-tectonic-hydrothermal system, particularly volcanic

eruptions, earthquakes, and hydrothermal eruptions, are the focus of this discussion. Certain other

geologic hazards, although likely to be of important concern to National Park Service managers in

the future, are not included within its scope. Examples of the latter would include landslides,

debris flows, or floods, except insofar as they are considered here as possible consequential

secondary hazards that might result from major volcanic, seismic, or hydrothermal activity.

The area considered in this hazard assessment is primarily Yellowstone National Park (fig.

1). Nevertheless, because most of the hazards considered here could have significant effects on

adjacent communities, the Park boundary does not constitute an absolute limit in our analysis.

The strategy is to establish an integrated view of the sources of relevant hazards and some

scenarios for how they might develop in time. This approach diverges from that of many previous

USGS volcano-hazard assessments in not emphasizing a catalogue of individual hazardous

processes and zonation maps for each of them. Rather, the emphasis is on (1) what kinds of

monitoring data might be of immediate concern to National Park Service managers or might be

considered premonitory to hazardous events, (2) the probabilities of recurring volcanic and acute

hydrothermal events, and (3) how multiple events might be related to one another. Although the

present open-file version of this report does not analyze earthquake hazards explicitly, a

forthcoming revision for more formal publication will include seismic hazards.

Current State of the Yellowstone system

Assessment of possible future activity in Yellowstone starts from an analysis of the current

state of the magmatic-tectonic-hydrothermal system. We first review the geologic framework and

follow with information on current activity and monitoring of the system.

Geologic background

The Yellowstone Plateau volcanic field of Wyoming, Idaho, and Montana (fig. 2) is one of

Earth's largest young volcanic systems, having erupted extraordinarily voluminous rhyolites

episodically over a little more than 2 million years (Christiansen, 1984; 2001). Its three largest

eruptions deposited sheets of mainly welded ash-flow tuffs regionally as well as coeval ash-fall

layers that fell over much of western and central North America (fig. 3). The volcanic field is the

current expression of a major sublithospheric mantle source that generates a melting anomaly that
6

has propagated northeastward at least 300 km relative to the North America craton since its initial

volcanism at 17-14 Ma (fig. 4). Researchers debate the nature of this melting source—whether, for

example, it represents a convective thermal plume from the base of the mantle or from the mantle

transition zone, an entirely upper-mantle response to plate-tectonic processes, or some other

mechanism (e.g., Pierce and Morgan, 1992; Smith and Braile, 1993, 1994; Pierce and others, 2000;

Christiansen and others, 2002; Camp and Ross, 2004; Waite and others, 2006). These varied

hypotheses are significant to interpretation of regional geophysical data but have only an indirect

bearing on the present analysis of volcanic hazards.

The volcanic field has evolved episodically in three cycles of rhyolitic activity. Each cycle

culminated in the rapid eruption of voluminous rhyolitic ash flows—hundreds to thousands of

cubic kilometers—and consequent catastrophic subsidence of the source areas to form large

calderas. Each climactic ash-flow eruption was preceded by a period of magmatic intrusion and

intermittent rhyolitic lava eruptions and was followed by a period of partial filling of the calderas

with rhyolitic lavas. During each of these cycles, basaltic lavas erupted around the margins of the

active rhyolitic source area but not within it. About a million years after their rhyolitic activity,

basalts finally erupted through the source areas of the first two cycles, but no basalts have yet

erupted within the youngest, the Yellowstone caldera.

Collectively, the voluminous rhyolitic ash-flow tuffs of the volcanic field are known

stratigraphically as the Yellowstone Group (Christiansen and Blank, 1972). The oldest and largest

of the caldera-forming eruptions produced the Huckleberry Ridge Tuff at 2.059±0.004 Ma

(Lanphere and others, 2002), covering an area of more than 15,000 km2 with ash flows having a

cumulative volume of at least 2,450 km3. The resulting caldera (fig. 5, purple line) spanned nearly

the entire width of the Yellowstone Plateau volcanic field (Christiansen, 1979, 2001). The Mesa

Falls Tuff of 1.285±0.004 Ma (Lanphere and others, 2002), smallest of the three ash-flow sheets,

erupted in the Island Park area west of the Yellowstone Plateau (Christiansen, 1982, 2001) and is

now exposed mainly in that vicinity, adjacent to its source caldera (fig. 5, blue line). The Mesa

Falls probably initially covered an area of more than 2,700 km2 and had an eruptive volume of

more than 280 km3. The youngest of the major caldera-forming eruptions at 0.639±0.002 Ma

(Lanphere and others, 2002) produced the Lava Creek Tuff and the Yellowstone caldera, now the

central feature of the volcanic field (Christiansen, 1984). Two distinct parts of the Lava Creek Tuff

erupted from separate segments of the caldera (Christiansen, 1979, 2001) but form a single

compound cooling unit that initially covered at least 7,300 km2 and had an eruptive volume of at

least 1,000 km3.

Precaldera rhyolitic lavas of the third volcanic cycle erupted from a growing ring-fracture

system in the area that subsequently became the Yellowstone caldera. Eight such precaldera

rhyolite flows are known, yielding isotopic ages from 1.22±0.01 to 0.609±0.006 Ma (table 1; see

note on the table regarding the accuracy of cited ages). Additional precaldera rhyolitic lava flows

may now lie entirely buried within the Yellowstone caldera. The precaldera lavas, named the

Mount Jackson Rhyolite and Lewis Canyon Rhyolite, and the growing ring-fracture system from

which they extruded (fig. 6) are interpreted as indicating magmatic intrusion and the growth of a

large rhyolitic magma chamber in the shallow crust during a period of nearly 600,000 years beneath

the area that would later erupt catastrophically to produce the Lava Creek Tuff and subside to form

the Yellowstone caldera (fig. 7).

Shortly following the climactic caldera-forming Lava Creek eruption, rhyolitic magma

again intruded the subcaldera region, uplifting segments of the caldera floor in a pair of resurgent

domes bounded by two inner ring-fracture segments enclosed by an outer ring-fracture zone

encompassing nearly the entire Yellowstone caldera (Christiansen, 2001). The western of these

resurgent domes is the Mallard Lake dome (fig. 7, ML); the eastern is the Sour Creek dome (fig. 7,

SC). Subsequently, additional rhyolitic lavas have erupted within the caldera during several
7

volcanic episodes. These rhyolites, mainly constituting large lava flows, are included

stratigraphically in the Plateau Rhyolite. The oldest of the Plateau Rhyolite flows, the Upper Basin

Member, erupted within the caldera from around the inner ring-fracture zones that bound the two

resurgent domes (fig. 7). These oldest postcaldera rhyolites yield isotopic ages that range between

516±7 and 479±10 ka (Gansecki and others, 1996) though there is a reasonable possibility that their

actual eruptive ages were even closer to that of the ~639-ka Lava Creek Tuff (see Lanphere and

others, 2002). Buried rhyolitic lavas identified in recent high-resolution sonar and seismicreflection mapping of Yellowstone Lake (Morgan and others, 2003b) and found as lithic clasts in

postglacial hydrothermal-explosion deposits south of the Sour Creek dome and in northern

Yellowstone Lake yield a 40Ar/39Ar age of 600±20 ka (Morgan and Shanks, 2005). Additional

postcaldera lava, probably as young as 275±11 ka is known (table 1). Other lavas may well have

extruded within the caldera during the time before about 170 ka but remain buried beneath younger

lavas and sediments.

Younger postcaldera rhyolitic volcanism dates from about 170 ka (Christiansen, 2001).

This activity probably began with extrusion of rhyolitic lava onto the Mallard Lake resurgent dome

in the western part of the Yellowstone caldera (the Mallard Lake flow) followed by renewed uplift

of that dome. Following those events, as much as several hundred cubic kilometers of rhyolitic

lava (figs. 8 and 9), named the Central Plateau Member of the Plateau Rhyolite, has nearly filled

the Yellowstone caldera from eruptive vents along two linear northwest-trending zones that project

across the caldera from neighboring tectonic fault zones. The western of these two zones extends

from the Teton fault zone (see fig. 11) to the tectonic basin of West Yellowstone. Lavas from these

vents form the Pitchstone and Madison Plateaus (fig. 1) and bury the western rim of the

Yellowstone caldera. The eastern zone extends northward across the caldera from the Sheridan

fault zone (see fig. 11) to the Norris Geyser Basin and erupted the lavas that form the Central

Plateau. The lava flows from this youngest intracaldera activity are typically quite large—some of

them exceeding 20 km3—and appear to have erupted in five episodes at about 164±5, 152±3,

114±2, 102±5, and 72±4 ka (table 2 and figs. 8 and 9). It is likely that rhyolitic pumice and ash

erupted during the opening of vents for each of these lava flows. In addition, concurrent with lava

eruptions of the Central Plateau Member, at least two significant pyroclastic eruptions occurred

within the caldera, depositing the tuffs of Bluff Point and Cold Mountain Creek. The former was

of sufficiently large volume to have caused subsidence of its source area to form a smaller caldera

(~10-km diameter) within the Yellowstone caldera, now partly preserved as West Thumb (fig. 8),

the western part of Yellowstone Lake (Christiansen, 2001). It is worth noting that this relatively

small caldera is itself as large as the well-known caldera of Crater Lake, Ore., which formed as a

result of the pyroclastic eruption of more than 50 km3

 of magma (Bacon, 1983).

Volcanism postdating the Yellowstone caldera also has occurred in areas outside the caldera

but within or near the boundaries of Yellowstone National Park. The individual eruptive products

of this extracaldera volcanism are generally smaller in volume than most of those within the caldera

but span nearly the same range of time as the intracaldera lavas. Both rhyolites and basalts occur

among these extracaldera lavas. K/Ar and 40Ar/39Ar ages obtained on the rhyolites range between

526±3 and 80±2 ka (table 1). Basaltic eruptions appear to have spanned much of the same time.

Basalts and rhyolites both erupted mainly in the southern part of the Norris-Mammoth corridor, a

zone of faulting, volcanism, and hydrothermal activity that extends northward from the caldera

margin near Norris Geyser Basin to just north of Gardiner, Mont. (figs. 10, 11). Other rhyolites and

basalts of similar ages have erupted northwest of the caldera, between the Madison River and

Cougar Creek, and near the southern boundary of Yellowstone National Park. Furthermore, basalts

erupted from at least 17 vents in the basin of Island Park, within 35 km of the west boundary of

Yellowstone National Park (fig. 10).
8

Hydrothermal activity occurs widely on and around the Yellowstone Plateau. Most of the

hottest and most active areas of geysers and other near-neutral (i.e., non-acidic) hot springs occur

within topographic basins situated above the ring-fracture zone of the Yellowstone

caldera—including Yellowstone Lake (Johnson and others, 2003; Morgan and others, 2003a;

Morgan and others, 2003b)—but others occur in the Norris-Mammoth corridor north of the caldera.

Other major sites of mainly acidic, sulfate-rich hot-springs and fumaroles occur in topographically

higher areas in and adjacent to the ring-fracture zone and the northeastern caldera rim. Smaller,

less concentrated zones of hydrothermal activity are distributed widely around the Yellowstone

National Park, but most areas deemed capable of producing violent events such as hydrothermal

explosions are in the major hydrothermal areas within or adjacent to the caldera and in the NorrisMammoth corridor.

Regional uplift and normal faulting have established the geologic framework within which

the Yellowstone Plateau volcanic field lies (figs. 4 and 11). Faults mapped within the caldera are

mainly small and seem to reflect strains associated with magmatic intrusion and, perhaps, shallow

hydrothermal activity. The principal faults south of the caldera that accommodate regional tectonic

extension generally trend nearly north-south and define the Teton, Sheridan, Flat Mountain, and

Upper Yellowstone fault zones. Major fault zones north of the caldera include the Lamar, East

Gallatin-Washburn, Hebgen, and Madison Valley fault zones, many of them trending more nearly

northwest. A little farther west, the Centennial fault zone trends east-west. The somewhat arcuate

Mirror Plateau fault zone northeast of the caldera (figs. 6 and 11) appears to have accommodated

both regional tectonic extension and displacements associated directly with the Yellowstone

caldera system (Christiansen, 2001).

Bathymetric and seismic-reflection studies have delineated several small faults beneath

Yellowstone Lake (Otis and others, 1977; Wold and others, 1977; Johnson and others, 2003;

Morgan and others, 2003b). At least the major sublacustrine faults appear to represent

continuations of the regional tectonic trends noted above (Morgan and others, in press-b).

Thermal surveys of Yellowstone Lake (Morgan and others, 1977), together with

geochemical indicators of the heat output of the Yellowstone hydrothermal system (Fournier,

1989), demonstrate that the total heat flux from the Yellowstone caldera exceeds 1500 mW/m2

,

more than thirty times the regional average. This great thermal flux is a direct reflection of the

magmatic heat source that produces and sustains Yellowstone’s hydrothermal activity.

Contemporary activity

The tectonic-volcanic-hydrothermal system of the Yellowstone region is vigorously active.

The hydrothermal system may be the largest on Earth, and the subsurface magmatic system

continually deforms the ground surface on time scales of months to a few years. Regional tectonics

and the magmatic system combine to produce some of the highest levels of earthquake activity in

the conterminous U.S. outside of California.

Seismicity

Although physiographically part of the Middle Rocky Mountains, Yellowstone lies along

the northeastern margin of the extensional basin-range tectonic region (figs. 4 and 11). Epicenters

of earthquakes associated with this tectonic extension define a belt of seismicity that extends north

from the Wasatch front in Utah to the Yellowstone Plateau, then branches to the northwest and

west into Montana and Idaho, termed the intermountain seismic belt by Smith and Sbar (1974).

Earthquake epicenters define a parabolic arc around the north, east, and south sides of the eastern

Snake River Plain through the Yellowstone Plateau, where the seismicity is most active (Smith and

Arabasz, 1991; Smith and Braile, 1994) and fault displacements are greatest (Anders and others,
9

1989; Pierce and Morgan, 1992). Earthquake distributions, focal mechanisms, and GPS

determinations of crustal strain in the Yellowstone Plateau area are all consistent with the generally

NE-SW direction of extension in the basin-range region (Waite and Smith, 2004; Puskas and

others, 2007).

Seismicity in the Yellowstone National Park region is monitored in real time by a network

of 26 seismic stations within or adjacent to the park (fig. 12), all of which are recorded at the

University of Utah as an integral part of the YVO program and archived in the Advanced Seismic

Network System. Of these seismic stations, 17 are single-component short-period seismometers

connected to the network by analog telemetry. Additionally, there are three short-period 3-

component seismometers with analog telemetry and six broadband 3-component seismometers with

digital telemetry (Yellowstone Volcano Observatory, 2006). Data are transmitted continuously

from the seismic stations via radio signals to a Federal Aviation Administration (FAA) radar site

located on Sawtell Peak, Idaho, west of Yellowstone National Park, where the signals are

multiplexed onto four FAA microwave lines for retransmission to the FAA control center in Salt

Lake City, Utah. From there, the data are transmitted to the University of Utah's central recording

laboratory via telephone lines. Earthquake data are processed using the USGS EarthWorm system

to produce automated real-time hypocenter determinations and emergency-response broadcasts.

Refined earthquake locations, magnitudes, and focal mechanisms are produced by seismic analysts

at the University of Utah Seismograph Stations from the digital data and are provided to YVO

users.

Given its average station spacing of 15-30 km and seismometer characteristics, the

Yellowstone network is optimally capable of detecting local earthquakes of magnitudes greater

than about -1.0 and locating their epicenters within about 1 km and their focal depths within about

0.5 km at depths below about 4 km (or about 1 km at shallower depths). Analysis of the digital

seismic data includes production of a revised catalog of Yellowstone earthquakes using new threedimensional P-wave velocity models determined by tomographic inversion of local earthquakes

(Husen and Smith, 2004) and a new magnitude scale for improved hazard assessments.

It is useful to consider seismicity of the Yellowstone region in terms of several categories of

activity. Most Yellowstone earthquakes (fig. 13) are of small magnitude (M3), but a few large

tectonic earthquakes have affected the region, including the deadly Hebgen Lake earthquake of

August 17, 1959, with a magnitude of 7.5 (U.S. Geological Survey, 1964; Trimble and Smith,

1975; Doser, 1985; Doser and Smith, 1989). Aftershocks of the Hebgen Lake earthquake were

numerous (Murphy and Brazee, 1964), and a high proportion of the earthquakes since then have

occurred in the same belt of seismicity as those early aftershocks (Trimble and Smith, 1975; Smith

and others, 1977; Smith and Braile, 1994). Numerous other earthquakes, including many within

the Yellowstone caldera, are scattered more widely, are typically shallow, and commonly occur in

the form of seismic swarms. Seismicity beneath the Yellowstone caldera is generally shallower (~5

km) relative to the deepest earthquakes (~20 km) on tectonic faults outside the caldera, probably

related to the effects of elevated crustal temperature above the magma that underlies the caldera

(Miller and Smith, 1999; Waite and Smith, 2002; Husen and Smith, 2004). Thus, seismicity can be

discussed in four categories: large earthquakes on regional tectonic faults, smaller tectonic

earthquakes, intracaldera earthquakes, and seismic swarms.

The 1959 Hebgen Lake earthquake had the largest magnitude of any historic earthquake in

the Rocky Mountains region (fig. 13). That earthquake produced a pair of fault scarps up to 6 m

high and 40 km long, caused a landslide that dammed the Madison River northwest of Hebgen

Lake about 30 km west of Yellowstone National Park, caused extensive damage to roads and

buildings within the park, and resulted in 28 deaths—all of them in the National Forest, outside the

park. The only other regional historic earthquake of roughly comparable magnitude was the M=7.3

Borah Peak, Idaho, earthquake of 1983, about 150 km southwest of Yellowstone National Park
10

(Dewey, 1983; Smith and others, 1985; Crone, 1987; Richins and others, 1987). It too produced a

long surface scarp and resulted in local damage. Both of these earthquakes caused discernable

changes in the hydrothermal features of Yellowstone (Watson, 1961; Marler, 1964; Marler and

White, 1975; Hutchinson, 1985). Prehistoric late Holocene scarps occur on several range fronts

near Yellowstone, including the Centennial, Madison, and Teton Ranges (pl. 1 of Pierce and

Morgan, 1992), indicating the possibility of further large tectonic earthquakes in the region.

Smaller tectonic earthquakes on regional faults occur widely in areas in and adjacent to

Yellowstone National Park (Smith and Sbar, 1974; Doser and Smith, 1983; Waite and Smith, 2002,

2004; White, 2006). Recurring seismicity of this type is common and must be expected in the

future. Most of these earthquakes are small and cause no significant damage. Others are felt

locally, and some have the potential to cause damage to structures or other facilities.

Intracaldera earthquakes are typically small (M3) and generally are shallower than about 6

km (Miller and Smith, 1999; Husen and Smith, 2004; Waite and Smith, 2004). A few however, are

larger and can be damaging. The largest earthquake recorded within the caldera area occurred

southeast of Norris Geyser Basin (fig. 13) in June, 1975 with a magnitude of 6.1 (Pitt and others,

1979). The earthquake, felt widely within Yellowstone National Park, caused damage to roads and

other visitor facilities. The area in which it was felt was more extensive to the north, outside the

caldera, than within the caldera, probably because of the attenuating effects of elevated crustal

temperatures beneath the caldera. Aftershocks occurred within a zone 10 km long that included the

main shock as well as on a parallel zone of seismicity about 5 km to the west that had been

intermittently active for the preceding year.

As is common in other volcanic and geothermal areas, many of the earthquakes in or

adjacent to the Yellowstone caldera occur in swarms (Smith and Braile, 1994; Waite and Smith,

2002; Farrell, 2007),. These swarms are characterized by numerous small, generally shallow

earthquakes clustered in both space and time, without a mainshock, and having magnitude

differences 0.25. Commonly these swarms are aligned on tectonic trends, though not necessarily

on recognized faults, and are generally considered to be related to migration of hydrothermal fluids

or magma in the shallow crust. A particularly intense earthquake swarm (fig. 14) occurred in

October and November of 1985 a short distance northwest of the Yellowstone caldera (Waite and

Smith, 2002). With time, the swarm activity migrated both away from the caldera and, to a lesser

extent, downward as the caldera itself underwent a change from uplift to subsidence. The swarm

was likely to have been related to the subsurface movement of hydrothermal fluids or magma from

beneath the caldera, deflating the caldera area and possibly penetrating the shallow crust nearby as

magmatic dikes that did not breach the surface in eruption (Waite and Smith, 2002). Possible

migration of magma may have induced local migration of hydrothermal fluids and magmatic gases

(Husen and others, 2004).

Induced or triggered earthquakes constitute another class of seismicity observed recently in

Yellowstone (Husen and others, 2004). Such earthquakes are triggered by the passage of largeamplitude surface waves from distant sources to produce local transient dynamic stress changes as

great as 0.15 MPa. A series of local earthquake swarms was observed, and clear changes in geyser

activity occurred immediately following local arrival of seismic waves from the 2002 Denali fault,

Alaska, earthquake (M=7.9) in the Yellowstone National Park area, 3100 km from the epicenter.

Beginning within hours of the arrival of surface waves, the YVO network located more than 250

earthquakes in the first days after the Denali event. The eruption frequency of several geysers was

altered, and numerous earthquake swarms close to major geyser basins were unusual in occurring

simultaneously in different geyser basins. This behavior is interpreted as having been induced by

dynamic stresses associated with the passage of large-amplitude surface waves, perhaps locally

altering the permeability of hydrothermal plumbing by unclogging existing fractures. Furthermore,

redistribution of hydrothermal fluids and locally increased pore pressures triggered earthquakes in
11

these swarms. Although earthquake triggering and changes in geyser activity have been

documented elsewhere, these notable changes at Yellowstone induced by a large-magnitude event

at such a great distance indicate the profound effects that remotely triggered seismicity can have on

hydrothermal and earthquake hazards and also suggests that earthquakes and hydrothermal

explosions at Yellowstone might be triggered by local earthquakes in the future.

Crustal deformation

Major surface deformation has been recognized in Yellowstone since precise leveling

surveys were repeated in 1975-1977 after a lapse of 50 years. Caldera-wide uplift to a maximum of

at least 73 cm (fig. 15) was documented in these surveys (Pelton and Smith, 1979, 1982). Similar

uplift continued, as measured repeatedly in several following years (fig. 16), averaging 23 mm/y

from 1976 to 1983, then slowing down for about a year before subsiding as much as 35 mm/y until

1987 (Dzurisin and Yamashita, 1987; Dzurisin and others, 1990).

Since these surveys, newer methods, including the Global Positioning System (GPS) and

Interferometric Synthetic Aperture Radar (InSAR) have shown that surface deformation remains

active over the entire area of the Yellowstone caldera and the seismically active area to the

northwest, including the epicenter of the 1959 Hebgen Lake earthquake. Both uplift and

subsidence occur in this area, sometimes in a caldera-wide fashion and at other times in more

complex patterns of more local uplift and subsidence (Meertens and Smith, 1991; Wicks and

others, 1998; Puskas and others, 2002; Wicks and others, 2006; Puskas and others, 2007).

Currently, deformation monitoring of the Yellowstone region is carried out by YVO mainly

through a network of 16 GPS receivers (Yellowstone Volcano Observatory, 2006), of which 12 are

located within Yellowstone National Park (fig. 17), continuously recorded at the University of Utah

and by the EarthScope Plate Boundary Observatory. Additional stations are planned for future

installation. Each station records a time series of 2 horizontal components and 1 vertical

component sampled every 5 seconds, which are analyzed at the University of Utah to produce

coordinate solutions. In addition, about 60 GPS stations are recorded in special surveys rather than

being monitored continuously. Repeat leveling surveys also continue intermittently.

Together, the geodetic data reveal complex patterns of crustal deformation over a period of

decades (fig. 14), as summarized by Puskas and others (2007). From 1923 to 1985 caldera-wide

uplift occurred at an average rate of 22 mm/yr, changing in 1986 to subsidence. Subsidence

continued within the Yellowstone caldera to 1995 at 14 mm/yr while uplift occurred northwest of

the caldera at 5 mm/yr. In 1995, caldera uplift resumed at 9 mm/yr, changing again in 2000 to

subsidence at 9 mm/yr while uplift continued northwest of the caldera at an average of 12 mm/yr.

A sudden change in 2004 to caldera uplift proceeded at unprecedented rates of up to 60 mm/yr,

continuing into 2007. These rapid changes in motion across the Yellowstone caldera over more

than a half century clearly reflect the importance of transient large-scale crustal deformation

sources that probably include both the subcaldera magma chamber and localized zones of

pressurized hydrothermal fluids that inflate and deflate the surface as the locations and properties of

the sources change with time.

Detailed analysis of ancient shorelines of Yellowstone Lake, both subaerial and

sublacustrine, indicates that similar changes of major uplift and subsidence date back at least as far

as the final retreat of Pleistocene glaciers from the lake basin (Meyer and Locke, 1986; Locke and

Meyer, 1994; Pierce and others, 2002; Johnson and others, 2003). Pierce and colleagues (2002)

conclude that cycles of uplift and subsidence have occurred repeatedly in the Yellowstone Lake

basin during the past 14,000 years; the record of these cycles is superposed on the record of a longterm decrease in lake level by downcutting of the lake outlet, but that little net elevation change

occurred between about 14 ka and 3 ka (fig. 18).
12

Together, earthquake focal mechanisms and both vertical and horizontal deformation

patterns suggest that the crustal stress field recorded both by earthquakes and by surface

deformation are generally consistent with regional basin-range extension. In the vicinity of the

Yellowstone caldera, however, the stress field is strongly modified by the effects of migrating

magmatic and hydrothermal fluids to produce changes in seismicity, uplift, and subsidence on time

scales of months to a few years (Dzurisin and others, 1990; Dzurisin and others, 1994; Puskas and

others, 2002; Waite and Smith, 2002; Husen and others, 2004; Waite and Smith, 2004; Wicks and

others, 2006; Puskas and others, 2007).

Hydrothermal and gas activity

At each eruption, immediately preceding, was an upheaval of some fifty feet high, followed by

one great explosion in which the water was thrown two hundred fifty to three hundred feet and

frequently hurling stone one foot in diameter five hundred feet from the crater. Joaquin Miller (in Muir, 1888), describing hydrothermal explosions of Excelsior Geyser.

Yellowstone National Park hosts one of the largest hydrothermal systems on Earth and, in

particular, has more geysers than all others in the world together. Most of these hot springs,

geysers, and fumaroles present hazards only to visitors who may stray from established pathways

into areas of unstable ground and hot water; it is impractical to consider such daily hazards in the

context of the present assessment. One of the most common acute geological hazards at

Yellowstone, however, consists of shallow-rooted explosions of steam, water, and rock without

any associated volcanism. These hydrothermal explosions (or eruptions) occur when hot

subsurface waters flash to steam, violently breaking the confining rocks and ejecting them from a

newly formed or existing crater. Individual events can last as briefly as a few seconds or as long as

several hours, and intermittent explosive activity can continue for years.

Hydrothermal-explosion craters are found in Yellowstone National Park at scales ranging

from less than a meter to several kilometers in diameter (Muffler and others, 1971; Morgan and

others, in review). Visitors to popular thermal areas like Upper and Lower Geyser Basins, Norris

Geyser Basin, and West Thumb Geyser Basin see circular, funnel-shaped pools, often with jagged

walls, filled with thermal water (fig. 19). Many of these pools appear to have formed in

hydrothermal explosions (Marler and White, 1975). Porkchop Geyser in Norris Geyser Basin,

Excelsior Geyser in Midway Geyser Basin, and Seismic Geyser in Upper Geyser Basin have

produced well-documented historical explosions (Marler and White, 1975; Fournier, 1989;

Whittlesey, 1990). Geologic research has identified many large craters formed before the time of

historical records, including a number at and near the north edge of Yellowstone Lake (fig. 20:

Mary Bay, Indian Pond, and Turbid Lake; also see fig. 24). If similar large explosions were to

occur today in that vicinity, they could threaten infrastructure and visitor safety along the East

Entrance Road, at Fishing Bridge and as far away as Lake Village. Geologic and historic evidence

in Yellowstone and elsewhere suggests that such hazardous activity could last days, weeks or even

years.

Browne and Lawless (2001) provide a lengthy summary of current knowledge of

hydrothermal “eruptions” around the world, preferring the word eruption over explosion to

reinforce the idea that the events were part of a continuum extending from geysering (no rocks

ejected) to very large rock-hurling eruptions creating craters hundreds to thousands of meters

across. While acknowledging the terminology of Browne and Lawless (2001) but to avoid

confusion with magmatic eruptions, we continue to use the term hydrothermal “explosions” in this

hazards assessment for events that can form craters or eject rocks.

We follow Browne and Lawless (2001) in distinguishing hydrothermal eruptions/explosions

from other types of ground water eruptions such as:
13

Phreatic eruption: An eruption of steam, water and rock caused when cool ground water is

heated rapidly to its boiling point by magmatic heat but no magma is actually erupted.

Phreatomagmatic eruption: Similar to a phreatic eruption but with evidence for eruption of

magma as well as fragments of the confining rock.

Magmatic-hydrothermal eruption: An eruption caused by magma heating a pre-existing

hydrothermal system; the magma may or may not reach the surface as a part of the ejected

materials. An excellent example occurred in New Zealand in 1886, when intrusion of basaltic

magma into a long fracture system caused hydrothermal explosions as far as 20 km away from the

lava vent (Hedenquist and Henley, 1985; Simmons and others, 1993).

In contrast, hydrothermal eruptions/explosions occur in pre-existing hydrothermal systems

without the proximate influence of magma. It appears that all explosions at Yellowstone since the

latest glaciation, roughly the past 16,000 years, are true hydrothermal explosions with no direct

participation of magma.

Pressures, temperatures, and fluids in geothermal systems

Understanding the causes of hydrothermal explosions requires understanding the structure

and dynamics of geothermal systems. Such systems are regions of anomalous heat flow in the

shallow crust of the Earth. Geothermal systems can be divided into two distinct types: hot-water

(or liquid-dominated) systems and vapor-dominated systems (White and others, 1971).

In a liquid-dominated system, the Earth’s subsurface is saturated with hot water, which is

located in interconnected pore spaces and fractures. Liquid-dominated systems can have two types

of surface manifestations: a) hot springs, which can represent the direct outflow of hot water from

the geothermal reservoir, and b) fumaroles, which represent steam and gas boiled off the reservoir.

At Yellowstone National Park, the hot springs of liquid-dominated systems are typically neutral or

slightly alkaline, Cl-rich, and are saturated with silica (SiO2). As these waters cool, they precipitate

amorphous silica as deposits called siliceous sinter. Waters that precipitate silica are common in,

but are not limited to, the Upper and Lower Geyser Basins, the Norris Geyser Basin, the West

Thumb Geyser Basin (Fournier, 1989) and sublacustrine hydrothermal vents in Yellowstone Lake

(Shanks and others, 2005; Shanks and others, in review).

Vapor-dominated systems lack sufficient water for complete liquid saturation of the

reservoir, resulting in low-density steam and gas being the interconnected phase. Vapor-dominated

reservoirs can still contain an appreciable amount of liquid, which remains within pore spaces in

the host rock. Surface manifestations above vapor-dominated reservoirs typically consist of “acidsulfate” features such as fumaroles, acid pools, and mud pots. Hot springs issuing neutral waters

are usually absent. Instead, acid pools and mud pots form due to condensation and oxidation of

acid gases and steam in perched bodies of surface and near-surface water. Most of the large

thermal areas in the eastern part of Yellowstone National Park (e.g., Mud Volcano, Hot Springs

Basin, Sulphur Hills, Josephs Coat) formed above vapor-dominated geothermal areas and have

acid-sulfate soils and widespread alteration of surface rocks, commonly to clays (White and others,

1971; White and others, 1975).

The simplest geothermal reservoir at Yellowstone consists of ground water at its boiling

temperature for a given hydrostatic head (fig. 21). In these situations, pressure is dictated by the

weight of the interconnected column of subsurface and surface water. Under normal conditions,

the confining pressure is sufficient to prevent the water from boiling catastrophically. Geyser

eruptions occur when subtle boiling at the top of the column induces local depressurization and

consequent boiling further down the column. Conditions such as a landslide, earthquake, or dambreak may cause a large and sudden decrease in pressure; geothermal water then may become

highly superheated relative to its boiling temperature at the new, lower pressure, causing rapid

conversion of some of the water to steam, forcing expansion, and triggering an explosion.
14

Hydrothermal explosions can be initiated more readily when pressures exceed hydrostatic. Ideally,

because the density of hot water is less than that of cold water, hydrothermal systems should be

slightly under-pressured with respect to the cooler surrounding ground water systems. However,

mineral-precipitation reactions, especially those forming silica, serve to clog permeable pathways

in the aquifer and can cause the geothermal waters to become isolated with respect to their

surroundings. Moreover, the presence of gas and steam can increase the pressure of the system. As

a result, many geothermal drill holes, including research holes drilled at Yellowstone in the 1960s

(White and others, 1975), are positively pressured with respect to the predicted hydrostatic gradient

(fig. 22).

An additional factor increasing the pressure within geothermal areas is the buildup of gases

such as carbon dioxide (CO2) and hydrogen sulfide (H2S), which do not condense upon ascent and

cooling and can build up in concentration beneath an impermeable caprock (Hedenquist and

Henley, 1985). At the vapor-dominated Ngawha system in New Zealand, for example,

extrapolations of deep gas pressures to near surface conditions could allow a hydrothermal eruption

that would lift and disperse about 130 m of overlying rock (Browne and Lawless, 2001). Thus,

vapor-dominated reservoirs also can experience hydrothermal explosions, either owing to buildup

of non-condensable gases or simply by release of superheated steam. In general, explosions from

vapor-dominated reservoirs are thought to be somewhat less dangerous than liquid-dominated

reservoirs because a given volume of steam will contain less potential energy than the same volume

of liquid at the same temperature. Even in vapor-dominated reservoirs, most of the potential

energy of the system resides within the residual liquid water (Browne and Lawless, 2001).

Mechanisms of hydrothermal explosion

The mechanisms for hydrothermal explosions recognized by Browne and Lawless (2001)

are summarized here in the context of the geothermal system at Yellowstone.

Pressures exceeding lithostatic: A hydrothermal explosion can occur if fluid pressure is

regulated by an impermeable caprock, such that pressures increase until they exceed the weight and

strength of the overlying rock. This mechanism is thought to be comparatively rare at geothermal

fields, where measured pressures are usually well below lithostatic, but eruptions from these fields

are relatively common. Moreover, the common hot springs and fumaroles at the surface above

geothermal fields imply that their caprock is not generally impermeable. At Yellowstone, it is clear

that areas affected by hydrothermal explosions are those having already-established reservoirs

connected to distinct surface expressions. Browne and Lawless (2001) infer that this mechanism is

most important above young geothermal reservoirs as they start to interact with overlying, as yet

unaltered, surface rocks. Phreatic eruptions at reactivating volcanoes also might occur when

transients in ground water pressures exceed the lithostatic load. Potentially, self-sealing due to

precipitation of minerals from hydrothermal fluids could decrease the permeability of hydrothermal

aquifers, causing pressures to approach lithostatic.

Slow accumulation of steam and/or gas: Pressures can more regularly exceed lithostatic if

steam is present in the system and can ascend, transmitting pressure to shallower regions. This is

thought to be the common mechanism for explosions in exploited geothermal fields (Browne and

Lawless, 2001) when withdrawal of geothermal fluid lowers subsurface pressures, thereby

triggering additional boiling. Because cooling of noncondensable gas has a negligible effect on its

volume, its presence will increase the likelihood of overpressure. Long-term drought and a drop in

ground water levels could cause a similar phenomenon wherein shallow steam pressures increase

with a decrease in the elevation of the ground water table (fig. 23, adapted from fig. 6 of Browne

and Lawless, 2001).

Rapid subsurface pressure release: Any geothermal system that follows the boiling point

with depth can be triggered into an eruption by a sudden release in pressure that causes flashing of
15

liquid water to steam. Such a pressure release could result from an earthquake, a landslide,

draining of a lake, deglaciation, or a hydrothermal fracturing event within the reservoir. At

Yellowstone, such events are known to have caused hydrothermal perturbations. The 1959 Hebgen

Lake earthquake (M=7.5) induced eruptions of some 289 springs, including 150 with no previous

record of geysering (Marler, 1973). Most of these eruptions were solely of liquid water and steam,

though Marler and White (1975) document the growth of a new feature, Seismic Geyser, whose

genesis involved a series of rock-hurling episodes. Muffler and others (1971) postulated that the

very late Pleistocene Pocket Basin explosion in the Lower Geyser Basin was initiated by drainage

of a postglacial lake, causing a pressure drop in the underlying geothermal reservoir.

Addition of magmatic heat or gas: Addition of external magmatic heat or gas yields what is

termed a magma-hydrothermal eruption. One might consider all heat at Yellowstone to be

ultimately magma-derived, but the term is used here to denote a rapid transfer of heat from a

shallow magmatic intrusion directly into the geothermal system. Though such a mechanism is

unlikely for the small shallow explosions most common at Yellowstone, its importance cannot be

ruled out for some of the earlier large hydrothermal explosions of the late Pleistocene and early

Holocene. If magma, however, did induce some of the large explosions, it did so without reaching

the surface.

Progressive flashing: Browne and Lawless (2001) conclude that this is the most common

type of hydrothermal explosion. It initiates close to the ground surface and works its way

downward with time as rocks fracture, causing the boiling front to move down, resulting in

increased boiling, brecciation, and depressurization of the underlying system. The model requires

that boiling water exists near the surface and overlies water at a boiling-point-for-depth temperature

gradient. As such, the initial confining pressure can be minimal, and open pools can be the site of

initiation. The eruption itself might be triggered by events like those described in the subsurface

pressure release model, but with the boiling water column destabilizing much closer to the surface.

Hydrothermal explosions in Yellowstone

At Yellowstone, a geologic record of hydrothermal explosions exists only for events

occurring after the most recent glaciation (ending about 16,000 years ago, Pierce and others, 2002),

which effectively erased evidence of any earlier such events. Evidence for eighteen large (>100-m)

hydrothermal explosions (table 3; fig. 24) is found within Yellowstone National Park (Morgan and

others, in review). Many large explosion craters are present on land, but others have been

identified in Yellowstone Lake (Wold and others, 1977; Morgan and others, 2003b; Morgan and

others, in press-b; Morgan and others, in review). Most of the largest explosion craters, such as

Mary Bay, Turbid Lake, Duck Lake, Indian Pond, and Pocket Basin (fig. 24), are found within and

along the margin of the 640-ka Yellowstone caldera. A few are present along the tectonically

controlled, north-trending Norris-Mammoth corridor. The hydrothermal-explosion craters appear

to be relatively shallow features affecting only the upper several hundred meters of substrata, which

typically have been previously affected by hydrothermal alteration (Muffler and others, 1971;

Morgan and others, in review). In Yellowstone, no large hydrothermal explosion is associated with

a volcanic event, and no evidence exists for a hydrothermal explosion triggering any volcanism.

Figure 25 displays historical hydrothermal-explosion sites identified through a literature

search. As with the prehistoric craters and deposits, the historical eruption sites are concentrated

within the caldera and the Norris-Mammoth corridor. Notably, the Upper and Lower Geyser

Basins hosted many of the observed events. Appendices 1 and 2 provide descriptions of some of

the historically observed explosions and the prehistoric explosion deposits studied.

The end result of most hydrothermal explosions is a crater, commonly water-filled,

surrounded by a berm of fragmental rocks with steep inward slopes and gentler outer slopes. The

deposits generally comprise pieces of hydrothermally altered or mineralized materials such as
16

siliceous sinter, mud, breccia, and subsurface lithologies such as lake, beach, or glacial sedimentary

rocks and rhyolitic rocks from deeper levels.

The individual fragmental deposits are quite varied and reflect the stratigraphy of

underlying rocks and the depths of rocks evacuated by the explosion, and they commonly record

complex, multiple-event histories. Silica, in the forms of quartz, chalcedony, opal, and amorphous

silica, is the most common cement of the explosion breccias and cross-cutting veins, but pyrite,

calcite, and zeolites also are common hydrothermal minerals.

The complexity of the deposits implies repeated explosion events over a wide range of

scales, with lithic fragments being ejected, falling back into or adjacent to the vent, and being

cemented with silica to form breccias. Multigenerational heterolithic breccias also are common,

indicating an extended process of repeated brecciation and hydrothermal cementation that likely

occurs in hydrothermal systems well below the explosion vents (Keith and Muffler, 1978; Morgan

and others, in review).

Factors contributing to hydrothermal explosions

Several factors contribute to the likelihood of hydrothermal explosions. Prior to the

eruption of Porkchop Geyser in 1989, Fournier and others (1991) had interpreted chemical

indicators as showing that the temperature of the water feeding the hot-spring pool had been

increasing with time. Increased temperature of the deep water would have had two primary effects.

First, it would have increased the amount of boiling and therefore steam production as the water

rose towards the surface. Second, it would have increased the amount of silica in the ascending

water and thus the supersaturation of silica as the water cooled during its rise. The first factor

would have increased pressure in the system as the steam/liquid ratio in the subsurface increased in

the relatively constant-volume system. The second factor would have decreased permeability in the

geyser’s conduit, potentially causing a decrease in the rate at which water could flow through the

system. Thus, increased heat to the system increased the likelihood of pressurization and

hydrothermal explosions.

Another influence on the likelihood of hydrothermal explosions is the depth of the vaporliquid interface above a liquid-dominated geothermal reservoir (fig. 23), as this depth controls the

pressure of any vapor reservoir near the surface. If rocks of low permeability overlie a vapordominated region, then the pressure can rise if any or several of the following things should

change: a) increased heat supplied to the system, b) increased amount of gas or steam rising

through the system, or c) reduced ground water recharge to the system causing the geothermal

water table to fall. Any of these factors would cause an increase in the thickness of the steam

reservoir, resulting in greater pressures transferred toward the surface through the vapor/steam

reservoir. This process occurs commonly at geothermal wells, where gas must be “bled off” to

prevent displacement of water and lowering of the water level within the well. If gas or steam is

allowed to accumulate in idle geothermal wells, the results can be disastrous. Because the ground

surface at Yellowstone is generally permeable above vapor-dominated areas, allowing abundant gas

flux (Werner and Brantley, 2003), this mechanism may not be a leading cause of explosions at

Yellowstone but might be so on occasion.

Rapid pressure reduction is commonly invoked as a cause of hydrothermal destabilization.

Withdrawal of fluids from geothermal wells is documented to have lowered subsurface pressures

and induced boiling and hydrothermal explosions in geothermal fields (Scott and Cody, 2000).

Earthquakes, deglaciation, or lake drainage all could cause sufficient depressurization of a boiling

aquifer to cause hydrothermal explosions at Yellowstone (Muffler and others, 1971; Morgan and

others, in review). Bargar and Fournier (1988) demonstrated that parts of the geothermal reservoir

were superheated by 20-50°C following deglaciation, at about 12 to 15 ka. This is presumably

because while glacial ice was present, areas beneath the ice-rock contact were pressurized
17

compared with ice-free conditions, allowing a higher boiling temperature for H2O. After

deglaciation, considerable boiling would be necessary to re-equilibrate the water column to the

new, lower-pressure conditions. Although this temporary instability would not likely be a direct

trigger for hydrothermal explosions, it could be a possible contributing factor; whether any

explosions resulted directly from deglaciation and attendant depressurization is not known.

Earthquakes, extremely common at Yellowstone, are also known to have strong effects on

geothermal features (Watson, 1961; Marler, 1964; Marler and White, 1975; Husen and Smith,

2004). Some of the large prehistoric hydrothermal explosions could have been triggered by

destabilization due to passage of large-amplitude seismic waves that dynamically increase local

stress in hydrothermal reservoirs.

Although several factors can be identified as potential triggers for hydrothermal explosions

at Yellowstone, the reality is that there are only sparse observational or scientific monitoring data

on past explosive events. Geologic evidence for triggering mechanisms for past events is rare and

ambiguous though current studies are aimed at detecting such information where it may exist.

Toxic gases

After steam, carbon dioxide (CO2) is the most common constituent of volcanic gas and can

be emitted in sufficient quantities to pose a hazard at many volcanic and geothermal systems

around the world (Baxter, 2005). After subtracting steam, CO2 typically constitutes 95 to 98% of

the gas emitted from Yellowstone’s fumaroles and bubbling pools (Werner and Brantley, 2003).

Carbon dioxide is non-toxic in low concentrations and makes up about 0.038% of the Earth’s

atmosphere. However, because it is about 50% heavier than normal air, it can accumulate to much

higher concentrations in soils and low or protected areas such as valleys and caves. Carbon dioxide

concentrations of >10% are toxic to humans and animals. When air contains over 20-30% CO2,

even a few breaths can quickly lead to unconsciousness and death from acute hypoxia, severe

acidosis, and respiratory paralysis (Hill, 2000). Hypoxia is a condition in the body resulting from

the displacement of oxygen such that it inhibits normal metabolism. Acidosis occurs when CO2

acidifies the blood, causing irreversible cellular damage. In the Yellowstone region, the local

concentration of CO2 to potentially toxic levels is generally only temporary and is restricted to

confined or topographically low areas of hydrothermal activity, generally under windless

conditions.

Volcanoes commonly emit acid gases like sulfur dioxide (SO2) and hydrogen chloride

(HCl). For example, at the volcanoes Kilauea in Hawaii and Masaya in Nicaragua, these acid gases

form aerosols that are a chronic hazard to both local vegetation and human populations (Baxter,

2005). At Yellowstone, the hydrothermal system and its host rocks act as a buffer that absorbs and

neutralizes acid gases, forming hydrogen sulfide (H2S) gas and soluble sodium chloride.

Additional H2S may rise directly off the magma. The concentration of H2S is typically about 20 to

200 times less than that of CO2, but its toxicity is much greater. Though H2S has an extremely

strong “rotten-egg” odor at levels of only a few parts per billion (ppb), concentrations of more than

10 ppm in the air will rapidly deaden the human sense of smell to its presence (Mandavi, 2005).

Concentrations of 100 ppm can cause severe eye and throat irritation, and at concentrations

exceeding 700 ppm loss of consciousness and death can occur rapidly. Hydrogen sulfide forms a

complex bond to iron in mitochondrial cytochromes, thereby arresting aerobic metabolism in an

effect similar to cyanide toxicity (Milby and Baselt, 1999).

The Hazards

The geologic setting, geophysical activity, and hydrothermal systems reviewed above

provide the framework within which to consider potential hazards arising from any future volcanic,
18

hydrothermal, or gas-emission activity. For each of these categories of potential hazards,

discussion is organized in terms of different types of activity that might develop as the Yellowstone

magmatic-tectonic-hydrothermal system continues to evolve.

Volcanic-eruption hazards

As is characteristic of many large continental magmatic systems, eruptive activity in the

Yellowstone Plateau volcanic field is highly episodic and involves long periods of time between

eruptions; many of the eruptions are quite large. Such systems are particularly difficult to evaluate

in terms of the probabilities of hazardous events and the consequent risks to people and resources.

In order to bring some coherence to this discussion, it is organized by considering first the smaller,

more likely future volcanic eruptions and proceeding to the larger and potentially most destructive

but least likely events.

Basaltic eruptions

As noted earlier, basaltic lavas have erupted around the margins of the active, mainly

rhyolitic Yellowstone Plateau volcanic field throughout its evolution. The absence of basalts from

within the rhyolitic source areas is interpreted to reflect the trapping within the crust of any basaltic

magmas that might have intruded from zones of partial melting in the upper mantle beneath crustal

rhyolitic magmas of lower density (Christiansen, 2001). Only after about a million years have

basalts erupted through the cooled, crystallized, and fractured upper-crustal magmatic sources of

the first and second rhyolitic cycles; no basaltic vents, however, occur within the third-cycle

Yellowstone caldera. A few small outcrops of basalt do occur on the northwest caldera wall near

Purple Mountain (Christiansen and Blank, 1974), but they are erosional remnants of lavas that

flowed down the steep slope from vents farther north. Additionally, some rare quenched inclusions

of basaltic magma were found within the basal part of the rhyolitic West Yellowstone flow near the

crest of the Madison Plateau west of Little Firehole Meadows (R. L. Christiansen and H. R. Blank,

Jr., unpubl. data), suggesting that basaltic magmas from lower-crustal levels might have played a

role in mobilizing some intracaldera rhyolitic magmas for eruption.

Most postcaldera basalts of the Yellowstone area are glaciated erosional remnants of oncelarger flows (fig. 10). Extrapolation to the likely outlines of initial distributions suggests that most

individual basaltic eruptions covered areas of less than about 5 km2

 and did not exceed 0.1 km3 in

volume. The largest single basaltic flow field, however, the Falls River Basalt, extends from the

southwestern caldera margin to Henrys Fork of the Snake River, a distance of about 60 km. The

Falls River Basalt flow field may have covered 900 km2

 and may account for an eruptive volume of

nearly 20 km3

. At least two vents and possibly more fed the flow field, only one of which is now

exposed (fig. 10). The next largest basaltic flow field near Yellowstone National Park, the Gerrit

Basalt covering about 100 km2

 in the area of Island Park, west of Yellowstone, erupted from at

least 13 vents. Basalts cover the Eastern Snake River Plain west of Island Park, and at least 4 vents

for basalts of the Snake River Group are within or immediately adjacent to the basin of Island Park

(Christiansen, 1982).

The postcaldera basalts of the Yellowstone region—generally pahoehoe flows—are mainly

olivine tholeiites having a range of K2O contents (but commonly <0.5%). Lavas producing such

flows would be expected to have low viscosities and to erupt rapidly. Probably most of the basalts

erupted in events lasting no more than a few weeks to a few months, but a large flow field like the

Falls River Basalt might have accumulated in multiple eruptions, each lasting many months. Most

of the recognized basaltic vents are localized agglutinated scoria accumulations or small lava

shields, but some basalts may have originated as dike-fed fissures. A few formed cinder cones.
19

A total of 33 postcaldera basaltic vents (fig. 10) have been recognized or inferred in the area

immediately surrounding the Yellowstone caldera (Christiansen, 2001; Smith and Bennett, 2006).

Of these, 17 are within or adjacent to the Island Park basin; the others are scattered around all

sectors of the Yellowstone-caldera margin. Because of their generally scattered distribution, most

of the vents probably remain preserved at the surface, but some might have been buried by younger

materials or not recognized during geologic mapping. If it is assumed that 80 percent of the actual

vents have been recognized, there could have been as many as 40 basaltic eruptions around the

periphery of the Yellowstone caldera in the past 640,000 years. Of these about half occurred in

Island Park. No postglacial basaltic eruptions have been recognized, indicating that none has

occurred within the past 16,000 years.

On the basis of the foregoing, the average period between basaltic eruptions in the area

around Yellowstone National Park during post-Lava Creek time is about 16,000 years. The

average annual probability (i.e., the number of events divided by the time period, in years, during

which they have occurred) of a basaltic eruption occurring somewhere around the periphery of the

Yellowstone Plateau volcanic field is therefore 6x10-5. However, it is unclear whether the basaltic

eruptions, like some of the Yellowstone rhyolitic eruptions, may have been clustered in time; if so,

the long-term average recurrence period may have little direct bearing on future eruption

probabilities. The most likely location for any such future basaltic eruption is within the basin of

Island Park but could be anywhere else within a band about 40 km wide surrounding the

Yellowstone caldera. Any such eruption is most likely to last between a few weeks and several

months. It is possible but unlikely that basalt could erupt from within the caldera; if such an event

were to occur, however, it would signal the demise of the large Yellowstone-caldera rhyolitic

magma chamber.

The principal hazard likely to result from a basaltic eruption around the periphery of the

Yellowstone caldera would be coverage of an area of several square kilometers by lava, one to a

few tens of meters thick. In addition, basaltic ash and cinders from the eruptive vent might blanket

areas of many hundreds of square kilometers to depths of a few meters to a few centimeters,

decreasing in thickness outward from the vent in directions determined by the prevailing winds. If

a basaltic vent were to emerge from beneath shallow water or a large area of saturated ground,

phreatomagmatic eruptions could produce pyroclastic surges within the proximal area that could

blast down trees and cause other similar destruction.

In addition to any primary hazards of lava inundation and ash blanketing, there are likely to

be secondary hazards from any basaltic eruption in the Yellowstone region. In particular, such an

event would be likely to start fires around an advancing flow front, particularly under dry

conditions. Debris flows or floods could be triggered by the melting of any significant snow pack

or by temporary blockages of major drainages and subsequent release of floodwaters as the

blockage was undermined or overtopped by rising waters.

Given the ongoing YVO monitoring program, it is likely that there would be recognizable

premonitory indicators of any impending basaltic eruption. In particular, multiple shallow

earthquake swarms focused in a small, probably linear area would be likely to be followed or

accompanied by volcanic tremor, as has been observed at many basaltic volcanoes as they prepare

to erupt. It is likely that surface ground cracks would open in the immediate stages prior to any

basaltic eruption as a dike approached the surface. Because YVO deformation monitoring is

focused on the Yellowstone caldera, initial uplift associated with the shallow intrusion of basaltic

magma peripheral to the caldera might not be recognized quickly.

The emission of magmatic gases to the surface would be a major indicator of impending

eruptive activity but might be quite difficult to recognize in the presence of Yellowstone’s massive

hydrothermal system, which tends to scrub out magmatic gases passing through it (Symonds and

others, 2001). Nevertheless, any locally increased emissions of CO2 or H2S should be monitored
20

closely; any appearance of SO2, presently absent at Yellowstone, would be especially indicative

that the hydrothermal system was becoming dried out by shallow magmatic intrusion. Even

without SO2, significant localized increases in the ratio of sulfur gases to carbon gases would

suggest the possibility of magmatic gas reaching shallow subsurface levels.

Because no basaltic eruptions have occurred in more than 16,000 years at Yellowstone,

there are no well-established magmatic pathways. Thus, it is most likely that premonitory

seismicity would be sufficiently prominent and of long enough duration to allow temporary

monitoring of ground deformation and gas emission to provide additional information for

interpreting possible locations and the nature of any such eruptive event. Among the earliest

indicators might be relatively deep long-period seismicity. The most immediate precursors

probably would occur as a shallow intrusion entered the hydrothermal system, generating very

active short-period seismicity and possibly hydrothermal explosions. However, the distinction

between an impending basaltic or small rhyolitic eruption might be difficult to evaluate before the

initial venting.

Because the basaltic lava flows of Yellowstone are virtually all tube-fed pahoehoe, their

advance would probably be slow enough to allow mitigating measures to be taken except, perhaps,

in areas close to erupting vents.

Large rhyolitic eruptions

At least 17 large rhyolitic lava flows, most of them with volumes of 10 km3 or greater, have

erupted within the Yellowstone caldera during about the past 170,000 years (Christiansen, 2001).

Stratigraphically they belong to the Central Plateau Member of the Plateau Rhyolite. Each of these

lava flows extruded through one of two linear vent zones that cross the caldera along the

extrapolated positions of extracaldera tectonic fault zones (figs. 8, 11), and activity has been

essentially contemporaneous on both zones. The largest of these lava flows cover areas greater

than 350 km2 and have volumes greater than 30 km3 (table 2). The constructional topography

formed by these flows is represented by the Pitchstone, Madison, and Central Plateaus (figs. 1 and

26). The Pitchstone-Madison Plateau alignment lies on the extrapolated position of the Teton

normal-fault zone (fig. 11) and extends to the southwest edge of the tectonic West Yellowstone

basin (fig. 8). The Central Plateau alignment (fig. 8) lies on an extrapolation of both the Sheridan

normal-fault zone (fig. 11) and the extracaldera Norris-Mammoth corridor (fig. 10).

As these young intracaldera rhyolite flows are petrographically, chemically, and isotopically

distinct from early intracaldera rhyolites (Hildreth and others, 1984; Hildreth and others, 1991) and

have an age range separated by about 80,000 years from the youngest known earlier intracaldera

flows, they are considered here together as a group separate from the older postcaldera lavas. The

hazard potential for a possible future intracaldera eruption may be best reflected in the

characteristics of this group of voluminous rhyolites of ~170 ka and younger.

Available K-Ar and 40Ar/39Ar dating of rhyolitic flows of the Central Plateau Member (table

2; figs. 9, 26) indicates that they erupted in a few major episodes. The earliest of these episodes

probably began with the Mallard Lake flow that erupted from the Central Plateau vent alignment

and covered the eastern part of the Mallard Lake resurgent dome (Christiansen, 2001). It was either

accompanied by or immediately followed by renewed uplift of the Mallard Lake dome.

Subsequent eruptions, within a period of less than 10,000 years, produced the Dry Creek, West

Thumb, and Mary Lake flows along the Central Plateau vent zone and the Buffalo Lake flow along

the Madison Plateau vent zone (fig. 9). Additional rhyolitic lavas, now buried, may well have

erupted during this time from either or both of those zones. In addition to these lavas, at least one

major pyroclastic eruption occurred along the Central Plateau vent alignment to produce the tuff of

Bluff Point. Although its volume cannot be reconstructed convincingly because of erosion and

burial by younger lavas, eruption of this pumiceous pyroclastic unit was sufficiently large to
21

produce a source caldera 10 km in diameter, now represented by West Thumb, the westernmost

basin of Yellowstone Lake (fig. 8). The eruptive volume almost certainly was several tens of km3

.

The actual time span represented by eruptions of these oldest units of the Central Plateau Member

is not defined precisely. The nominal span of weighted-mean ages for the dated lava flows and tuff

of this group is 173±11 to 160±3 ka, but disagreements between individual age determinations and

the stratigraphic order interpreted from geologic mapping suggest that some of the isotopic ages are

incorrect. The group mean age and standard deviation of the individual weighted-mean isotopic

age determinations on all of these units is 167±5 thousand years, and all the weighted-mean ages of

the individual units of this group overlap that range. Thus, the range might be a closer

representation of the actual span of time involved in erupting them than the nominal span of

individual age determinations.

Distinct in time but only shortly after the first group of eruptions along the Central Plateau

vent alignment (fig. 8) were the Aster Creek, Elephant Back, Spruce Creek, and Nez Perce Creek

lava flows (fig. 9). The undated Spring Creek flow may have vented on the Madison Plateau vent

alignment during this same episode. At least one significant pyroclastic unit, the tuff of Cold

Mountain Creek, is also interstratified among lavas erupted along the Madison Plateau alignment

although its source location is not known. This tuff may represent early pyroclastic activity from

the vent for a large lava flow of the Madison Plateau. The age of this group of rhyolitic eruptions is

about 150±5 ka, calculated similarly to that of the earlier group.

Another relatively brief episode of large-volume rhyolitic eruptions within the Madison

Plateau vent zone formed the Summit Lake, Bechler River, and West Yellowstone flows (fig. 9);

the smaller Douglas knob and Trischman Knob domes also probably erupted during this episode,

possibly as late-stage vent domes of the Bechler River flow (Christiansen, 2001). The area covered

by flows of this age group is smaller than for the older groups, and all flows erupted during this

sequence probably are represented by surface exposures. Nominal weighted-mean K-Ar ages of

these lavas range from 124±10 to 114±1 ka, but just as for the older groups, stratigraphic relations

suggest that the actual time span of eruption was short. A mean age and standard deviation of the

individual weighted-mean age determinations on these flows is 118±5 ka.

Yet another episode of relatively large rhyolitic lava eruptions vented along the Central

Plateau at about 102-103 ka, producing the Hayden Valley and Solfatara Plateau flows (fig. 9).

The age of these flows, calculated similarly to those of the older groups, is 103±1 ka.

The most recent episode of large intracaldera rhyolitic lava eruptions is represented by the

adjacent Grants Pass and Pitchstone Plateau flows (fig. 9), with weighted-mean ages of 72±3 and

79±11 ka, respectively. Both flows may well represent a single eruptive event, the linear vent for

the Grants Pass flow representing an early dike phase, and the Pitchstone Plateau flow issuing from

a longer-lived central-vent phase. The weighted-mean age and standard deviation of these flows is

76±5 ka.

Several factors suggest a model for systematic evolution for rhyolitic eruptions of the

Central Plateau Member that may have implications for related volcanic hazards. Volumes of the

Central Plateau lavas (table 2) were calculated from geologic maps using a cut-and-fill estimator in

ArcInfo, a Geographic Information System; those calculations proceed by estimating the original

extent and a computer-generated extrapolation of a flat bottom to each flow, resulting in a

minimum-volume estimate and, thus, probably an underestimate of the total volume of all the

flows. The aggregate volume of the oldest group, at 167±5 ka, is by far the greatest, at least 138

km3

. That volume, however, is possibly even greater, ~400 km3 as suggested by a separate

calculation of the total volume for all Central Plateau Member lavas of >600 km3

, with the

difference between these calculations listed as “unobserved units” in table 2. (An earlier estimate

of the total volume of the Central Plateau Member by Christiansen (2001) was 900 km3

). The

aggregate volume of the second group of lavas, at 150±5 ka is at least 62 km3, and that of the next
22

group, at 118±5 ka, is about 86 km3. The volume of the 103±1-ka pair is about 9 km3, and that of

the youngest pair, at about 76±5 ka, is about 70 km3

. (These volume estimates do not include the

tuffs of Bluff Point and Cold Mountain Creek, for which relevant data are inadequate). Thus, large

volumes of rhyolite erupted in episodes separated by roughly 15,000 to 30,000 years from each of

the two linear extensional zones across the Yellowstone caldera where they intersect the caldera

ring-fracture system. More than 70,000 years have ensued since the latest eruptions.

Isotopic analyses of oxygen on most stratigraphic units of the Yellowstone Plateau volcanic

field were reported by Hildreth and others (1984). Oxygen, being by far the most abundant

chemical element in any igneous rock, is particularly useful in tracing processes that affect the

magmatic system as a whole. Although quartz phenocrysts in the oldest rocks of the volcanic field

have 18O >7‰—more or less typical igneous values—quartz in lavas erupted shortly after the

first-cycle caldera-forming eruption had sharply depleted 18O, to values as low as ~4‰, resulting

from massive assimilation of low-18O hydrothermally altered rocks or meteoric water. Lavas and

tuffs erupted over nearly the next 1.5 million years recovered gradually in 18O indicating that

infusions of magma from greater depths successively augmented the upper-crustal magma

chamber; by the time of the climactic Lava Creek Tuff eruption at 639±2 ka, 18O had recovered to

values in quartz as high as 6.8‰. Formation of the Yellowstone caldera, however, was again

accompanied by massive assimilation of low-18O materials; quartz in early intracaldera lavas

drastically reduced in 18O to values as low as ~1‰. Since then, 18O in intracaldera lavas and tuffs

has again gradually recovered with successive additions of high-18O magma; during the

intracaldera eruptions of ~170,000-70,000 years, 18O in quartz has increased irregularly from ~4.6

to 5.1‰.

Trace-element compositions of the Central Plateau Member lavas became progressively

more evolved with time (Hildreth and others, 1991; Vazquez and Reid, 2002). Vazquez and Reid

(2002) showed that zircons in the Central Plateau Member lavas have 238U/230Th model ages that

generally range from their eruption ages back to about 200 ka—compared to U-Pb zircon data from

the Lava Creek Tuff that show its magma to have been generated <100,000 years before the

caldera-forming eruption (Bindeman and others, 2001). Thus the Central Plateau Member lavas

can be regarded as representing an incrementally augmented body of upper-crustal rhyolitic magma

beneath the Yellowstone caldera. Furthermore, these late postcaldera lavas show evidence of

progressive differentiation throughout the time of their eruption. The ratio Rb/Sr, for example,

generally increases with decreasing ages of the flows. Vazquez and Reid (2002) modeled mineral

and whole-rock trace-element chemistry as being consistent with the sequence of lavas postdating

formation of the Yellowstone caldera representing at least 40% sanidine-dominated fractional

crystallization of the residual magma body. This magma, initially low in 18O, was augmented

progressively by infusions from deeper magma. Through most of this time, in the Vazquez and

Reid model, the lavas erupted intermittently while sanidine and other major phenocryst phases

fractionated progressively to accumulate below in a crystal-liquid mush but zircon did not. In fact,

it may well be that any preexisting zircon and plagioclase were resorbed in the residual magma.

After about 120 ka, new zircon crystals grew in magmatic liquid separated from the underlying

mush for as long as 50,000 years before eruption to produce the youngest of the Central Plateau

lavas, the Pitchstone Plateau flow.

The successive eruption of these voluminous, progressively more evolved lavas from a

single magma body suggests a model in which magma is augmented incrementally but slowly

crystallizes and differentiates in a postcaldera Yellowstone magma chamber while the chamber roof

is progressively cooled from the top by hydrothermal circulation. A mixture of magmatic liquid

and crystals is envisioned to remain in the chamber until fractures driven by regional tectonic

extension in the cooling chamber roof become large enough to extend downward toward the level

of magma storage. These fractures then enable the intrusion of rhyolitic dikes, allowing the magma
23

to extrude and erupt to the surface. The intrusions heat the chamber roof, reducing its elastic

strength, sealing and inhibiting further development of large, deep fractures, and tending to retard

further eruption. Only as hydrothermal circulation again cools the rocks of the chamber roof can

fractures of sufficient size again start to develop and promote renewed tapping of liquid from the

magma chamber, renewed diking, and surface eruption. With time, magmatic liquid in the chamber

continues to differentiate, but smaller volumes remain sufficiently liquid to allow mobilization and

eventual eruption. Thus, the rhyolites erupt in episodes of slightly more chemically evolved

compositions. This model implies that any magmatic increments introduced into the upper-level

chamber between ~170 and 70 ka intruded at a rate lower than the rate of differentiation within the

magma body.

Renewal of large rhyolitic eruptions within the Yellowstone caldera would present two

major classes of hazards, explosive pyroclastic ejections and effusive lava eruptions. The initial

phases of an eruption would undoubtedly involve pyroclastic eruption through an opening vent.

Pyroclastic deposits probably lie buried beneath each of the large Central Plateau Member flows

but remain mostly unexposed. Their distal equivalents, however, have been recognized in

sedimentary basins surrounding Yellowstone, perhaps including ash beds designated Hebgen

Narrows and Natural Trap by Izett (1981). Pumiceous ash mapped as the tuff of Cold Mountain

Creek may also represent pyroclastic materials erupted in vent-opening phases of one or more of

the Central Plateau Member lavas. Burial by pumice lapilli and ash from such vent-opening

eruptions could be widespread and locally as thick as many meters. Pyroclastic flows could

devastate areas of many tens of square kilometers. Subsequent lava extrusion could last many

years, covering areas as great as 350-400 km2

 to thicknesses of tens or hundreds of meters and

volumes of 5 to more than 50 km3. Because such voluminous rhyolitic lava flows are

unprecedented in the global historical record and their extrusion rates may be considerably greater

than those of observed smaller flows, the actual length of time that an individual eruption might

continue is uncertain. Certainly no infrastructure in areas affected by such lavas would survive.

Note should also be taken here of the tuff of Bluff Point, the only voluminous ash-flow tuff

of the Central Plateau Member; its eruption resulted in formation of the 10-km caldera now

represented by the West Thumb. The Bluff Point eruption was characterized by ejection, not only

of pumice, ash, and small lithic fragments, but also of vitrophyric lava blocks as large as 1 m and

possibly involved copious amounts of water from a predecessor of Yellowstone Lake. Based on

this example, it is reasonable to suggest that any future voluminous rhyolitic eruption within or

immediately adjacent to Yellowstone Lake might have the potential to produce an explosive

pyroclastic eruption, much smaller in volume than a major caldera-forming event like the Lava

Creek Tuff but nevertheless capable of producing high-speed flows of hot ash, rhyolitic debris, and

volcanic gas that could affect virtually all of the Yellowstone Plateau area and produce regionally

widespread finer downwind ash.

Secondary hazards associated with large intracaldera rhyolitic eruptions could include

wildland fires, disturbed drainages, and possibilities of catastrophic flooding, and they could be

especially severe and widespread given their great volumes. In particular, if a large-volume lava

flow should displace a considerable volume of Yellowstone Lake, overflow could inundate large

parts of the Yellowstone River drainage. Other large drainages could be impounded by lava, then

released as catastrophic floods overtopped or undermined the lava dams.

Estimating probabilities for one or more large intracaldera rhyolitic eruptions involves

considerable uncertainty. Considering that 17 known voluminous lava eruptions and 2 moderately

large pyroclastic eruptions are documented and that additional flows probably are buried within the

caldera, the raw probability might be estimated on the basis of approximately 25 events between

about 170,000 and 72,000 years ago, an average intereruptive period of about 3900 years during

that time. Alternatively, if it is assumed that the probability of such an eruption has remained
24

uniform from ~170 ka to the present, the average period between eruptions would be 6800 years,

the latest of which was ~72,000 years ago, or an annual probability of 1.5x10-4. Clearly, however,

such calculations are misleading. The voluminous eruptions were highly episodic, with numerous

events occurring within periods of a few thousand years and longer periods of ~12,000 to 38,000

years ensuing without eruption and about 72,000 years since the latest of them. Furthermore,

generally decreased numbers of eruptions in successive episodes, a doubling of the post-eruptive

interval compared to the longest previous intereruptive intervals, and the fact that the latest of the

five eruptive episodes was possibly a single voluminous event might even suggest that the

repeating cycle has ended.

The foregoing calculated probabilities are based implicitly upon the assumption that the

probability of an eruption occurring within a given time does not depend upon the time elapsed

since the last previous eruption (statistically a Poisson, or simple exponential process). Because the

large intracaldera eruptions of about the past 170,000 years are clustered in time, however, a more

meaningful probabilistic calculation might be based upon a different statistical basis, a mixed

exponential process (Cox and Lewis, 1966; Nathenson, 2001). Because of the isotopic evidence for

reorganization of the magma at about 200 ka and distinct changes in eruptive behavior and

petrologic characteristics of intracaldera lavas since about 170 ka, only the time since 170 ka is

considered in the following analysis. (A complete discussion of this analysis and the relevant

figures is given in appendix 3; only the important conclusions are summarized here).

The chronology for all known Yellowstone intracaldera eruptions since 170 ka is shown

versus event number in figure 27. Because of uncertainties in the ages of these eruptions, this

figure shows three possible models for the ages for individual lava flows within each episode.

Intereruptive intervals during certain periods are short, averaging less than 8 ky, and may be much

less than that. There are several longer interruptions of activity: one is 38±4 ky (thousand years),

another is 30±6 ky, and two others are 12±6 ky. The resulting probability distribution (fig. 28)

illustrates this behavior; the mixed exponential model curve matches the data well.

Using the mixed exponential probability distribution, the conditional probability that an

intracaldera lava eruption will occur in the next year (fig. 29) decreases as the time since the last

eruption increases, changing very little after about 20 ky. The mixed exponential probability of an

intracaldera rhyolitic lava eruption in the next year, ~72 ka after the last eruption, is equivalent to

an annual rate of 5x10-5 or a recurrence rate of 1 in 20,000 years.

In addition to the probability of an eruption occurring in a given time, an important

consideration is its likely size. Figure 30 shows the probability of a lava eruption within the caldera

having a volume greater that any particular value, based upon volumes estimated for the 19 known

late intracaldera lava flows or domes (table 2). The data are logarithmic with volume over most of

the range, with large-volume lava flows more probable in the caldera than smaller ones. The

probability that the next intracaldera lava flow will have a volume greater than 5 km3

 is about 60%;

the probability of a volume greater than 10 km3

 is about 40%.

If a large intracaldera rhyolitic eruption were to occur, there would almost certainly be

significant premonitory indications. Seismic swarms would probably be concentrated in any area

of diking as magma rose from the magma chamber toward the surface. Such seismicity would

most likely embrace the entire thickness of the brittle part of the crust above the chamber and might

well display a pattern of decreasing focal depths with time. Furthermore, an initial intrusive phase

would probably be indicated by a linear array of epicenters although they might or might not

centralize toward a smaller area before the opening of an eruptive vent. Surface deformation

should be readily detectible, and the volume of localized inflation—distinguished from such areally

generalized inflation or deflation as measured within the caldera in recent decades—would likely

presage the arrival of lava at the surface. Vent opening might well involve the local ground water

system to produce phreatic or phreatomagmatic explosions. Initial pyroclastic eruption could
25

extend to stratospheric heights, and its duration and magnitude might be considered as indicative of

the volume of possible subsequent lava extrusion.

Small rhyolitic eruptions

The possibility of future rhyolitic lava eruptions in the Yellowstone system involves two

more or less distinct scenarios. In addition to the large intracaldera eruptions just discussed,

relatively small eruptions (a few cubic kilometers or less) might occur either within or outside the

Yellowstone caldera.

Most of the small postcaldera eruptions of rhyolitic lava have occurred at locations outside

the Yellowstone caldera. Of the 13 such eruptions known (fig. 10), all but one vented north of the

caldera, either in the Norris-Mammoth corridor or between the Madison River and the southern

Gallatin Range. A single postcaldera rhyolitic lava flow erupted south of the caldera, near the

south entrance to Yellowstone National Park (fig. 10). Most of these extracaldera rhyolites are

volcanic domes with diameters of 0.5-2.5 km and volumes less than 0.5 km3

. Only four are larger

flows that cover between 2 and 20 km2 or have volumes as great as 4 km3.

Several of the small rhyolites are intimately associated with basalt that occurs both as

quenched magmatic inclusions and larger masses within the rhyolite; the largest of these are the

Gardner River and Grizzly Lake rhyolite-basalt mixed-lava complexes. Two rhyolites, the

Obsidian Cliff and Crystal Spring flows, completely lack phenocrysts, indicating that they were

emplaced at or above their liquidus temperatures; the Crystal Spring flow also contains abundant

quenched magmatic basaltic inclusions near its base. These observations suggest a role for basaltic

magma in mobilizing the rhyolitic magma for intrusion to shallow depths. Geologic, geochemical,

and isotopic data show that the extracaldera rhyolites erupted from individual magma sources that

were related to local basaltic magmas but unconnected to the large rhyolitic magma chamber

beneath the Yellowstone caldera (Hildreth and others, 1984; Hildreth and others, 1991).

Although individual small rhyolitic eruptions have vented from several locations around the

Yellowstone caldera, by far the greatest number erupted within the Norris-Mammoth corridor. All

of those younger than about 170,000 years (table 1), equivalent in age to the youngest intracaldera

rhyolites, are located either in the Norris-Mammoth corridor or near the South Entrance to

Yellowstone National Park, within 25 km of the caldera margin. Unlike the major normal fault

zones such as the Teton, Sheridan, and Hebgen Lake zones (fig. 11), faults in the Norris-Mammoth

corridor are generally of small to moderate displacement. By contrast, the principal fault along

which the east side of the Gallatin Range is relatively uplifted, had a net late-Cenozoic stratigraphic

throw of about 2 km. Nevertheless, it has been inactive during postglacial time or longer. It was

suggested by Pierce and others (1991), that magmatic diking along the Norris-Mammoth corridor

postdating the Lava Creek Tuff, both rhyolitic and basaltic, effectively takes up a significant

proportion of the extensional strain that otherwise would be expressed in normal displacement on

the fault zone bounding the east side of the Gallatin Range.

Two distinct types of primary hazards might be associated with a small rhyolitic eruption in

the Yellowstone area. The initial vent-opening phase of such an eruption would almost certainly be

an explosive ejection of rhyolitic pumice. Most of the coarser pumice would be deposited close to

the vent, but finer ash probably would be dispersed downwind over distances of many kilometers.

Thicknesses of the downwind deposits could range from several meters to a few centimeters or less.

In addition, an explosive phase could generate pyroclastic flows, high-speed mixtures of hot ash

and gas capable of devastating areas within several kilometers of the erupting vent. Ash falls

could cover large areas and cause damage to structures, power lines, and machinery by loading

them, especially if such an eruption should occur under wet atmospheric conditions. Water

supplies also could be damaged by substantial ash falls. If pyroclastic flows occurred, they would

be highly dangerous and destructive to the local environment or to any structures or people in the
26

vicinity of the eruption. Such initial pyroclastic eruptions would likely last between a few hours

and a few weeks but quite likely would be succeeded by extrusion of viscous rhyolitic lava to form

a dome or small lava flow. Areas up to several square kilometers could be covered by lava tens of

meters thick. Although the immediate area affected by the primary hazards of deep burial by

pumice, ash, or lava might be relatively small, the slow extrusion of lava would be likely to

continue for many months or even years. Structures would not be able to withstand the lava

advance, and any nearby roads or other infrastructure would certainly be affected. Advance of the

viscous rhyolitic lava would be much slower than for a basaltic eruption; most facilities could be

safely evacuated, and some of them, as well as vehicles, could be relocated during advance of a

typical flow front.

Secondary hazards from a small rhyolitic eruption would be likely to include fires started by

contact with the extruding lava. Other secondary hazards might include flooding if drainages were

temporarily blocked or subsequently released suddenly. Because of the relatively slow advance of

such lavas, however, effective mitigation of such effects might be somewhat less difficult than for a

rapidly advancing basaltic lava flow.

The probability of a small rhyolitic eruption outside the Yellowstone caldera is quite low.

Fourteen have occurred in postcaldera time (Christiansen, 2001; Spell and others, 2004) with an

average intereruptive period of about 46,000 years. During the past 170,000 years there have been

6 small extracaldera rhyolitic eruptions, equivalent to an average period of 28,000 years. Eruptions

in the past 170,000 years have been larger, typically producing flows of several km3

, in comparison

to the domes of less than 0.5 km3 extruded earlier in the postcaldera history. Insufficient age data

exist to evaluate any possible episodic behavior of these youngest eruptions. The average annual

probability of a small extracaldera rhyolitic eruption in the Yellowstone Plateau volcanic field is

small. Calculated on the basis of the entire postcaldera period, it probably is on the order of 2x10-5;

on the basis of the record of the past 170,000 years it would be about 4x10-5. Although less

probable, a similar small rhyolitic eruption might also be possible within the Yellowstone caldera

as well as in the extracaldera zone. Such extrusions, for example, formed the Douglas Knob and

Trischman Knob domes on the Madison Plateau though these might be late-stage vent domes

related to the much larger Bechler River flow.

It might be difficult in the early stages to distinguish typical premonitory seismic,

deformation, or gas-emission indicators of a small rhyolitic eruption from those for a basaltic

eruption. Because of more viscous magma, uplift and tilting of the ground surface above a zone of

intensifying swarm seismicity might likely be more pronounced for a rhyolitic eruption. Also, as

such seismic signals are typically associated with subsurface separation of a gas phase from the

magma, the time required for any relatively deep long-period seismicity that might develop could

be longer for a viscous rhyolitic magma than for basaltic magma. Detecting any such signals

might, however, require instrumentation sensitive to a broader frequency range than is currently

available at Yellowstone, such as borehole strain meters.

Large caldera-forming eruption

Of all the possible eruptive hazards that might occur in the region of Yellowstone National

Park, by far the least likely is that of another major caldera-forming pyroclastic eruption of 100 km3

or greater. Three such events have occurred in about the past 2 million years, each associated with

a cycle of precaldera and postcaldera rhyolitic volcanism lasting on the order of a million years. In

the Island Park area, west of the 639±2-ka Yellowstone caldera, the older rhyolitic source areas

have subsequently produced basaltic lava eruptions. In contrast, contemporaneous basaltic magmas

surround the Yellowstone caldera, but none have erupted within the caldera. This pattern strongly

suggests that the crust where rhyolitic magma chambers existed during the previous two major

caldera-forming eruptions and their associated rhyolitic volcanism has cooled and solidified
27

sufficiently to fracture and allow basaltic magmas to intrude from below, precluding the possibility

of large volumes of eruptible rhyolitic magma remaining there. However, the great heat flow

represented by the massive long-lived hydrothermal circulation system of Yellowstone (Fournier,

1989) as well as significant delays in seismic-wave travel times and wave attenuation imaged in the

shallow crust beneath the Yellowstone caldera (Benz and Smith, 1984; Miller and Smith, 1999;

Husen and others, 2004) strongly suggest the continued presence of magma. What remain most

uncertain are (1) the percentage of melt in the remaining, partly crystallized magma, (2) its degree

of interconnection, and (3) its potential eruptibility. The more than 600 km3

 of highly differentiated

magma that has erupted as lava flows within the caldera between ~170 and 72 ka represents a

volume equivalent to a large caldera-forming eruption. Those eruptions perhaps partly degassed

and depleted the magma sufficiently slowly without triggering voluminous pyroclastic eruptions

that they may have rendered another major caldera-forming eruption from the present subcaldera

chamber unlikely.

Despite the seeming improbability of another large caldera-forming eruption at

Yellowstone, the severe consequences of any such eruption, if one should occur, render it important

to consider its potential hazards. The distribution of basalt around the Yellowstone caldera, as

noted above, would seem to restrict the area of potential accumulation of sufficient rhyolitic

magma for any such eruption to the vicinity of the Yellowstone caldera. It might seem that

propagation of the Yellowstone hotspot toward the northeast through late Cenozoic time (compare

fig. 4) could make a future caldera-forming eruption more likely to occur northeast of the present

caldera center. Indeed, crustal density is notably low, hydrothermal heat flow high, and seismic

wave speeds slow in the area of Hot Springs Basin, ~13 km northeast of the Sour Creek dome (fig.

7). These characteristics, however, probably are related to crustal changes associated with the deep

vapor-dominated hydrothermal system of that area (Miller and Smith, 1999). The pattern of young

postcaldera rhyolitic magmatism of the Central Plateau Member suggests the possibility

(Christiansen, 2001) that, since ~170 ka, rhyolitic magma could have been rejuvenated or

remobilized beneath the western, or Mallard Lake, caldera segment in a manner analogous to the

relation of the second-cycle Henrys Fork caldera to the older first-cycle Big Bend Ridge caldera

segment (fig. 5). Although no particular site can be shown to be either more or less likely to be the

source of any possible future voluminous pyroclastic eruption, only a site within or immediately

adjacent to the present caldera seems reasonable for a large pyroclastic eruption at Yellowstone.

Spell and others (2004) and Smith and Bennett (2006) have suggested that the basaltic and

rhyolitic volcanism of the Norris-Mammoth corridor postdating the Lava Creek Tuff represents an

evolving magma chamber that could presage a new caldera-forming eruption from that area. That

scenario, however, seems to be contradicted by several lines of evidence. Geochemical and

isotopic data, especially 18O, show that all three caldera cycles of the Yellowstone Plateau volcanic

field erupted from a continuously evolving shallow-crustal magma-storage system and that the

magmas erupted in the Norris-Mammoth corridor are unrelated to them (Hildreth and others, 1984;

Hildreth and others, 1991). Furthermore, volumes of rhyolite in that corridor are small, and

basaltic magma was available for eruption continuously during eruption of those rhyolites, as

indicated by the undated Osprey Basalt that fills deep canyons cut into the youngest dated basalt

(174±46 ka) as well as abundant quenched mafic inclusions in the 80-ka Crystal Spring rhyolite

flow. Additionally, whereas geophysical indicators of inflation and deflation of the Yellowstone

caldera are commonly coherent over the caldera area, the Norris-Mammoth corridor shows no

comparable indications of systematic deformation or seismicity. These factors argue against any

large-scale rhyolitic magma-storage sufficient to promote and sustain a caldera-forming eruption in

the Norris-Mammoth corridor.

Although the probability of another major caldera-forming eruption at Yellowstone is

extremely low, the hazards from any such eruption cannot be overstated. Once any such eruption
28

began, it probably would proceed quickly to a climax, which would last for many days. Any

caldera-forming eruption probably would begin at one or a few individual sites with plinian

columns of hot volcanic ash rising into the stratosphere, and likely would spread to encompass a

ring-fracture system encircling the newly collapsing caldera. Lacking historical precedents for

such an eruption, the length of time that an initial plinian phase might last is uncertain and could be

a few hours to many weeks. Once the eruption developed past the initial phase, areas within and

adjacent to the eruptive source would be overrun quickly by lateral flows of mixed volcanic ash,

rocks, and gases at temperatures of several hundred degrees Celsius and speeds of a hundred km/h

and greater as ash fall continued both within the source area and to great distances, even around the

globe. Ash-accumulation rates estimated for the well-studied Bishop Tuff of eastern California,

roughly comparable to the Lava Creek Tuff though approximately half its volume, indicate that the

Bishop Tuff erupted continuously for several days (Wilson and Hildreth, 1997). A major calderaforming eruption at Yellowstone probably would have a similar duration. All structures and living

things within the areas overrun by ash flows would be destroyed and possibly vaporized. The

atmosphere above the eruption would be choked with pumiceous volcanic ash and lapilli,

producing an extreme hazard to aviation, and the ash and associated fallout would quickly spread

far beyond the eruptive area. Because of the complex patterns of upper-atmospheric wind

circulation, there would likely not be a single simple plume of ash fallout in one downwind

direction, but large parts of western and central North America might well be buried by ash falls,

ranging from many meters in thickness within tens of kilometers of the eruption to many

centimeters at continental distances (compare fig. 3). Fine ash and other particulates would circle

the globe.

The possible secondary hazards from a potential large caldera-forming eruption at

Yellowstone could themselves be greater than the primary hazards from other types of eruption.

Suspended ash would continue to circulate in the upper atmosphere for many weeks, and could,

together with volcanic gases associated with the eruption, affect global climate for several years.

Reduced solar irradiation resulting from suspended ash and sulfuric-acid droplets could cool the

lower atmosphere for several years, at least throughout the northern hemisphere and probably

globally as well (Rampino and others, 1988). Other global climatic effects might last even longer,

depending on the sensitivity of certain environments (Rampino, 1991), with possibly severe

ecological consequences (Rampino and Ambrose, 2000). In addition, the primary ash blanket over

large regions, especially of the Western United States, would be eroded by rainfall and surface

drainage to be redistributed from higher areas into lower-lying places, mantling some of them with

many meters of ash even where initial fall was only millimeters to centimeters thick. Many

reservoirs would be silted up, some of them becoming unusable; some dams might fail as a result of

overtopping by their reservoirs. Most of these secondary effects would be distributed over large

regions, far from Yellowstone; the devastating primary effects would far outweigh them in the

proximal area.

Although the probability of a large caldera-forming eruption at Yellowstone is exceedingly

small, it is exceedingly difficullt to make a defensible quantitative estimate of that probability. As

there have been three such eruptions in about the past 2,100,000 years, there are only two intereruptive periods from which to gauge any additional possible interval between the third and a

potential fourth such event. The first interval, between the Huckleberry Ridge (2.059±0.004 Ma)

and Mesa Falls (1.285±0.004 Ma) caldera-forming events, was 774,000±5700 years. The second

interval, between the Mesa Falls and Lava Creek (0.639±0.002 Ma) events, was 646,000±4400

years. A statement, widely repeated in popular media, regards such eruptions as occurring at

Yellowstone “every 600,000 years” with the latest eruption having been “600,000 years ago”. This

is commonly taken to imply that another such eruption is “overdue”. Such a statement is

statistically indefensible on the basis of the extrapolation of two intervals. (Even the simple
29

arithmetic average of the two intervals is 710,000 years, not 600,000 years). From the line of

reasoning outlined here, the probability of a fourth large caldera-forming event at Yellowstone can

be considered to be less than 1 in a million, below the threshold of hazards interest unless future

premonitory phenomena, probably more severe than those recorded historically in caldera systems

around the world (Newhall and Dzurisin, 1988), were to be recognized.

Premonitory indications of an impending major caldera-forming eruption at Yellowstone

would include intense swarm seismicity, perhaps localized near the site of an impending outbreak,

but initial indications might not be greatly different from those for a smaller eruption. A magma

body large enough to sustain a major caldera-forming eruption would, however, be expected

eventually to respond as a whole. By the time seismicity and ground fracturing spread to

encompass a larger area, equivalent to a potential caldera and perhaps to outline a ring-fracture

system, major eruption might already be well underway. Based upon the geologic record of the

Lava Creek eruption, magma rising to shallow levels almost certainly would produce significant

uplift of both a locus of possible outbreak and also a larger area of shallow rising magma. Quite

possibly ground fracturing would accompany intrusion to shallow crustal levels and might even

begin to outline a ring-fracture system (compare to figure 6). Magmatic gases venting to the

atmosphere before any ash or lava were to erupt, including CO2, various sulfur species (but

especially SO2), and halogens, might well be more evident and more copious from such a large

shallow magma body as it ascends into the brittle upper-crustal zone than would be expected for a

smaller body that might lead to a single central-vent rhyolitic eruption. Although many of the

specific premonitory events for such an eruption might resemble precursors to a smaller rhyolitic

eruption, the magnitudes would be expected to be correspondingly greater, there might be

geophysical activity over an areally large source, and the course of events would be expected to be

more complex and of longer duration than for a smaller eruption.

Hydrothermal-explosion hazards

The potential hazards and probabilities of future hydrothermal-explosion events in

Yellowstone National Park are considered in this section, based upon information summarized

earlier about previous such events at Yellowstone and elsewhere. (Also see appendices 1 and 2).

How often do they occur?

For several reasons, it is difficult to estimate the frequency with which hydrothermal

explosions occur at Yellowstone National Park. First, there is no clear demarcation between a

forceful geyser eruption that ejects mud and rock fragments and a more energetic explosion that

creates a new crater. Second, no comprehensive catalog of explosion events has been compiled

heretofore. Table 4 lists the results of a literature review that uncovered 26 examples of

hydrothermal explosions that have been described in the past 126 years. They refer to events in

which a new feature was created or the size of an existing thermal feature was increased

significantly by explosive excavation. There are many other examples of geysering where mud and

rock were thrown from the vent, including an eruption near Wall Pool in the Upper Geyser Basin

during the Summer of 2006, while this report was being prepared. However, the catalog in table 4

refers to events where the vent walls or conduit are shattered and fragments are dispersed outward.

Nearly all of the explosions in table 4 occurred at thermal areas close to roads, where visitors are

most frequent and the likelihood of observation of an actual explosion or of recent deposits is

highest. It is likely that similar explosions are much more frequent than tabulated in table 4. A

conservative estimate would be that at least one rock-hurling explosion occurs every two years at

Yellowstone. Because these events are typically small and can occur any time of year, including

when few visitors are present, the likelihood of harm to any individual park visitor is small. For
30

example, while the annual probability of a small hydrothermal explosion within the park, assuming

one eruption every two years, is 0.5, if visitors are generally absent from geyser basins at night and

between December and March, then the annual probability of an explosion that could cause

personal injury falls to 0.5 x 0.67 (8 months/year) x 0.58 (14 hours/24-hour day), or 0.19.

Moreover, explosions can occur away from trails and roads where people observe thermal features.

When such factors are considered, the annual probability of an injury due to a small hydrothermal

explosion falls to somewhere between 0.01 and 0.1.

Larger events clearly represent a greater threat to park visitors and infrastructure. Muffler

and others (1971) and Morgan and others (in review) identified geologic deposits formed through

hydrothermal explosions within the past 14,000 years (table 3). Within the park, there are at least

18 such features, all with diameters greater than 300 m. Two craters, Mary Bay and Turbid Lake,

have diameters greater than 1500 m.

One can generalize the probability of hydrothermal explosions by plotting explosion

magnitude (crater size) versus frequency, simplified as annualized probability as shown in table 5.

In figure 31, these data plot as a straight line on a logarithmic plot of annualized probability

versus the area of the newly formed crater (a proxy for eruption magnitude). Interpolation along

this line implies that one could expect an explosion large enough to produce a crater 100 m in

diameter every 200 years.

Such an analysis is consistent with the energetics of hydrothermal explosions. Browne and

Lawless (2001) estimated that the energy releases of even large explosions, producing craters 1000

m across, represent only a few weeks to months of the heat flow from the system. Hedenquist and

Henley (1985) produced a similar assessment for the prehistoric Waiotapu eruptions in New

Zealand (craters ranging from 60 to 250 m in diameter). They concluded that the eruption that

formed the present Champagne pool (now 70 m in diameter) released the energy equivalent of 30

hours of natural heat flux from that feature. Using such reasoning, Browne and Lawless (2001)

advised that hydrothermal explosions should be relatively common events in long-lived geothermal

areas and do not require unusual conditions.

Potential effects

As stated above, small hydrothermal eruptions creating craters a few meters across are

unlikely to cause frequent injury to park visitors or damage to infrastructure. Larger explosions,

however, are more energetic, and their effects could be correspondingly greater. We have

estimated that an explosion creating a 100-m-diameter crater might occur in Yellowstone about

once every 200 years. A series of similar-sized craters formed between one and two thousand years

ago in the Waiotapu geothermal system of New Zealand. As described by Hedenquist and Henley

(1985), deposits from these explosions extend 5 to 20 times greater than the diameters of their

source craters; thus, rock fragments might fly as far as 2000 m from a 100-m crater. Areas near the

explosion source would be buried beneath thick fragmental deposits. At the Okaro crater in New

Zealand, deposits are 2 or more meters thick within 400 m of the originally 250-m-wide crater, and

ejecta deposits extend up to 1000 m from the vent (Hedenquist and Henley, 1985). Morgan and

others (in review) note that rock fragments were ejected at least as far as 3.5 to 4 km from the Mary

Bay crater rim during its period of explosive activity, ~13,600 years ago. Fragments from the

Indian Pond explosion crater are found as far as 3 km from the vent (Morgan and others, 2003b).

Another potential hazard from hydrothermal explosions is the release of substantial amounts

of toxic gases such as H2S and CO2. The topic of toxic gas is discussed later in this hazard

assessment, but it is worth pointing out here that gas released during hydrothermal explosions at

Dieng, Indonesia, in 1979, killed 142 people. The gas was discharged from several adjacent areas

that had first ejected rocks and mud, forming several craters, the largest of which was 90 m wide

and 100 m deep. At Yellowstone, there is no evidence that any hydrothermal explosions have been
31

accompanied by significant toxic-gas release. On the other hand, the Yellowstone caldera is a

prolific source of CO2 (Werner and Brantley, 2003), and it remains possible that explosions from

vapor-dominated regions, areas without a record of historic explosions, may release significant

quantities of gas.

Precursory signals

Prior to 1989 (when it exploded), Porkchop Geyser exhibited a variety of characteristics that

could be taken to indicate a potential explosion. Its behavior changed from that of an irregular

geyser to a perpetual spouter (i.e., in constant geyser-like eruption), and the temperature of its deep

feeding waters increased (Fournier and others, 1991). With the spring emitting a roar audible from

2 km away, few visitors in the vicinity would have failed to notice its presence. Today, as then, a

geyser behaving in such a manner would be monitored closely by YVO staff and collaborators.

Other hydrothermal features in Yellowstone have experienced explosions after first displaying

similar signs of instability. After the M=7.5 1959 Hebgen Lake earthquake, park geologist George

Marler inventoried hundreds of features with altered behavior. One such feature, across the

Firehole River from Biscuit Basin, originated as a set of newly formed fractures emitting steam

with apparent superheat, at ~95°C. A fumarole exploded from this site sometime in the early

spring of 1963, forming a crater 2.7 x 4.9 m in size (Marler and White, 1975) that was subsequently

named Seismic Geyser. Currently, the feature is a non-erupting hot spring. Hydrothermal

explosions following earthquakes would be fundamentally unpredictable.

The Dieng, Indonesia, hydrothermal explosion of February 20, 1979 was preceded for

several days by shallow seismic swarms (Allard and others, 1989). During the three hours prior to

the beginning of the explosions, three earthquakes were felt by nearby villagers 3 km distant from

the explosion source. Once initiated, the explosive activity lasted for about 2 days. Other thermal

features at Dieng, as elsewhere, have exploded without conspicuous precursory activity (Allard and

others, 1989). White (1955) described an eruption at the Lake City Hot Springs northeastern

California in March 1951 that dispersed 300,000 tons of mud, greatly modifying 20 acres of hotspring area and dispersing fine dust at distances more than 6 km from the center. The activity

occurred without any warning to local residents and farmers. Hydrothermal explosions occur

nearly every year from hot springs and geothermal wellheads in the Taupo Volcanic Zone of New

Zealand. In particular, Rotorua, a developed geothermal area and tourist attraction, has produced

scores of such eruptions over the past 100 years without obvious precursors (Scott and Cody, 2000;

Browne and Lawless, 2001).

Clearly, some hydrothermal eruptions can initiate quickly, including but not limited to those

triggered by earthquakes. Others seem to follow years of anomalous activity of hot springs and

fumaroles. Because hydrothermal explosions typically appear to be shallow-rooted phenomena,

initiating close to the surface, geophysical signals are most likely to affect localized areas within

only a few hundred meters of the actual explosion site. At most volcanoes and hydrothermal

systems, including Yellowstone, the station spacing of seismometers, tilt meters and other

monitoring equipment will only rarely yield sufficient information to detect such precursory

activity. Local monitoring networks with multiple sensors at a geyser basin might be sufficient to

detect precursors such as ground tilting or subtle tremor, but as no such experiment has yet been

undertaken, this supposition remains unproven.

Where are hydrothermal explosions most likely to occur?

Both historic (table 4 and fig. 25) and prehistoric (table 3 and fig. 24) hydrothermalexplosion craters are located predominantly in the Firehole River geyser basins, in and around

Yellowstone Lake, and in the southern part of the Norris-Mammoth Corridor. The preponderance
32

of historic explosion sites within the Firehole River and Norris basins may partly reflect greater

visitation and scrutiny of thermal features in those areas but also is consistent with liquiddominated systems having near-surface temperatures near boiling being more likely to experience

explosions than vapor-dominated areas that are not continuously saturated with liquid.

The north half of Yellowstone Lake and the adjacent area has long-been known as a site of

large explosion craters and hosts an active liquid-dominated geothermal system (Muffler and

others, 1971; Morgan and others, 1977; Otis and others, 1977; Wold and others, 1977; Morgan and

others, 2003b). Besides the explosion craters listed in table 3, numerous small sublacustrine

explosion craters have been identified within linear fissure zones by high-resolution imaging

(Johnson and others, 2003; Morgan and others, 2003b; Morgan and others, in press-a). Johnson

and others (2003) and Morgan and others (2003b) considered domal areas on the lake floor to be

the most likely sites of future hydrothermal explosions. These areas are manifested by gently

warped areas of the lake floor having at most 30 meters of apparent uplift (Morgan and others, in

press-b). Hydrothermal alteration is inferred from the nature of seismic reflections; doming is

interpreted to have occurred as a result of hydrothermal pressurization beneath a sealed zone of low

permeability resulting from precipitation of siliceous sinter or silicification (Morgan and Shanks,

2005). Some of these domal areas are breached by vents, the formation of which would have

relieved any notable overpressure. Domes without any evidence of breaching were considered by

Johnson and others (2003) as the most likely to experience future explosive activity.

Figures 24 and 25 identify a significant number of historic and prehistoric explosion craters

within the acid-sulfate terrains that overlie vapor-dominated systems. As such, one cannot negate

the possibility that these areas too are susceptible to explosive activity and may present a

significant hydrothermal-explosion hazard. It remains possible, however, that many of the

observed hydrothermal-eruption craters in vapor-dominated areas, such as Roaring Mountain, Fern

Lake, Hot Springs Basin, and Sulphur Hills, were formed at a time preceding present vapordominated activity. White and others (1971) noted that the Mud Volcano region showed evidence

for having earlier been a liquid-dominated system. Many inactive sinter terraces were apparently

encroached upon by acid-sulfate mineralization, indicating a local increase in the depth to the

boiling ground water table and cessation of surface flow of silica-saturated waters. Fournier (1989)

noted similar evidence of a previous liquid-dominated system that precipitated siliceous sinter at

Hot Springs Basin, an area today characterized almost solely by acid-sulfate mineralization and

without outflow of appreciable neutral-Cl waters (Allen and Day, 1935). Plausibly, the large

explosion craters in these areas occurred prior to or during a transition from liquid- to vapordominated conditions.

Though Yellowstone’s silica-depositing geyser basins and acid-sulfate thermal areas are all

known to host hydrothermal-explosion features, the same cannot be said for calcium carbonate

travertine deposits in areas such as Mammoth Hot Springs, Terrace Spring, and Calcite Springs.

Typically, the waters issuing in these areas are 20-30°C below the local boiling temperature and,

therefore, pose less threat of destabilization. No evidence for previous explosions has been found

in these areas. The greatest hazards related to travertine-forming areas are buildup of toxic or

unhealthy levels of CO2, as discussed later, and local subsidence of the ground surface due to

formation of sinkholes in travertine areas. Unlike silica, the solubility of calcite increases with

decreasing temperature, so that migration of cold ground water, especially when it is acidic, can

create caves and sinkholes. Such features are common at Mammoth Hot Springs and pose a

continual threat to park infrastructure.

In summary, two types of thermal areas at Yellowstone could potentially host future

hydrothermal explosions. The geologic and historic records show that the thermal areas with active

discharge of neutral, high-chloride waters, especially those hosting geysers, are most likely to

experience explosions. Vapor-dominated regions covered with acid-sulfate soils may also
33

experience explosions, though perhaps less regularly than areas with abundant shallow boiling

ground water. Travertine-forming regions do not appear to experience hydrothermal explosions.

Seasonal and long-term effects on hydrothermal explosions

Fournier and others (1991) noted that the 1989 explosion of Porkchop Geyser corresponded

with the onset of a seasonal disturbance common in the Norris Geyser Basin (summarized by White

and others, 1988), typically occurring in late summer and characterized by increased heat

discharge, increased boiling, and chemical changes. The disturbance is akin to an underground

geyser eruption, where an increase in local heat and boiling migrates through the shallow

subsurface (White and others, 1988; Fournier and others, 1991). White and others (1988) speculate

that the disturbances are due to changes in the local ground water pressure, typically related to

seasonal fluctuations. The temporal correlation of the Porkchop explosion and the basin-wide

Norris disturbance suggests that hydrothermal explosions might be more likely during periods of

increased heat discharge and seasonal or longer-term fluctuations in the pressure of the ground

water table. The feedback between ground water pressure and earthquakes was recently discussed

by L. B. Christiansen and others (2005), who noted seasonal variations in seismicity beneath

Yellowstone Lake, increasing during periods of reduced lake levels. The observed drop in the level

of Yellowstone Lake in late summer and early fall lowers the hydrostatic head on sublacustrine

hydrothermal vents, apparently sufficient to cause increased release of H2S-rich gas bubbles and

entrainment of fine-grained sediments in upwelling hydrothermal fluids (Morgan and Shanks,

2005).

The potential effects of longer-term variations in climate on stability of hydrothermal

systems can be inferred from the data presented by Muffler and others (1971) and Bargar and

Fournier (1988), showing fluid-inclusion filling temperatures in secondary inclusions that were

higher than present-day hydrostatic boiling temperatures. After departure of glacial ice from the

Yellowstone region, the hydrothermal system apparently remained at least temporarily superheated

relative to the newly lowered hydrostatic pressure. Under such conditions, the Yellowstone

hydrothermal systems might have been especially susceptible to explosion-triggering mechanisms

such as earthquakes or drainage of glacial lakes. The established chronology of large

hydrothermal-explosion craters is insufficient to show that large explosions were more common

during the very latest Pleistocene, shortly following ice melting. Recent work (Pierce and others,

2002; Morgan and others, in review) suggests that most large hydrothermal explosions in

Yellowstone occurred at times that preclude a direct link to deglaciation.

Future climate variability, even in the absence of glaciation, might be expected to affect the

stability of hydrothermal reservoirs. Drops in the water level of Yellowstone Lake or reduction in

the ground water table at the Norris or Firehole River Geyser Basins could increase the likelihood

of explosive hydrothermal activity in those regions. Many vapor-dominated reservoirs display

evidence for conversion from an earlier stage as a liquid-dominated system (White and others,

1971; Moore and others, 2000). This evolution may occur due to explosive venting, reduced

recharge, or other potential mechanisms. As discussed above, the increase in boiling associated

with such a conversion could result in explosions. It is possible that long-term drought and

associated reduced recharge could induce hydrothermal explosions in reservoirs in the process of

converting from liquid-dominated to vapor-dominated conditions.

Hazard mitigation

Infrastructure: Hydrothermal explosions at Yellowstone occur solely in geothermal areas.

Foremost among these areas are the geyser basins where silica-saturated, neutral, Cl-rich waters

emerge at or close to their boiling point. The Upper and Lower Geyser Basins, Norris Geyser
34

Basin, West Thumb Geyser Basin, and the area within and adjacent to the north half of

Yellowstone Lake are the areas most likely to experience hydrothermal explosions. Park

infrastructure should be minimized in these areas, and new construction would best be concentrated

away from hydrothermal basins.

Documentation: Areas exhibiting superheat (water or fumarole temperatures above the

local boiling point) should be monitored carefully by YVO and park staff. Subsequent to large, felt

earthquakes, park staff should document changes to thermal features using a formal archival system

that can be made widely available and accessible.

Monitoring: The hazard from small hydrothermal explosions is relatively low. Small

events, usually producing craters less than two meters across, happen every few years and are only

rarely witnessed. However, events forming craters over 50 to 100 m in diameter do occur on a

societally relevant timescale and could appreciably impact park infrastructure and visitor safety.

Observational monitoring can be augmented by limited geophysical monitoring of ground

deformation and seismicity in selected geyser basins and around the northern half of Yellowstone

Lake and, perhaps, repeated geochemical sampling of hydrothermal fluids. Existing seismic and

GPS stations are placed so as to minimize the “noise” from thermal features and maximize the

signal from deep earthquakes and magmatic processes. More local monitoring designed to “listen”

to geyser basins would initially be useful as a tool to understand the dynamics of geothermal areas

on monthly to annual timescales. It is quite possible that precursors to some hydrothermal

explosions would become evident in such monitoring.

Gas-emission hazards

There has been no systematic study of the concentrations of CO2 and H2S in the air at

Yellowstone. Nevertheless, because human death or serious illness from toxic gases in the air is

exceedingly rare at Yellowstone, it is clear that gas concentrations are rarely at levels that could

cause severe illness. In general, the flat topography in most of Yellowstone’s thermal areas and the

light to moderate winds are sufficient to readily disperse hydrothermal gases. Reconnaissance

efforts to record the abundances of CO2 and H2S in thermal areas show that concentrations at waist

level are rarely above 0.1% and 2 ppm respectively, even adjacent to fumaroles (Jacob Lowenstern

and Henry Heasler, unpublished data). Much higher and even lethal concentrations occasionally

can be found in cracks, holes, and locations within 10 cm of the ground surface; such air, however,

generally is unlikely to be inhaled.

Though concentrations of toxic gases are generally low in Yellowstone, it is clear that they

can build up in valleys, caves, and tunnels and during windless conditions. The impact of gas

concentrations on humans living in the park was documented as early as 1883, when a woman was

overcome by vapors in the basement of one of the homes in Mammoth Hot Springs. The woman’s

absence was quickly noticed, and she was rushed into fresh air (Whittlesey, 1995, p. 66). Caves

and sinkholes near Mammoth Hot Springs are known to harbor relatively high concentrations of

CO2, which bubbles out of the local hot-spring waters. Poison Spring, Poison Cave, and the

Stygian Cave are all known as places where dead birds, insects, and small mammals have been

found (Whittlesey, 1995, p. 68).

“Death Gulch,” is a well-known source of CO2 in the northeast section of Yellowstone

National Park, along a short tributary of Cache Creek, 3 km above its confluence with the Lamar

River. Temperatures in this area are well below boiling, but springs yield copious bubbles of gas,

predominantly CO2, H2S, and methane (CH4) (Werner and Brantley, 2003). During his visit in

1888, USGS geologist Walter Weed found the remains of six bears, an elk, squirrels, hares and

numerous butterflies and insects in the gulch (Weed, 1889). All were concluded to have died

through asphyxiation where CO2 and H2S had accumulated to toxic levels. One of the bears had

not yet begun to decompose, and Weed concluded that the death had been very recent. Death
35

Gulch was visited in 1897 by T. A. Jaggar, who found the carcasses of seven grizzly bears and one

black bear (Traphagen, 1904). Traphagen visited Death Gulch in 1903 and recorded more than

10% CO2 in air near the bottom of the gulch and verified that gas emitted from openings in the

ground was mainly CO2 but contained over 10,000 ppm H2S. Flies placed into these openings died

within six seconds (Traphagen, 1904).

Yellowstone’s only documented human death by gas inhalation occurred in 1939 when

workers were lowered in a bucket into a 26-foot-deep pit near the Yellowstone River at Tower

Junction, dug to gather information about geological conditions relevant to nearby construction

(Whittlesey, 1995, p. 67). In trying to rescue a fellow construction worker who was overcome by

the vapors, another worker, Bill Nelson, lost consciousness and later died. Later tests revealed H2S

concentrations of 200 to 400 ppm in the pit and CO2 levels of 20%. The combination was

sufficient to cause another rescuer to lose consciousness after only a few short breaths. After this

incident, Bureau of Mines engineers also tested the air at Devils Kitchen at Mammoth Hot Springs.

Their finding of 7% CO2 spurred them to recommend closing of that and other nearby caves to

visitation (Whittlesey, 1995). The caves have remained closed since then.

A more recent episode of toxic-gas inhalation occurred in March 2004, when five bison

died along a section of the Gibbon River adjacent to the Norris Geyser Basin (Heasler and

Jaworowski, 2004). The bison were all lying in similar positions, with their legs stiffly

perpendicular to their bodies, and showed no signs of poor health prior to their deaths. Park

biologists and geologists concluded that toxic gas inhalation was the most likely cause of death

(Heasler and Jaworowski, 2004). Though the area is adjacent to thermal features rich in CO2 and

H2S, dangerous concentrations of these gases were not found in ambient air at the time of discovery

of the animals, or since then. Consideration of the previous month’s weather records caused

Heasler and Jaworowski (2004) to conclude that toxic levels of gas accumulated along the river

bottom during a brief time of cold weather and associated atmospheric inversion about one week

before the animals were found. It remains unknown whether a spike in outgassing from the

hydrothermal system may have contributed to the probable gas accumulation.

Relevant examples of toxic volcanic or hydrothermal gas hazards

As noted above, Yellowstone differs from most volcanoes in having very low emissions of

acid gases such as SO2 and HCl. A notorious hazard at some volcanic systems is buildup of CO2 in

deep volcanic lakes, which can result in catastrophic degassing of the deep waters. Such events

happened twice in Cameroon from Lakes Monoun and Nyos in 1984 and 1986, respectively. The

latter event occurred when a landslide into the lake caused dissolved CO2 to be released rapidly

from solution, causing a cloud of CO2 to flow downhill from the crater lake and killing 1700 people

and even more livestock (Sigurdsson, 1988). Following these events, engineers constructed

vertical pipes in the lakes to ensure continual degassing of the deep gas-charged waters. The

Cameroon Lakes are meromictic, meaning they are stratified into different layers that do not

normally mix, a consequence of the geometry of the lakes (deep and of small diameter), the

temperate climate, and the fact that CO2-rich water is denser than the fresh water above.

Yellowstone Lake and the other lakes of Yellowstone National Park are dimictic; their deep and

shallow waters mix twice per year, in the spring and fall (Thompson and others, 1998).

Additionally, hot springs on the floor of Yellowstone Lake tend to increase the temperatures of

bottom waters, enabling mixing as the warm waters rise. Because of the continual mixing, it is

highly unlikely that dissolved CO2 concentrations in the deep waters of Yellowstone lakes could

increase to levels sufficient to cause catastrophic degassing.

Though CO2 is unlikely to be released suddenly due to lacustrine processes, it remains

possible that dangerous amounts of this gas could be released by an explosion from the

hydrothermal system. A precedent for such an event occurred in the Dieng Plateau, a volcanic area
36

in Indonesia known for accumulations of CO2 in depressions and hollows, as well as for frequent

eruptions of mud from volcanic craters. In 1979, 134 people were killed while attempting to escape

from the area following several days of earthquakes, hydrothermal explosions, and an

accompanying mud flow that blocked access to the village (Allard and others, 1989). Another

death by CO2 poisoning at Dieng occurred in March 1992 (Global Volcanism Network, 1992).

Several deaths have been attributed to CO2 poisoning at Mammoth Mountain, a ski resort in

a volcanic region of eastern California. Carbon dioxide discharge was observed there in the early

1990s and resulted in several tree-kills that covered 70 hectares. High concentrations of CO2 have

been measured there in depressions in the snow around trees and adjacent to buildings (Sorey and

others, 1996). In 1998, a skier is believed to have died as a result of CO2 poisoning after falling

into a snow well adjacent to a U.S. Forest Service restroom (Hill, 2000). In March 2006, three

additional gas-related deaths occurred at Mammoth Mountain as members of the Mammoth

Mountain ski patrol made repairs to fencing that surrounds a near-boiling fumarole high on the

mountain (Becerra, 2006). Although the area is normally roped off during the winter because the

fumarole melts a hole in the snow—a hazard for skiers—a heavy snowfall partially obscured the

fence. During repairs, two individuals fell into the hole and quickly succumbed to the CO2 gas.

Another team member died during a subsequent rescue attempt. The tragedy was clearly

engendered by the unusually thick snow pack, but the ultimate hazard was the fumarole itself and

its CO2 emissions.

Hazard mitigation

Given the prodigious CO2 and H2S emissions at Yellowstone, the former estimated at

45,000 tons per day (Werner and Brantley, 2003), it is clear that toxic gases pose a potential hazard

to park visitors. Normally, the gases dissipate quickly into the atmosphere, and only rarely are

visitors present where gas concentrations can cause harm. Nevertheless, incidents such as those at

Dieng and Mammoth Mountain are possible at Yellowstone. Also, unless safety measures are

utilized, incidents similar to the one that caused a worker’s death at Tower Junction could occur

again. It is important that National Park staff working in suspect areas be made aware of the

hazards from toxic gases, carry appropriate equipment to detect anomalous gas concentrations, and

follow recommended protocols. Caves near thermal areas should continue to be closed to visitors,

including park staff. As is park policy, visitors in thermal areas should stay on maintained

pathways and away from thermal features and enclosed areas where dense gases can accumulate.

Because so few data on gas concentrations in thermal areas have been collected, we

recommend further reconnaissance studies, including continuous monitoring to assess the effects of

wind, air temperature, and other climatic factors on gas concentrations in and around specific

thermal areas. These types of studies would aid in determining whether occasional gas surges from

thermal features produce high concentrations of toxic gas in the surrounding air and what hazard

such events might pose to park staff and visitors. Monitoring of buildings, particularly basements,

might also be justified in areas of possible toxic-gas accumulation.

Conclusions

Yellowstone National Park is the site of a large, active, and integrated tectonic, magmatic,

and hydrothermal system. Potential hazards related to possible future activity of this system, as

evaluated in this report, vary greatly among the various possible types of future activity.

Earthquake hazards, although they are among the most commonly recurring hazards, are not

evaluated explicitly in this preliminary open-file version of the assessment. Their assessment,

however, is planned for a forthcoming, more complete version. Of the remaining hazards discussed

here, the most likely to occur are hydrothermal explosions, with an average annual probability from
37

as high as 0.5 (equivalent to an average recurrence of 2 years) for small explosions to perhaps

5x103

 (an average recurrence of 200 years) for explosions large enough to form a 100-m-diameter

crater. Hydrothermal explosions that might result in potential risks to people have a probability no

greater than 0.1 per year (an average recurrence of 10 years) for small explosions.

Potential for a volcanic eruption is much lower than that for hydrothermal explosions.

Probabilities range from about 6x10-5 per year (an average recurrence of 16,000 years) for a basaltic

eruption in the region to about 5x10-5 (average recurrence of 20,000 years) for a large rhyolitic lava

eruption within the Yellowstone caldera or 2x10-5 per year (average recurrence of 50,000 years) for

a small extracaldera rhyolitic eruption. The probability of a large caldera-forming eruption is very

much smaller than any of these but is not readily quantifiable; it probably is less than 10-6. The

aggregate annual probability of any volcanic eruption occurring from the Yellowstone magmatic

system is ~1x10-4, an average recurrence of 10,000 years.

Continued monitoring by YVO is likely to enable recognition of premonitory indications

before any volcanic eruption. There are likely to be few if any indications of an impending

hydrothermal explosion although certain conditions, such as major earthquakes or seasonal or longterm lowering of local water tables, tend to favor the occurrence of hydrothermal explosions. The

YVO monitoring system should at least provide rapid information about any such event once it

occurs. Toxic gas releases, especially of CO2 or H2S, would probably be recognized only after they

occur, mainly by indirect evidence such as dead or distressed plants or animals.

Advance preparation for responses to potential hazards from future volcanic eruptions or

hydrothermal explosions can help mitigate their effects on Yellowstone National Park and its staff

and visitors. Despite the low probability of a future volcanic eruption at Yellowstone, the

hazardous consequences possible from even a small eruption render important the preparation of

plans for appropriate responses to any such future event.

Acknowledgments

We appreciate the help of Jeff Cross and Ralph Taylor, both members of the Geyser

Observation and Study Association (GOSA), who provided documentation of some of the historic

hydrothermal explosions at Yellowstone. Nicole Nastanski and Terry Spell of the University of

Nevada, Las Vegas and John Obradovich provided unpublished 40Ar/39Ar age data for Yellowstone

volcanic rocks. Andy Calvert provided unpublished 40Ar/39Ar dates of Yellowstone rhyolites and

advised us more broadly on issues related to geochronology. Tom Olliff and John Varley of

Yellowstone National Park were supportive of this effort and helpful throughout its planning and

preparation. Kathryn Flynn assembled much of the geographic, monitoring and scientific data used

to construct the figures. Shaul Hurwitz, Bill Evans, Pat Shanks, and Deborah Bergfeld of the

USGS reviewed portions of this report relevant to hydrothermal and gas hazards. Wes Hildreth and

Dave Hill provided detailed technical reviews of the entire document.

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Mammoth Hot

Springs

Hebgen

Lake

Canyon

Village

M O N TA N A

M O N TA N A

W Y O M I N G

W Y O M I N G

I D A H O

I D A H O

MIDWAY GEYSER

BASIN

UPPER GEYSER

BASIN

West Yellowstone

Gardiner

Ashton

Jackson

Lake

PITCHSTONE

PLATEAU

CENTRAL

PLATEAU

MADISON

PLATEAU









Mammoth Hot

Springs

Hebgen

Lake

M O N TA N A

M O N TA N A

W Y O M I N G

W Y O M I N G

I D A H O

I D A H O

NORRIS GEYSER

BASIN

Yellowstone

Lake

Canyon

Village

Madison

Junction

West Yellowstone

Gardiner

Ashton

Jackson

Lake




Figure 20. Indian Pond, just north of Yellowstone Lake, is about 500 meters in maximum diameter and was 

formed about 3000 years ago by a hydrothermal explosion (Morgan and others, in review). Photo by Jim 

Peaco, 2001 (NPS stock photo).

65











77

Table 1. Ages of volcanic units related to the third cycle of the Yellowstone Plateau volcanic field.

Stratigraphic Unit Subunit Age±1 (Ma)* Source**

Third-cycle intracaldera units

Central Plateau Member Pitchstone Plateau flow 0.0791±0.0105 12

Grants Pass flow 0.072±0.003 1

Solfatara Plateau flow 0.1028±0.0076 13

Hayden Valley flow 0.102±0.004 3

West Yellowstone flow 0.1140±0.0012 12

Trischman Knob dome

Douglas Knob dome

Bechler River flow 0.116±0.002 2

Summit Lake flow 0.124±0.010 13

Tuff of Cold Mountain Creek 0.143±0.005 4

Spring Creek flow

Nez Perce Creek flow 0.1483±0.0051 13

dSpruce Creek flow

Elephant Back flow 0.153±0.002 1

Aster Creek flow 0.155±0.003 1

Buffalo Lake flow 0.160±0.003 1

Mary Lake flow 0.165±0.004 1

West Thumb flow 0.173±0.011 13

Tuff of Bluff Point 0.1731±0.0049 13

Dry Creek flow 0.166±0.009 4

Mallard Lake Member Mallard Lake flow 0.164±0.014 13

Upper Basin Member Scaup Lake flow 0.257±0.013 12

Dunraven Road flow 0.486±0.042 8

Canyon flow 0.484±0.015 7

Tuff of Sulphur Creek 0.473±0.009 9

Tuff of Uncle Toms Trail

Biscuit Basin flow 0.516±0.007 7

Third-cycle extracaldera units

Central Plateau Member Moose Falls flow 0.0806±0.0046 5

Roaring Mountain Member Crystal Spring flow 0.080±0.002 1

Obsidian Cliff flow 0.106±0.001 10

Gibbon River flow 0.118±0.010 10

Obsidian Creek Member Gibbon Hill dome 0.134±0.003 10

Roaring Mountain Member Norris Basin flow*** 0.138±0.003 1

Swan Lake Flat Basalt Panther Creek volcano 0.174±0.046 11

Snake River Group

Gerrit Basalt Hatchery Butte flow 0.390±0.038 14

Obsidian Creek Member Paintpot Hill dome 0.208±0.005 10

Osprey Basalt Lamar River flow 0.221±0.041 1

Obsidian Creek Member Landmark dome 0.226±0.006 10

Grizzly Lake complex 0.263±0.003 4
78

Gardner River complex 0.301±0.003 10

Apollinaris Spring dome 0.316±0.002 10

Willow Park dome 0.326±0.002 10

Madison River Basalt Flow 0.358±0.016 11

Roaring Mountain Member Cougar Creek dome 0.358±0.002 10

Riverside flow 0.526±0.003 10

Madison River Basalt "Vent area" flow 0.530±0.060 11

Falls River Basalt

Swan Lake Flat Basalt Tower Road shield 0.590±0.065 11

Basalt of Geode Creek

Basalt of Mariposa Lake

Third-cycle precaldera units

Undine Falls Basalt Upper flow 0.584±0.026 1

Mount Jackson Rhyolite Big Bear Lake flow

Mount Jackson Rhyolite Mount Haynes flow 0.609±0.006 1

Basalt of Warm River Warm River flow 0.759±0.052 1

Mount Jackson Rhyolite Harlequin Lake flow 0.839±0.008 2

Lewis Canyon Rhyolite Lewis Canyon flow 0.853±0.007 2

Mount Jackson Rhyolite Flat Mountain flow 0.929±0.034 2

Mount Jackson Rhyolite Wapiti Lake flow 1.17±0.01 2

Mount Jackson Rhyolite Moose Creek Butte flow 1.22±0.01 2

The Narrows Upper flow 1.30±0.35 2

*Shown in known or assumed stratigraphic order. Ages shown are selected as the “best” ages available on the basis of analytical

precision, method of analysis, and fit to stratigraphic order. Error limits shown are 1, as provided by the source, based on precision

of the analytical data. In some instances, geologic constraints demonstrate that the actual accuracy of the ages shown may reflect

errors somewhat larger than the stated analytical precision.

**Data sources: 1. K-Ar sanidine age determination (Obradovich, 1992)

2. Weighted mean of 2 K-Ar sanidine age determinations (Obradovich, 1992)

3. Weighted mean of 3 K-Ar sanidine age determinations (Obradovich, 1992)

4. 40Ar/39Ar total-fusion sanidine age determination (J. D. Obradovich, 1997, unpubl. data)

5. Weighted mean of 3 40Ar/39Ar total-fusion sanidine age determinations (J. D. Obradovich, 1997, unpubl.

data)

6. 40Ar/39Ar total-fusion sanidine age determination (Muffler and others, 1971; Gansecki and others, 1996)

7. 40Ar/39Ar total-fusion sanidine + plagioclase age determination (Gansecki and others, 1996)

8. 40Ar/39Ar total-fusion plagioclase age determination (Gansecki and others, 1996)

9. Weighted mean of 2 40Ar/39Ar total-fusion sanidine + plagioclase age determinations (Gansecki and others,

1996)

10. 40Ar/39Ar age determination (N. Nastanski and T. L. Spell, 2004, unpubl. data)

11. 40Ar/39Ar age determination (Smith and Bennett, 2006, from unpubl. data by N. Nastanski and T. L. Spell,

2004)

12. 40Ar/39Ar plateau age determination (A. T. Calvert, 2005, unpublished)

13. 40Ar/39Ar isochron age determination (A. T. Calvert, 2005, unpublished)

14. 40Ar/39Ar plateau age determination (Tauxe and others, 2004)

*** Identified in the original source as part of the Gibbon River flow.
79

 Table 2. Age, area, and volume of late postcaldera lavas of the third cycle of the Yellowstone

Plateau volcanic field.

Lava flows shown in most probable order of age, as based on data shown in figure 9. Cumulative volume of all flows

excludes the estimated volume of unobserved older flows.

*Volumes for the two pyroclastic units cannot be calculated from existing data; the volumes shown with asterisks are

rough estimates based only on the extents of known outcrop distributions and the size of the West Thumb caldera,

source of the tuff of Bluff Point, as noted in the text.

Name of flow Age ±1

(ka)

Mean Group

Age±1

(ka)

Area

(km2)

Volume

(km3)

Cum.

Volume

(km3)

Group Cum.

Volume

(km3)

Avg. Thickness

(m)

Pitchstone Plateau 79±11 351 70 366 200

Grants Pass 72±3 76±5 14 0.5 296 71 37

Solfatara Plateau 103±8 123 7 295 57

Hayden Valley 102±4 103±1 111 2 288 9 17

West Yellowstone 114±1 389 41 286 105

Trischman Knob 0.8 0.014 245 17

Douglas Knob 0.3 0.010 245 32

Bechler River 116±2 236 8 245 35

Summit Lake 124±10 118±5 402 37 237 86 92

Cold Mountain Creek 143±5 10* 200

Spring Creek 115 7 190 63

Nez Perce Creek 148±5 157 6 183 41

Spruce Creek 95 3 176 32

Elephant Back 153±2 325 25 173 77

Aster Creek 155±3 150±5 332 10 148 62 31

Buffalo Lake 160±3 552 54 138 97

Mary Lake 165±4 33 2 84 46

West Thumb 173±11 245 11 83 43

Bluff Point 173±5 50* 72

Dry Creek 166±9 178 9 22 53

Mallard Lake 164±14 225 13 13 138 56

Unobserved units 167±5 2985 256 256 86
80

othermal explosion craters in Yellowstone National Park (data from Muffler and others, 1971; Morgan and others,

Max.. length Max.. width* Diameter (avg.) Area Max. Rim height Max. water depth Total depth rim to floor Geologic Setting Age (est.)

(m) (m) (km) (km2) (m) (m) (m) (ka)

er

644 504 0.574 0.255 58 0 58

In Lava Creek Tuff (LCT), along NorrisMammoth corridor <16

ater 314 289 0.302 0.071 33 0 33

In Lava Creek Tuff, along Norris-Mammothcorridor <16

sin 114 105 0.162 0.009 16 unknown 16

In Lava Creek Tuff, along Norris-Mammothcorridor <16

rater

In Lava Creek Tuff, along Norris-Mammothcorridor

s:

758 418 0.588 0.249 41 0 41

In basin of Quaternary sediments surrounded byrhyolite lavas <16

342 240 0.291 0.064 23 unknown 23

In basin of Quaternary sediments surrounded byrhyolite lavas <16

641 624 0.633 0.314 129 unknown 129

Thermal kame, in basin surrounded by rhyolitelavas <16

733 500 0.617 0.288 35 18 53 Edge of Dry Creek flow 4-6

553 544 0.549 0.236 27 42 27 High heat flow, edge of Aster Creek flow 4-6

one Lake area:

1685 1502 1.594 1.988 85 42 127

In alluvium, along topographic marginYellowstone caldera 10.3

495 418 0.457 0.163 11 27 38

In alluvium, along Weasel Creek-Storm Pointlinear trend 3.0

2400 2824 2.612 5.323 60 53 113 High heat flow region of Yellowstone Lake 13.6

938 727 0.833 0.536 52 60 52 High heat flow, inside edge of lava flow 8.0

770 712 0.741 0.431 21 50 21

Edge Aster Creek flow, topo. marginYellowstone caldera

354 248 0.301 0.069 50 0 50

LCT, along resurgence-related fault in the SourCreek dome <16
81

1095 540 0.818 0.464 38 8 46

Edge Upper Basin Member near topo. marginYellowstone caldera <16

289 268 0.279 0.243 39 0 39

LCT and thermal kame, along topo. marginYellowstone caldera <16

397 341 0.369 0.106 35 0 35

Thermal kame, edge Canyon Flow & topo.margin Yell. caldera <16

red perpendicular to maximum length.
82

ric hydrothermal explosion craters in Yellowstone National Park.

Thermal area

Maximum

length

(m)

Maximum

width

(m)

Age

(estimated)

Reference

Biscuit Basin n.d. n.d. 1957-1958 (Marler and White, 1975)

Biscuit Basin 60 15 1918, 1925,

1934, 1953

(Marler and White, 1975; Paperiello, 1998)

Biscuit Basin 10 6 1959 (Marler and White, 1975)

Biscuit Basin n.d. n.d. 1963 (Marler and White, 1975)

Cascade Group 12 1959-1971 (Marler and White, 1975)

Geyser Hill n.d. n.d. 1969 (Marler and White, 1975)

Geyser Hill n.d. n.d. 1922, 1972 (Marler, 1973)

Basins

Fairy Group 3 3 1980? (Bryan, 1995)

Fountain Group 1.5 1.2 1969 (Marler and White, 1975)

Fountain Lake n.d. n.d. 1959 (Marler and White, 1975)

Kaleidoscope Group 4.5 4.5 1959 (Marler, 1973)

ser Kaleidoscope Group 3 3 1963 (Marler and White, 1975)

Kaleidoscope Group 3.7 1.8 1960 (Marler and White, 1975)

Midway Geyser Basin 10.5 9 btw 1904-1940 (Marler and White, 1975)

Midway Geyser Basin 107 55 1881, 1882,

1888, 1890

(Whittlesey, 1990)

Norris Geyser Basin 9 2 1971 (White and others, 1988)

Norris Geyser Basin 75 21 2003 http://volcanoes.usgs.gov/yvo/2003/NorrisTherm03.html

Norris Geyser Basin 13.9 11.7 1989 (Fournier and others, 1991)

Unnamed 16 11 1986, 1987 (Hutchinson, 1987; Hobart, 1989)

Thermal area west of Elk

Park

n.d. n.d. 1988 (Paperiello, 1998)

Mushpots 1.9 1.4 1985 (Hutchinson, 1990)

e

Divide Group 1 1 1999 or 2000 Jeff Cross (personal commun., Feb. 2004)

Shoshone Geyser Basin 5 5 1929 (Davis, 1930)

Potts Basin n.d. n.d. 1965, 1982 (Taylor and Grigg, 1999)

Potts Basin n.d. n.d. 1998 (Taylor and Grigg, 1999)

Roaring Mountain n.d. n.d. 1922 (Whittlesey, 1988)
83

Table 5. Estimates of annualized probability of events greater than a given magnitude.

Diameter (m) Area (m2) Events in last 14

thousand years

Annualized

Probability

>2 3.1 7000 (estimated) 0.50

>300 70,700 16 0.0013

>2000 3,140,000 2 0.00014
84

Table 6. Ages of late intracaldera lavas and tuffs. Group ages for episodes calculated from the

means and standard deviations of the nominal ages. Model I arbitrarily spreads individual events

for each episode over 1000-year time intervals. Model II spreads individual events over plus and

minus one standard deviation for the three oldest episodes and, for the two youngest episodes

(each with only two events), at the event age and one standard deviation. Model III is an estimated

chronology based as closely as possible on the nominal ages and analytical standard deviations of

the individual events.

No. Name Ages and sd (ka)

Group mean age

and sd (ka)

Model I age

(ka)

Model II age

(ka)

Model III age

(ka)

1 Pitchstone Plateau 79±11 72 72 72

2 Grants Pass 72±3

76±5

73 76 73

3 Solfatara Plateau 103±8 102 102 102

4 Hayden Valley 102±4

103±1

103 107 103

5 West Yellowstone 114±1 113.5 112 114

6 Trischman Knob 113.8 113 116

7 Douglas Knob 114 114 116.2

8 Bechler River 116±2 114.3 115 116.4

9 Summit Lake 124±10

118±5

114.5 116 124

10 Cold Mountain Creek 143±5 151.5 149 143

11 Spring Creek 151.7 150 145

12 Nez Perce Creek 148±5 151.9 152 148

13 Spruce Creek 152.1 153 150

14 Elephant Back 153±2 152.3 154 153

15 Aster Creek 155±3

150±5

152.5 155 155

16 Buffalo Lake 160±3 163.5 159 160

17 Mary Lake 165±4 163.7 161 165

18 West Thumb 173±11 163.9 163 172

19 Bluff Point 173±5 164.1 165 173

20 Dry Creek 166±9 164.3 167 174

21 Mallard Lake 164±14

167±5

164.5 169 175

22 Unobserved units 165 171 176
85

Table 7. Conditional probabilities for an eruption in the next year for Models I, II, and III for the

mixed-exponential and Weibull distributions.

Mixed-exponential distribution

Time since last

eruption (y) 0 1000 20000 72000

Model

I 2.5x10-3 5.6x10-4 4.6x10-5 4.6x10-5

II 4.3x10-4 4.1x10-4 3.5x10-5 3.4x10-5

III 3.5x10-4 3.3x10-4 5.3x10-5 5.1x10-5

Weibull distribution

Time since last

eruption (y) 0 1000 20000 72000

Model

I 2.5x10-2 3.5x10-4 7.5x10-5 3.9x10-5

II 3.3x10-4 2.4x10-4 2.2x10-4 2.1x10-4

III 4.7x10-4 2.4x10-4 1.9x10-4 1.7x10-4
86

Appendices

Appendix 1. Description of representative historic hydrothermal

explosions

Porkchop Spring/Geyser

The most carefully documented hydrothermal explosion at Yellowstone occurred in 1989 at

Porkchop Geyser in Norris Geyser Basin. The following description is summarized from Fournier

and others (1991).

Prior to 1985, Porkchop Geyser was a relatively obscure 3-m-long pool with a triangular

orifice at the surface, a few cm on each side. In 1985, a change occurred wherein the pool drained

and was replaced by a 6- to 9-m-high “perpetual spouter” with continual pulsing of fluid. For four

years, such behavior continued, with sufficient discharge to produce a roar audible up to 2 km away

from the vent. At 14:40 MDT on September 5, 1989, eight park visitors witnessed a tripling of the

eruption column height, immediately followed by explosive disruption and ejection of the sinter

sheet surrounding the pool; the explosion was over in seconds. It resulted in formation of an 11-mwide pool surrounded by an ejecta ring of sinter that was between 1 and 1.4 meters high. Material

was ejected as far away from the pool as 66 meters to the south. The debris was not dispersed

evenly as materials traveled only about 25 meters eastward. The maximum diameter of the largest

ejected block measured 1.9 meters. Fournier and others (1991) interpreted breccia blocks in the

ejected sinter as evidence of earlier explosions. In addition, a fluted fragment appeared to be part of

the throat (or conduit) of the geyser. The throat was at least 10 cm in diameter, considerably

greater than the size of the orifice through which Porkchop had been venting to the surface.

Excelsior Geyser

The most witnessed, and most impressive of all historic hydrothermal explosions at

Yellowstone took place over about ten years between 1881 and 1890 at Excelsior Geyser in the

Midway Geyser Basin. These explosions highlight the continuum between geyser eruptions and

hydrothermal explosions. In essence, Excelsior exhibited geysering that was sufficiently intense

that rock fragments were commonly ejected and the size of the crater was markedly enlarged. One

hundred years after the cessation of its phenomenal geysering, Lee Whittlesey, the Yellowstone

National Park historian and archivist summarized the eyewitness accounts of the Excelsior

eruptions (Whittlesey, 1990). The following material summarizes that information.

Excelsior was already a large and active 60 x 100 meter spring when the park was

discovered in 1872. Though eruptive activity may have started earlier, it was first described in

1881, jetting to heights of 30 to 90 meters and, according to Superintendent P. W. Norris,

sufficiently exceeding the activity of any other geyser that it merited the title of Excelsior. Activity

continued into 1882, with occasional eruptions over 110 meters in height. Whittlesey (1990)

quotes tour guide Nestor Henderson:

“The close of each eruption was accompanied by violent earthquake shocks that tore down the

geyserite walls and added much both to the danger and [to the] sublimity of the spectacle.

These masses of broken wall were at each eruption hurled into the air several hundred feet

above the topmost waves, clashing together in their descent into the yawning abyss with a

deafening noise that was most terrific.”
87

By the Fall of 1882, Excelsior became dormant. It may have erupted between 1883 and

1888, but is clearly documented to have returned to activity by May 1, 1888, when blocks 0.3 m in

diameter were hurled over 150 m from the crater. Eruptive activity continued throughout the

summer and fall. The last known eruption is believed to have been in 1890, until Excelsior erupted

again for two days in 1985, over one hundred years after its first described eruptions.

Water discharge from Excelsior’s larger eruptions is said to have doubled the outflow of the

Firehole River. The plume from Excelsior was visible at distances up to 160 km because the steam

from its eruptions would create its own clouds.

West Nymph Creek Thermal Area

The West Nymph Creek Thermal Area is an informal name for the unnamed backcountry

thermal area approximately 3 km north-northwest of the Norris Geyser Basin. The area was the

site of a series of hydrothermal explosions during 1987. In June, some time after activity had

started, two park employees witnessed a vent ejecting mud as high as six meters and as far away as

ten meters from the crater rim. As discussed in Hutchinson (1987), bursts from the pool were

“churning freshly-killed tree trunks as if they were oversized swizzle sticks.” The crater size was

9.1 x 10.5 m, with a depth as great as 3.45 m. This crater was 27 meters downslope from another

crater inferred to have formed on January 26, 1986 as an effect of the autumn 1985 seismic swarm,

in the northwest part of the Yellowstone caldera. Several other older but similar craters are found

within this thermal area. A report from a snow-coach driver combined with on-site observations

led the park geologist to conclude that the 1986 event was over in a matter of a few hours.

Black Opal/Wall Pool and Sapphire Pool

Explosions that created the present Black Opal Pool and Wall Pool in the Upper Geyser

Basin may have begun as early as 1902, but are clearly documented in 1918, 1925, 1934 and 1953.

By 1934, the opening of what is today “Wall Pool” was essentially complete. Subsequent activity

was entirely at Black Opal Pool. In January of 1934, a violent explosion created a pool 12 m in

diameter that had not existed a week before. Several tons of rock were displaced, and some 300 kg

rocks had been transported 9 m (Paperiello, 1998).

As noted above, explosive activity occurred in other years as well, and into the 1950s.

After the August 17, 1959 Hebgen Lake earthquake, Black Opal Pool remained quiet, but nearby

Sapphire Pool exhibited impressive changes. For about three weeks after the earthquake, it surged

constantly to a height of 3-4 meters. On September 5, massive bursts up to ~ 50m from an area

about 60 m across began a period of activity that washed away the famous silica biscuits for which

Biscuit Basin was named (Marler, 1964).

Historic hydrothermal explosions elsewhere

On April 19, 2005, a 50m-wide crater resulted from a hydrothermal eruption in a remote

area of New Zealand between Taupo and Rotorua. Mud and water were dispersed 70-100 meters

away from the site, and the eruptive column, which reached 200 meters into the air, was visible

from 10 km away. The eruption was one of many that have occurred over the past 40 years in New

Zealand. In 2000-2005 alone, eight different events were reported for New Zealand in the Bulletin

of the Global Volcanism Network (http://www.volcano.si.edu). Scott and Cody (2000) reported

that 91 hydrothermal explosions of natural features had occurred in Rotorua city between 1845 and

1998, with the majority occurring at times of greatest human impact on the local hydrology and

geothermal systems.

In October 1990, a thermal area near El Barro, El Salvador exploded without apparent

warning, killing 25 people and injuring 15. The area was within a small community adjacent to the
88

producing Ahuachapán geothermal field. The eruption lasted a few minutes, created a crater 40 m

in diameter and dispersed rock fragments over a 200-m-diameter area (Handal and Barrios, 2004).

Historical and geological studies indicate that the area had experienced previous hydrothermal

explosions.

Appendix 2. Description of large prehistoric hydrothermal eruption sites

at Yellowstone

Pocket Basin

Muffler and others (1971) were the first to describe hydrothermal-explosion craters in

Yellowstone National Park, focusing primarily on Pocket Basin, a late Pleistocene ring of

fragmental deposits within the Lower Geyser Basin. Their study describes Pocket Basin as a flatfloored basin surrounded by an elliptical ridge of rock 360 x 790 m in minimum and maximum

diameters. The height of the ridge above the basin floor ranges from 4 to 20 m, with the inner

slopes as steep as 20 to 25° and the shallower outer slopes less than 10°. The ridge is composed of

unconsolidated and unsorted materials that are predominantly angular fragments of yellow-stained

sandstone, siltstone and conglomerates. Most blocks in the deposit are less than 0.3 m in diameter,

though a few range up to 2.5 m. The deposits are apparently only from shallow sources as only

clasts from glacial deposits are present, with none of the underlying rhyolitic lava that was

encountered at depths of 40 meters in a nearby drillhole (Y3). Ejecta fragments are found as far as

1200 meters from the center of Pocket Basin, and thus much farther than the actual berm of

eruption deposits. The presence of glacial ice in the adjacent basin was inferred by Muffler and

others (1971) as an explanation for incision of the Firehole River through the Pocket Basin deposits

rather than adjacent lowlands. This helps constrain the time of eruption to that of the latest

glaciation, which ended ~16 ka. These authors speculated that catastrophic release of water from a

glacial lake overlying areas within the Lower Geyser Basin may have caused an abrupt pressure

decrease within the geothermal system, thus triggering the eruption.

Mary Bay

In the northern basin of Yellowstone Lake, Mary Bay contains an approximately 2.4-km by

2.8-km area of coalesced explosion craters (Wold and others, 1977), making it the world’s largest

known hydrothermal-explosion crater (Browne and Lawless, 2001). The Mary Bay hydrothermalexplosion crater extends from the northern basin of Yellowstone Lake onto land and into the lower

Pelican Valley where steep (~35o

 slope) cliffs of explosion breccia are exposed. In the lake, a welldefined crater rim on the lake floor rises about 10 m above the flat-bottomed crater; the total

elevation difference between the flat-bottomed main crater floor to the top of the outer, subaerial

crater rim is as much as 70 m (Morgan and others, 2003b; Morgan and others, in review). About

100 individual hydrothermal-vent structures are located within the larger main crater; individual

vents are as deep as 35 m below the main crater floor, making the total elevation difference from

bottom of an individual vent to the top of the rim close to 105 m. Investigations using a

submersible Remotely Operated Vehicle (ROV) show that fluids from a narrow, 53-m-deep

hydrothermal vent in Mary Bay have temperatures near the 120°C limit of the temperature probes

used, reflecting extremely high heat-flow values in this area (Morgan and others, 1977). Charcoal

fragments collected from the soil layer below the Mary Bay ejecta deposits as well as from within

the hydrothermal explosion deposit have an average calibrated 14C age of 13.6 ka BP (Pierce and

others, 2002).
89

Elliott’s crater

Within Yellowstone Lake, 1 kilometer southwest of the Mary Bay crater complex is a large

(~800-m-diameter) composite depression informally referred to as Elliott’s crater (Morgan and

others, 2003b), named after Henry Elliott, who mapped Yellowstone Lake during the 1871 Hayden

survey. The northern rim of the crater rises about 40 m above the flat-bottomed main crater floor,

whereas the southern crater rim is about 30 m above the crater floor. About 10 smaller individual

craters are present on the main crater floor and along the southwest rim; several of these are active

and have fluid temperatures ranging from 51oC to as high as 91oC. The presence of two younger

craters at the south end of the main crater floor and the presence of many smaller hydrothermal

vents further indicates more recent hydrothermal activity and possibly younger explosions. Along

the southeast and eastern rim outside of the main crater, more than 3 km south from shore, rocks of

varied lithologic compositions lie above sediments on the lake floor; these rocks may be products

of more recent hydrothermal explosions.

Stratigraphic relations indicate that the main hydrothermal explosion occurred between 13

and 8 ka, based on sedimentation rates in the lake (Johnson and others, 2003). Seismic reflection

techniques identify “opaque areas” within the sedimentary fill that are interpreted to indicate the

presence of hydrothermal fluids and (or) gases (Johnson and others, 2003; Morgan and others,

2003b).

Evil Twin explosion crater

Morgan and others (2003b) described a 500-m-diameter, sublacustrine explosion crater,

referred to informally as the Evil Twin explosion crater, in the western part of West Thumb basin

near the currently active West Thumb Geyser Basin and 300 m northeast of Duck Lake, a

postglacial (subaerial) hydrothermal-explosion crater just outside Yellowstone Lake (Muffler and

others, 1971). The 500-m-wide Evil Twin explosion crater is surrounded by 12- to 20-m-high,

nearly vertical, walls and has several smaller nested craters along its eastern edge. These nested

craters are as deep as 40 m and are younger than the main crater. Temperatures of hydrothermal

fluids emanating from the smaller northeastern nested crater have been measured at 72°C by ROV

(Morgan and others, in review).

Frank Island explosion crater

An oval, steep-walled, flat-bottomed hydrothermal-explosion crater more than 700 m wide

is located south of Frank Island (Morgan and others, 2003b). Previous lower-resolution seismic

reflection profiles interpreted this structure as a topographic edge to the Yellowstone caldera (Otis

and others, 1977); however, higher-resolution swath sonar and shallower seismic-reflection profiles

indicate that this structure, while on the slumped margin of the Yellowstone caldera, has

characteristics very similar to those of other hydrothermal explosion craters in Yellowstone Lake

(Morgan and others, 2003b). Muted topography suggests that this explosion crater is one of the

oldest still recognizable in Yellowstone Lake. Further, this crater occurs in an area where heat-flow

values are, at present, relatively low. Submersible ROV investigations do not indicate hydrothermal

activity within the crater (Morgan and others, in review).

Indian Pond

Indian Pond, formerly referred to as Squaw Lake (Muffler and others, 1971), is an oval,

500-m-wide, lake-filled crater (Table 3; Figure 20). The pond is rimmed by an apron of explosion

breccia that rises about 11 m above present-day lake level. Low-resolution bathymetric surveys of

Indian Pond (Unpubl. Data, Yellowstone National Park, 1966) indicate the pond has steep inward-
90

dipping slopes around its perimeter and has at least three elongate, east-west-trending basins on the

crater floor. The average depth of the crater floor is about 13 m below lake level; depths of the

individual basins or smaller craters range from 20 to 27 m below lake level. The lake is breached

on its southwest edge where Little Indian Creek flows out of Indian Pond into northern

Yellowstone Lake.

The Indian Pond explosion breccia is a poorly sorted, matrix-supported breccia deposit; the

matrix is light medium-brown clay that has a pervasive orange-yellow stain where groundwater has

flowed. Lithic clasts in the breccia are generally angular to sub-angular and are composed

primarily of cemented beach gravels with subordinate angular clasts of silicified lake sediments

with hydrothermal clinoptilolite and well-sorted, fine-grained indurated golden-tan siltstone.

Maximum diameter lithic fragments of cemented beach gravels and sands are up to 1.5 m at the

crater rim; in the wave-cut terrace exposures along Mary Bay, lithic fragments have maximum

diameters of 30 cm and average around 3 to 7 cm. The thickness of the Indian Pond breccia in the

wave cut terraces along Mary Bay is less than one meter, although northwest of the crater,

thicknesses are estimated to be as much as several meters.

Turbid Lake

Turbid Lake, along the eastern edge of the topographic margin of the Yellowstone caldera

was recognized by Muffler and others (1971) as a hydrothermal explosion crater. It is the second

largest hydrothermal explosion crater in Yellowstone, having a maximum crater diameter of 1685

meters and an area close to 2.0 km2 (Morgan and others, in review). A large primary crater is

located in the main central portion of the lake and is rimmed by an apron of explosion breccia that

rises about 33 m above present-day lake level on its northern, western, and southern shores. Along

its eastern edge, evidence for a smaller second crater is present where a north-south-trending ridge

of explosion breccia is deposited inside the main crater wall. The height of the eastern main crater

wall rises about 85 m above present-day level of Turbid Lake. Low-resolution bathymetric surveys

of Turbid Lake (Unpubl. data, Yellowstone National Park, 1977) indicate that the lake has steep

inward-dipping slopes around its central deep crater and has multiple smaller craters around the

periphery of the deep crater. Maximum depth of the central deep crater is 42 m below lake level;

depths of the smaller craters range from 17 m to 5 m along the northern and western edges to as

deep as 27 m along the eastern edge (table 3, Morgan and others, in review). The crater rim is

breached on the north by the Sedge Creek inlet and on the south by the inlet of Bear Creek. At its

western crater rim, Sedge Creek flows out of Turbid Lake and into northern Yellowstone Lake.

Acidic hot springs are present along the eastern and southeastern rim; a thermal area is present

along the southern rim at lake level. Mud pots there have pH of 1.1-1.5 and temperatures of 49-

57°

C.

The hydrothermal breccia is exposed along the creeks, varying in thickness from 2 to 10 m.

The matrix-supported breccia is composed of angular fragments of hydrothermally altered Lava

Creek Tuff, cemented gravels and sands, moderately and poorly sorted clastic sulfidic sandstones,

cemented pebble conglomerate, and chert breccias (Morgan and others, in review). The matrix is

generally whitish fine-grained clay. As noted by Muffler and others (1971), a broad constructional

outer ramp extends from Turbid Lake toward the northwest and may be indicative of the primary

flow direction of the explosion deposit. The explosion breccia may be distributed toward the

northwest and west as much as 4.5 km from its source. Charcoal fragments from within the

hydrothermal explosion deposit from the southern rim have a radiometric age of 9.4 ka (Pierce and

others, 2002).
91

Appendix 3. Probabilities of episodic volcanic eruptions and application

to the young intracaldera volcanic history of Yellowstone

By Manuel Nathenson and Robert L. Christiansen

An underlying assumption of USGS volcano hazards assessments for Cascade volcanoes in

Oregon and Washington has been that the probability distribution of volcanic eruptions may be

treated as a Poisson process. Time histories for some volcanoes match this assumption well (e.g.

Klein, 1982). The probability of an eruption during any particular period of time is calculated from

the relation for the occurrence rate. For a Poisson process, this relation is obtained from the

exponential distribution for the probability P{ T  t} that an eruption will occur in a time T less

than or equal to the time period t:

P{ T  t} = F(t) = 1 - e-μ t

 μt , for μt small,

where F(t) is the probability distribution function, and μ is the mean occurrence rate (events

per year) for the exponential distribution. Investigators such as Scott and others (1995) cite low

occurrence rates in the Cascades as justification for use of this relation there.

Given a set of n eruption time intervals ti

, the average recurrence interval (the reciprocal of

the occurrence rate) may be estimated by:

1

μ =

1

n 

i=1

n

t i 

The properties of a Poisson process include the characteristic that the conditional

probability of waiting a time until an eruption occurs does not depend on the time already waited

but only on the time period selected (e.g. 1 year, 30 years, etc.) to calculate a conditional

probability. For some volcanoes, the time history contains disparate time intervals between

eruptions, some being short and others much longer. Some examples of time histories having such

disparate eruption-time intervals are those of Mount Rainier and Mount St. Helens in Washington.

Mullineaux’s (1974) data for tephra layers at Mount Rainier include three long intervals (>2000

years) and seven short intervals (<600years) between eruptions. Mullineaux's (1996) data for

Mount St. Helens include one interval of 8600 years, one of 1500 years, and 34 less than 640 years.

In such instances, other probability distributions more accurately represent the data, the conditional

probabilities based on these distributions depending on the time since the last eruption.

Bebbington and Lai (1996) proposed using the Weibull distribution to model eruption times

that vary with the preceding time interval:

P{ T  t} = F(t) = 1- e-μ(t)

where μ(t) = (t

 )



 

and T is the time, less than the time period t, when an eruption will occur. Parameters  and

 are referred to as the scale and shape parameters, respectively; when  = 1, this reduces to the

exponential distribution.

For eruption intervals that can be divided into two populations, one of short intervals and

one of long intervals, a relevant model is the mixed exponential (Cox and Lewis, 1966; Nathenson,

2001):

P{ T  t} = F(t) = 1- p1 e-μ1 t - p2 e-μ2 t

where p1 = 

n1

n1+ n2
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and 1

μ1

 = 

1

n1



i=1

n1

t i

where p1 is the fraction of short intervals, μ1 is the mean occurrence rate for the short

intervals, n1 is the number of short intervals, and p2, μ2, and n2 are equivalent parameters for the

long intervals. The basic notion embodied in the mixed exponential is that there are two states.

The probability of an eruption occurring in each of these states is governed by an exponential

distribution. If one knew that the volcano were currently in a particular state (a difficult judgment

to make), then the probability of an eruption could be calculated using the appropriate simple

exponential relation for that state only.

Of direct interest for application here is the conditional probability P{t  T  t + t | T >

t} of an eruption occurring between time t and time t + t, (e.g. the next year or the next 30

years), after already waiting a time t since the last eruption. This conditional probability can be

calculated from the distribution function F(t) as

P{t  T  t + t | T > t} = 1 - 

1- F(t+ t)

1- F(t)

For the simple exponential distribution, the conditional probability reduces to:

P{t  T  t + t | T > t} = 1 - e-μ t

Thus, for the simple exponential distribution, the passage of past time does not change the

probability of the time to a future eruption. (In the engineering language of time to failure, there is

no wear or fatigue). For the Weibull distribution, the conditional probability is:

P{t  T  t + t | T > t} = 1 – exp {(

t

 )



- (

t + t

 )



}

For the mixed exponential, the conditional probability is:

P{t  T  t + t | T > t} = 1 – [p1 e-μ1 (t + t) + p2 e-μ2 (t + t) ] /

[p1 e-μ1t + p2 e-μ2t ]

Thus, unlike for the simple exponential distribution, the conditional probability for the

mixed exponential does depend on the time since the last eruption, t.

As discussed in the section on hazards from large rhyolitic lava eruptions, the chronology of

these eruptions is problematic. Table 6 reproduces the ages for these events as given in tables 1

and 2. Some of these rhyolites yield nominal ages that are younger that those of rhyolites that are

stratigraphically above them, but all such discrepancies with stratigraphic order remain within the

range of analytical uncertainty, reported as plus or minus one standard deviation. The combined

record of mapped stratigraphy and isotopic ages suggests that these rhyolites represent a series of

episodes. The times between episodes are tens of thousands of years, but the times between events

within each episode are short—possibly tens, hundreds or a few thousand years. With available

isotopic-age data, it is not possible to discern the actual time intervals between eruptions within an

episode, and there could be intervals in one episode that are a few tens of years and in another

episode that are hundreds or thousands of years. We use the available data to calculate both a mean

age for each episode, weighted by the analytical uncertainty for each event within the episode, and

a standard deviation for the age of the episode (table 6). The mean recurrence interval for the five

episodes is 23 ky, with a range from 12 ky to 38 ky; treating each episode as an independent event,

the probability in one year of a new episode occurring is 4.3x10-5.

To evaluate the variation in eruption probabilities resulting from the uncertainty of the

various time intervals between eruptions within an episode, we assess three models. Model I

assumes that each episode lasts a thousand years; the assumed individual eruption ages are spread

evenly within that thousand years for each episode, as listed in table 6. Model II assumes that the
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calculated standard deviation for the mean age of each episode represents a measure, not only of

the analytical uncertainty, but also the actual duration of the episode, spreading the individual

events evenly over the calculated standard deviation. For the two youngest episodes, each with

only two events, the age of one event in this model is taken to be the mean age, and the other is

taken to be the mean age plus one standard deviation; for the other three episodes, the ages are

spread over a time equal to plus and minus one standard deviation. Model III is an estimated

chronology based as closely as possible on the reported ages and standard deviations for each

eruptive event. When the nominal ages are in correct stratigraphic order, they are assigned directly.

Other ages are assigned to be as close as possible to the nominal ages and both to be within the

reported standard deviation and to comply with the proposed stratigraphic order. (For a few events,

where stratigraphic order is ambiguous, as illustrated in figure 9, the order for model III is chosen

rather than definitively established).

The three models tabulated in table 6 are illustrated in figure 27. Model I, with its short

duration for each episode, readily resolves graphically into the five episodes discussed here. Model

II, with its longer duration of each episode, still resolves graphically into the episodes but less

clearly than model I. The episodes are not well resolved graphically in model III. Some of the

episodes correspond graphically more to a change in rate of occurrence than to distinctly different

times. All three models, however, preserve the observation that there are long periods with no

intracaldera rhyolitic eruptive activity.

Probability distributions calculated for the three models (fig. 28) all show that most

intervals are grouped at relatively short times and only a few at longer or much longer times.

Model I (fig. 28A) has a large number of very short time intervals between eruptions and four

longer intervals. Model II (fig. 28B) spreads the eruption time intervals more uniformly, with only

two long intervals. Model III (fig. 28C) spreads the time intervals more uniformly than either of

the other two models. The three analytical distribution functions are also shown for each model.

As expected, because of the disparate time intervals, the simple exponential distribution is a poor fit

to the data in models I and II, but it is a reasonable fit to the model III data. The Weibull

distribution is a reasonably good fit to model I but a poor fit to both models II and III for the longinterval data. The mixed exponential distribution provides a good fit to both the short-interval and

the long-interval data for each model.

The conditional probability of an eruption occurring in the next year is shown in figure 29

for each of the three models. Some representative values are listed in table 7, including some of the

probabilities at short times after an eruption that are off the scale of the plots in figure 29. The

conditional probability based on the simple exponential distribution does not depend on the time

since the last eruption. Because each of the three models has a somewhat different total duration,

(table 6), the probabilities based on the exponential distribution vary slightly among the three parts

of figure 29. The conditional probabilities based on the Weibull distribution for long times since

the last eruption are considerably different for model I than for models II and III (fig. 29 and table

7). The mixed exponential distribution yields similar results for long times since the last eruption

in all three models. Because of the poor match of the Weibull distribution for long time intervals

with models II and III (fig. 28) and the variable results for the conditional probability at long times

(fig. 29 and table 7), the mixed exponential distribution is a better choice for the calculation of

conditional probabilities for this data set.

The greatest differences for conditional probabilities among the three models for the mixed

exponential distribution are at short times. For the probability of an eruption in the next year

immediately following an eruption, the mixed exponential distribution yields a probability almost

ten times higher for model I than for models II and III (table 7). This result accords with the

intuitive recognition that the many short time intervals in model I would lead one to anticipate

another eruption shortly after one had occurred. At the present time, a value for the probability of
94

an intracaldera rhyolitic eruption within the next year of 5x10-5 reasonably represents all three

models based on a mixed exponential distribution. This value is quite close to the value of 4x10-5

calculated independently for a new episode. Because both calculations are model dependent, their

very similar results are reassuring; it has been ~72 ky since the last eruption, which is also the time

since the last episode. The major differences among the models are in the estimation of the

probabilities of another eruption to follow if an eruption were to occur now (table 7). The

limitations of the available data preclude a definitive resolution of those differences and heighten

the uncertainty associated with their use in estimating eruption probabilities.

In addition to the probability of a future intracaldera rhyolitic eruption occurring within any

given time period, an important consideration would be its likely size. Figure 30 illustrates the

probability of a future lava eruption within the caldera having a volume greater that any particular

value, based upon volumes estimated for 20 of the known late intracaldera eruptions (91 % of the

known events), excluding the two pyroclastic eruptions, whose volumes are indeterminate (table 3).

The data are distributed logarithmically with volume over most of the range. The probability that a

future intracaldera lava flow would have a volume greater than 2 km3 is about 80%; the probability

of a volume greater than 10 km3 is about 40%. Thus, the probability of a large-volume future lava

flow in the caldera is greater than the probability of a small one.