U.S. Department of the Interior

U.S. Geological Survey

History of Surface Displacements at the Yellowstone 

Caldera, Wyoming, from Leveling Surveys and InSAR 

Observations, 1923–2008

Professional Paper 1788
COVER:

Hydrothermal features in the vicinity of Mushroom Geyser in the Back Basin of Norris Geyser Basin, Yellowstone National Park. Norris Geyser Basin is the 

hottest and most changeable thermal area in Yellowstone. One of three deformation centers discussed in this report is located along the north rim of the Yellowstone caldera near Norris Geyser Basin. Pressure changes within Yellowstone’s hydrothermal system likely play a role in causing episodic ground-surface 

deformation, together with inflation/deflation cycles of the underlying basalt-rhyolite magmatic system. (U.S. Geological Survey photograph by Dan Dzurisin.)
History of Surface Displacements at the 

Yellowstone Caldera, Wyoming, from 

Leveling Surveys and InSAR Observations, 

1923–2008

By Daniel Dzurisin, Charles W. Wicks, and Michael P. Poland

Professional Paper 1788

U.S. Department of the Interior

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

KEN SALAZAR, Secretary

U.S. Geological Survey

Marcia K. McNutt, Director

U.S. Geological Survey, Reston, Virginia: 2012

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Suggested citation:

Dzurisin, D., Wicks, C.W, and Poland, M.P., 2012, History of surface displacements at the Yellowstone Caldera, 

Wyoming, from leveling surveys and InSAR observations, 1923–2008: U.S. Geological Survey Professional Paper 

1788, 68p.

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TITLE PAGE 

Photograph of a U.S. Coast and Geodetic Survey benchmark in Yellowstone National Park.
iii

Contents

Abstract...........................................................................................................................................................1

Introduction.....................................................................................................................................................2

Introduction to Yellowstone Geography and Place Names...........................................................2

Previous Work ................................................................................................................................................4

Nature of This Report....................................................................................................................................9

Historical Context...........................................................................................................................................9

Results of Leveling Surveys and InSAR Observations, 1923–2008 ......................................................10

Initial Leveling Surveys, 1923 and 1934 ...........................................................................................10

Procedures and Accuracy for Leveling Surveys...........................................................................11

1959 Ms

 7.5 Hebgen Lake Earthquake, Coseismic and Postseismic Displacements...............11

1975 M 6.1 Yellowstone Park Earthquake .......................................................................................13

Discovery of Caldera Floor Uplift, 1923 to 1975–1977....................................................................13

Continuing Caldera Floor Uplift, 1976–1984.....................................................................................16

A Pause in Vertical Surface Motion, 1984–1985 ............................................................................16

Caldera-Floor Subsidence, 1986–1995.............................................................................................17

Variable Subsidence and Resumption of Uplift from InSAR Observations, 1992–1997...........19

North Rim Uplift and Caldera-Floor Subsidence, 1997–2002 .......................................................22

Creation of the Yellowstone Volcano Observatory, May 2001.....................................................22

Rapid Uplift of the Sour Creek Dome and North Rim Subsidence, 2004–2008..........................22

Discussion.....................................................................................................................................................28

Observational Constraints on Deformation Models......................................................................29

Constraints from Deformed Shorelines and Microgravity Measurements...............................32

Evidence from Deformed Shorelines......................................................................................32

Evidence from Repeated Microgravity Surveys...................................................................33

Deformation Mechanisms and Models...........................................................................................34

Previous Work ............................................................................................................................34

Our Preferred Model .................................................................................................................36

Current and Future Research Topics........................................................................................................41

Nature of Fluids Involved in Surface Deformation........................................................................41

Structural Control of Deformation Sources....................................................................................43

Temporal Gravity Observations as a Means to Distinguish Between Magma 

and Other Fluids.....................................................................................................................44

Four-Dimensional Numerical Simulations and Modeling—A Way Forward............................46

Conclusions...................................................................................................................................................47

Acknowledgments.......................................................................................................................................48

References....................................................................................................................................................48

Appendix A. Uncertainties in Surface Displacement Measurements from Repeated Leveling 

Surveys and InSAR ........................................................................................................................56

Uncertainty in Leveling Surveys.......................................................................................................56

Uncertainty in InSAR Measurements..............................................................................................56

Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone 

National Park...................................................................................................................................58
iv

Figures

 1. Map of Yellowstone National Park showing the Yellowstone caldera and its two resurgent 

domes, the park’s five entrances and public road system with major junctions, large 

lakes, and other place names used in the text............................................................................4

 2. Panoramic photograph looking south across the Yellowstone caldera, Yellowstone National 

Park, from Mount Washburn, a Tertiary volcanic assemblage that forms part of the 

northeast caldera rim ......................................................................................................................6

 3. Panoramic photograph looking west across the Yellowstone caldera, Yellowstone National 

Park, from Lake Butte, an intrusive mass of Tertiary age that forms part of the east 

caldera rim.........................................................................................................................................6

 4. Oblique map diagram of eastern Snake River Plain and Yellowstone Plateau with outlines 

of three youngest calderas: (I) 2.1 Ma stage 1; (II) 1.3 Ma stage 2; and (III) 0.64 Ma 

Yellowstone caldera (Christiansen, 2001).....................................................................................7

 5. Map of level lines in the Yellowstone-Hebgen Lake region, after Holdahl and Dzurisin (1991). 

See text for dates of leveling surveys along various lines. TJ, Tower-Roosevelt Junction; 

CJ, Canyon Junction; NJ, Norris Junction; MJ, Madison Junction; ML, Mary Lake; OF, 

Old Faithful; IL, Isa Lake; WTJ, West Thumb Junction; FBJ, Fishing Bridge Junction; LB, 

Lake Butte; SL Sylvan Lake; LF, Lewis Falls ...............................................................................12

 6. Map of Yellowstone National Park leveling network and vertical surface displacements 

measured by comparison of leveling surveys in 1923 and 1975–77.......................................14

 7. Diagram showing profiles of topography and vertical displacements during 1976–83 and 

1976–84 along the level line between Lake Butte and Canyon Junction by way of Le 

Hardys Rapids in Yellowstone National Park.............................................................................17

 8. Diagram showing profiles of topography and cumulative vertical displacements from 1984 to 

1995 along the level line between Lake Butte and Canyon Junction by way of Le Hardys 

Rapids in Yellowstone National Park ..........................................................................................18

 9. Diagram showing profiles of topography and vertical displacements along the level line 

between Lewis Falls and Madison Junction by way of West Thumb Junction and the Old 

Faithful area in Yellowstone National Park from 1976 to 1986, 1986 to 1987, 

and 1987 to 1996..............................................................................................................................19

10. Diagram showing profiles of topography and vertical displacements along the level line 

between Lake Butte and Canyon Junction by way of Le Hardys Rapids in Yellowstone 

National Park from 1976 to 1984, from 1984 to 1985, and from 1985 to 1993 .........................20 

11. Four interferograms of the Yellowstone National Park region produced by Wicks and others 

(1998) from ERS-1 and ERS-2 radar images ...............................................................................21

12. Four interferograms superimposed on digital terrain showing range changes (ground 

deformation) at the Yellowstone caldera, Yellowstone National Park, between 1996 and 

2003...................................................................................................................................................23

13. Stack of two deformation interferograms of the Yellowstone caldera, Yellowstone National 

Park, superimposed on digital terrain map, that collectively span the period from 

September 22, 2004, to August 23, 2006......................................................................................24

14. ENVISAT IS2 mode interferogram of the Yellowstone caldera, Yellowstone National Park, 

superimposed on digital terrain map; interferogram was produced from images acquired 

on August 3, 2005, and July 4, 2007..............................................................................................25

15. ENVISAT IS2 mode interferogram of the Yellowstone caldera, Yellowstone National Park, 

superimposed on digital terrain map; interferogram was produced from images acquired 

on September 12, 2007, and August 27, 2008 .............................................................................26
v

16. Diagram showing profiles of topography and vertical displacements along the level line 

between Lake Butte and Canyon Junction by way of Le Hardys Rapids in Yellowstone 

National Park from 1995 to 1998, from 1998 to 2005, and from 2005 to 2007 .........................27

17. Diagrams of estimates of subsurface volume decrease at 6–8 km depth within the 

Yellowstone caldera, Yellowstone National Park, based on a best-fit model of GPS and 

InSAR observations .......................................................................................................................37

18. Map of structural, thermal and volcanic features in and around the Yellowstone caldera, 

Yellowstone National Park............................................................................................................38

19. Diagram showing recent history of vertical surface displacements from leveling and lineof-sight surface displacements from InSAR at Yellowstone’s Sour Creek resurgent dome 

and north caldera rim area, Yellowstone National Park..........................................................39

 B1. Map showing location and status of benchmarks in Yellowstone National 

Park and vicinity .............................................................................................................................59

 B2. Location and status of benchmarks between the West Entrance of Yellowstone National 

Park and Madison Junction .........................................................................................................60

 B3. Location and status of benchmarks between Madison Junction and 

Norris Junction...............................................................................................................................60

 B4. Location and status of benchmarks between Norris Junction and Mammoth Hot Springs 

Junction ...........................................................................................................................................61

 B5. Location and status of benchmarks between the North Entrance of Yellowstone National 

Park and Tower-Roosevelt Junction by way of Mammoth Hot Springs Junction...............62

 B6. Location and status of benchmarks between Tower-Roosevelt Junction and Canyon 

Junction ...........................................................................................................................................63

 B7. Location and status of benchmarks between Canyon Junction and Norris Junction...........64

 B8. Location and status of benchmarks between Canyon Junction and Fishing 

Bridge Junction ..............................................................................................................................64

 B9. Location and status of benchmarks between Fishing Bridge Junction and West Thumb 

Junction ...........................................................................................................................................65

B10. Location and status of benchmarks between West Thumb Junction and Old Faithful 

Junction ...........................................................................................................................................65

B11. Location and status of benchmarks between Old Faithful Junction and 

Madison Junction...........................................................................................................................66

B12. Location and status of benchmarks along the Mary Mountain Trail..........................................66

B13. Location and status of benchmarks between the Northeast Entrance of Yellowstone 

National Park and Tower-Roosevelt Junction...........................................................................67

B14. Location and status of benchmarks between the East Entrance of Yellowstone National 

Park and Fishing Bridge Junction ...............................................................................................67

B15. Location and status of benchmarks between the South Entrance of Yellowstone National 

Park and West Thumb Junction...................................................................................................68

Tables

1. Average vertical displacement rates near Le Hardys Rapids, Yellowstone National Park, from 

leveling surveys..............................................................................................................................15
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History of Surface Displacements at the Yellowstone 

Caldera, Wyoming, from Leveling Surveys and InSAR 

Observations, 1923–2008

By Daniel Dzurisin1, Charles W. Wicks2, and Michael P. Poland3

Abstract

Modern geodetic studies of the Yellowstone caldera, 

Wyoming, and its extraordinary tectonic, magmatic, and 

hydrothermal systems date from an initial leveling survey 

done throughout Yellowstone National Park in 1923 by the 

U.S. Coast and Geodetic Survey. A repeat park-wide survey 

by the U.S. Geological Survey (USGS) and the University 

of Utah during 1975–77 revealed that the central part of the 

caldera floor had risen more than 700 mm since 1923, at an 

average rate of 14±1 mm/yr. From 1983 to 2007, the USGS 

conducted 15 smaller surveys of a single level line that crosses 

the northeast part of the caldera, including the area where 

the greatest uplift had occurred from 1923 to 1975–77. The 

1983 and 1984 surveys showed that uplift had continued at an 

average rate of 22±1 mm/yr since 1975–77, but no additional 

uplift occurred during 1984–85 (-2±5 mm/yr), and during 

1985–95 the area subsided at an average rate of 19±1 mm/yr. 

The change from uplift to subsidence was accompanied by an 

earthquake swarm, the largest ever recorded in the Yellowstone area (as of March 2012), starting in October 1985 and 

located near the northwest rim of the caldera.

Interferometric synthetic aperture radar (InSAR) images 

showed that the area of greatest subsidence migrated from the 

northeast part of the caldera (including the Sour Creek resurgent dome) during 1992–93 to the southwest part (including 

the Mallard Lake resurgent dome) during 1993–95. Thereafter, uplift resumed in the northeast part of the caldera during 

1995–96, while subsidence continued in the southwest part. 

The onset of uplift migrated southwestward, and by mid-1997, 

uplift was occurring throughout the entire caldera (essentially 

1

U.S. Geological Survey, David A. Johnston Cascades Volcano Observatory, 

1300 S.E. Cardinal Court, Building 10, Suite 100, Vancouver, WA 98683-9589.

2

U.S. Geological Survey, 345 Middlefield Road, MS 977, Menlo Park, CA 

94025.

3

U.S. Geological Survey, Hawaiian Volcano Observatory, 1 Crater Rim Drive, 

Hawaii National Park, HI 96718.

rim to rim, including both domes). Consistent with these 

InSAR observations, leveling surveys indicated 24±3 mm of 

uplift in the northeast part of the caldera during 1995–98. The 

beginning of uplift was coincident with or followed shortly 

after an earthquake swarm near the north caldera rim during 

June–July 1995—the strongest swarm since 1985. Rather than 

a single deformation source as inferred from leveling surveys, 

the InSAR images revealed two distinct sources—one beneath 

each resurgent dome on the caldera floor.

Subsequently, repeated GPS surveys (sometimes referred 

to as “campaign” surveys to distinguish them from continuous GPS observations) and InSAR images revealed a third 

deformation source beneath the north caldera rim. The northrim source started to inflate in or about 1995, resulting in as 

much as 80 mm of surface uplift by 2000. Meanwhile, motion 

of the caldera floor changed from uplift to subsidence during 1997–8. The north rim area rose, while the entire caldera 

floor (including both domes) subsided until 2002, when both 

motions paused. Uplift in the northeast part of the caldera 

resumed in mid-2004 at a historically unprecedented rate of 

as much as 70 mm/yr, while the north rim area subsided at a 

lesser rate. Resurveys of the level line across the northeast part 

of the caldera in 2005 and 2007 indicated the greatest average 

uplift rate since the initial survey in 1923—53±3 mm/yr. Data 

from a nearby continuous GPS (CGPS) station showed that 

the uplift rate slowed to 40–50 mm/yr during 2007–8 and to 

near zero by September 2009. Following an intense earthquake swarm during January–February 2010, this one near the 

northwest caldera rim and the strongest since the 1985 swarm 

in the same general area, CGPS stations recorded the onset of 

subsidence throughout the entire caldera.

Any viable model for the cause(s) of ground deformation 

at Yellowstone should account for (1) three distinct deformation sources and their association with both resurgent domes 

and the north caldera rim; (2) interplay among these sources, 

as suggested by the timing of major changes in deformation 

mode; (3) migration of the area of greatest subsidence or uplift 

from the northeast part of the caldera to the southwest part 

during 1992–95 and 1995–97, respectively; (4) repeated cycles 

of uplift and subsidence and sudden changes from uplift to 
2 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

subsidence or vice versa; (5) spatial and temporal relationships 

between changes in deformation mode and strong earthquake 

swarms; and (6) lateral dimensions of all three deforming 

areas that indicate source depths in the range of 5 to 15 km.

We prefer a conceptual model in which surface displacements at Yellowstone are caused primarily by variations in the 

flux of basaltic magma into the crust beneath the caldera. Specifically, we envision a magmatic conduit system beneath the 

northeast part of the caldera that supplies basalt from a mantle 

source to an accumulation zone at 5–10 km depth, perhaps at 

a rheological boundary within a crystallizing rhyolite body 

remnant from past eruptions. Increases in the magma flux 

favor uplift of the caldera and decreases favor subsidence. A 

delicate equilibrium exists among the mass and heat flux from 

basaltic intrusions, heat and volatile loss from the crystallizing 

rhyolite body, and the overlying hydrothermal system. In the 

absence of basalt input, steady subsidence occurs mainly as a 

result of fluid loss from crystallizing rhyolite. At times when 

a self-sealing zone in the deep hydrothermal system prevents 

the escape of magmatic fluid, the resulting pressure increase 

contributes to surface uplift within the caldera; such episodes 

end when the seal ruptures during an earthquake swarm. To 

account for the north rim deformation source, we propose that 

magma or fluid exsolved from magma episodically escapes the 

caldera system at the three-way structural intersection of (1) 

the northern caldera boundary, (2) an active seismic belt to the 

north-northwest that is associated with the Hebgen Lake fault 

zone, and (3) the Norris–Mammoth corridor—a zone of faults, 

volcanic vents, and thermal activity that strikes north from the 

north rim of the caldera near Norris Geyser Basin to Mammoth Hot Springs near the northern boundary of Yellowstone 

National Park. Increased fluid flux out of the caldera by way 

of this intersection favors subsidence of the north rim area, 

and decreased flux favors uplift. This model does not account 

for poroelastic and thermoelastic effects, nonelastic rheology, 

or heat and mass transport in the hot and wet subcaldera crust. 

Such effects almost surely play a role in caldera deformation 

and are an important topic of ongoing research.

Introduction

Even the largest and deadliest volcanic eruption on Earth 

during historical time, the 1815 eruption of Mount Tambora on 

the island of Sumbawa, Indonesia, pales in comparison to the 

greatest eruptions in our planet’s geologic history. The Tambora 

eruption (1) ejected 30–33 km3 (magma volume) of trachyandesite magma (Self and others, 2004); (2) formed a caldera 6 

km in diameter; (3) killed ~10,000 people directly by bomb 

impacts, tephra fall, pyroclastic flows, and tsunamis; and (4) 

produced the “year without a summer” in 1816, which resulted 

in an additional ~100,000 deaths by starvation, disease, and 

hunger (Stothers, 1984; Rampino and others, 1988). For 

comparison, the Yellowstone eruption that produced the Lava 

Creek Tuff at 0.64 Ma (millions of years before the present) 

ejected more than 1,000 km3

 of rhyolitic magma and formed a 

caldera 45 by 85 km across. Yellowstone’s 2.1 Ma Huckleberry 

Ridge Tuff eruption was even larger, with more than 2,450 km3

of magma ejected and collapse of a caldera measuring perhaps 

100 by 60 km across (Christiansen, 2001). There is ample evidence that the Yellowstone magmatic system is still active and 

likely to erupt again, although the timing, nature, and size of 

the next eruption cannot be anticipated. Therefore, understanding the current state of the Yellowstone system—including the 

implications of its dynamic hydrothermal activity, pervasive 

seismicity, rapid ground motion, high heat flow, and copious 

gas emissions—is important for both scientific and hazardassessment reasons.

With reference to ground motion, we have learned in 

recent decades that restless silicic caldera systems, such as 

Yellowstone (Wyoming), Long Valley (California), Rabaul 

(Papua New Guinea), and the Phlegraean Fields (Italy) can 

be remarkably dynamic—with or without an ensuing eruption 

(Newhall and Dzurisin, 1988). At Yellowstone, much of this 

new knowledge has been garnered through repeated leveling 

surveys, GPS observations, and interferometric synthetic aperture radar (InSAR) deformation imagery. The historical record 

at Yellowstone is not nearly as long as at the Phlegraean Fields, 

nor has the amount of surface displacement at Yellowstone 

been as great as at the Phlegraean Fields or at Rabaul. Although 

Yellowstone has not erupted during historical time, as have the 

Phlegraean Fields (1537) and Rabaul (1937, 1994, and other 

smaller events), we are fortunate to have a relatively complete 

record from surficial geology of surface deformation throughout postglacial time, a relatively long historical record from 

leveling, and very detailed information in recent years from 

both InSAR and a large network of continuous GPS (CGPS) 

stations. As a result, we have learned that (1) parts of this largest of all restless calderas can move up or down at rates of tens 

of millimeters per year; (2) the transition from uplift to subsidence or vice versa can occur in a matter of weeks or months; 

(3) such changes in deformation mode generally are associated 

with strong earthquake swarms; and (4) there are three distinct 

deformation sources beneath the caldera that interact in ways 

that are poorly understood. For these reasons and many others, 

Yellowstone is a special place that deserves careful scientific 

study. Here, we compile and discuss results of an ongoing 

study of Yellowstone’s remarkable ground motions, using one 

technique that is more than a century old (leveling) and another 

at the forefront of modern geodesy (InSAR).

Introduction to Yellowstone Geography and 

Place Names

Today the 85-by-45 km Yellowstone caldera is nearly 

invisible to the untrained eye—largely buried beneath post-caldera rhyolite flows with a veneer of glacial deposits that support 

verdant subalpine forest and meadows, punctuated by thermal 
Introduction 3

basins and home to Yellowstone’s renowned wildlife. With its 

winding roads and locally dense forest, Yellowstone National 

Park can be difficult to navigate for the first-time visitor. Likewise, grasping the geography of the park and the relationships 

among place names, structural features, and overlapping time

scales for leveling and InSAR observations could be a daunting 

task for the reader. To make the job easier for those unfamiliar 

with the area, we offer the following primer on relevant Yellowstone geography and place names. Readers who know their 

way around Yellowstone might want to skip to the next section, 

where we describe a large body of previous scientific work as 

context for the leveling and InSAR results.

Established by the U.S. Congress and signed into law by 

President Ulysses S. Grant on March 1, 1872, Yellowstone 

National Park—the world’s first national park—was included 

in the World Network of Biosphere Reserves by the United 

Nations Educational, Scientific, and Cultural Organization 

(UNESCO) in 1976 and was designated a UNESCO World 

Heritage Site in 1978. The park spans an area of 8,980 km2

at the northeast end of the Snake River Plain in the states of 

Wyoming, Montana, and Idaho. It is accessible by road from 

five entrances: (1) the North Entrance near Mammoth, Wyoming; (2) the Northeast Entrance near Silver Gate and Cooke 

City, Montana; (3) the East Entrance west of Cody, Wyoming; 

(4) the South Entrance north of Jackson, Wyoming; and (5) the 

West Entrance near West Yellowstone, Montana (fig. 1). The 

entrance roads connect to the 230-km-long Grand Loop Road, 

a figure-eight pattern of roads that provide access to the park’s 

front-country attractions by way of Mammoth Hot Springs 

Junction, Tower-Roosevelt Junction, Canyon Junction, Fishing 

Bridge Junction, West Thumb Junction, Madison Junction, and 

Norris Junction. The Old Faithful area is located between West 

Thumb Junction and Madison Junction. The area named Lake 

is about 3 km southwest of Fishing Bridge Junction along the 

north shore of Yellowstone Lake; Grant Village is about 5 km 

southeast of West Thumb Junction. Most of the park is designated wilderness and accessible only by trails or, in areas of 

critical wildlife habitat, is off-limits to human visitation.

The Yellowstone caldera occupies approximately onethird of the area of Yellowstone National Park and is accessed 

by the southern half of the Grand Loop Road (fig. 1). Two 

resurgent domes are located within the caldera—the Mallard 

Lake dome in the southwest part and the Sour Creek dome 

in the northeast part. Madison Junction, Norris Junction, and 

Canyon Junction are located near the north topographic rim of 

the caldera. Fishing Bridge Junction and West Thumb Junction 

are located in the central part of the caldera near the north and 

west shore of Yellowstone Lake, respectively. With an area of 

352 km2

 and maximum depth of 118 m, Yellowstone Lake is 

the largest body of water in Yellowstone National Park. Other 

large lakes in the park include Shoshone Lake, Lewis Lake, 

and Heart Lake. Hebgen Lake is located just outside the park’s 

western boundary; the Hebgen Lake fault zone strikes northwest–southeast and extends to the northwest caldera rim near 

Madison Junction. The Yellowstone River at Fishing Bridge 

is the outlet for Yellowstone Lake; an important part of the 

Yellowstone leveling network follows the road that parallels 

the course of the Yellowstone River across the northeast part 

of the caldera between Fishing Bridge Junction and Canyon 

Junction. Another part crosses the southwest part from Lewis 

Falls near the south shore of Lewis Lake to Madison Junction, 

by way of West Thumb Junction and the Old Faithful area.

The Yellowstone leveling network generally follows 

the public road system throughout the park, including the 

entire Grand Loop Road and spurs to each of the park’s five 

entrances (fig. 1). In addition, a backcountry line crosses 

the southern half of the Grand Loop figure-8 along the Mary 

Mountain trail (figs. 5 and 6). The part of the network that has 

been surveyed most often is the cross-caldera line between Lake 

Butte, located on the topographic rim of the caldera along the 

East Entrance road, and Mount Washburn on the opposite rim. 

The line follows the East Entrance Road along the northeast 

shore of Yellowstone Lake from Lake Butte to Fishing Bridge 

Junction. There, it joins the Grand Loop Road and swings northwestward along the Yellowstone River by way of Le Hardys 

Rapids, the Mud Volcano area, Hayden Valley, and Canyon 

Junction. Le Hardys Rapids is an important location because the 

greatest amount of uplift or subsidence along the Lake Butte–

Mount Washburn line typically occurs within a few kilometers 

of the rapids.

To simplify the discussion, we use generalized locations and a small set of specific place names to describe the 

observations and modeling results whenever possible. For 

example, we refer to the southwest, central, and northeast 

parts of the caldera, to the Norris area along the north caldera 

rim, and to the Mallard Lake and Sour Creek resurgent domes. 

Of particular interest are two roughly parallel level lines that 

skirt the resurgent domes—(1) across the northeast part of the 

caldera between Lake Butte and Mount Washburn by way of 

the Sour Creek resurgent dome at Le Hardys Rapids and (2) 

across the southwest part of the caldera from Lewis Falls to 

Madison Junction by way of the Mallard Lake resurgent dome 

at Isa Lake and the Old Faithful area. Roughly perpendicular 

to those two level lines and coincident with the long axis of 

the elliptical caldera is an axis of maximum surface uplift and 

subsidence (fig. 6). Le Hardys Rapids and Isa Lake are located 

near that axis, so we speak of uplift in the northeast part of 

the caldera centered near Le Hardys Rapids or subsidence of 

the southwest part centered near Isa Lake, or (equivalently) 

we refer to uplift of the Sour Creek dome or subsidence of the 

Mallard Lake dome.

An additional level of specificity is required to report 

leveling results at individual benchmarks—for example, the 

amount of uplift or subsidence for a particular time period at 

mark DA3 1934 near Le Hardys Rapids or 51 MDC 1976 near 

Isa Lake (for example, figs. 7–10, 16). Throughout the text, a 

four-digit number at the end of a benchmark name indicates 

the year in which the mark was installed. The year is part 

of the benchmark name, and in a few cases the full name is 

necessary to distinguish between mark names that otherwise 
4 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

would be ambiguous (for example, 11 MDC 1976 and 11 

MDC 1977, both in Yellowstone National Park).

Previous Work

Yellowstone is one of the most intensively studied large 

silicic magma systems on Earth. Its structure and eruptive history are known from detailed geologic mapping and 

petrologic studies (Christiansen, 2001), and the configuration 

of its magmatic system has been inferred from seismic and 

other geophysical investigations (Smith and others, 1982; 

Smith and Siegel, 2000). Conditions within its hydrothermal 

Grand

Loop

Road

Grand Loop

Loop

Road

Road

Grand

20

89

14

89

20

287

287

191

212

Hebgen

Lake

MONTANA

MONTANA

WYOMING WYOMING

IDAHO

IDAHO

Gardiner

Silvergate

Cooke City

LeHardys 

Rapids

Mud Volcano

Hayden

Valley

Lake

Old 

Faithful

Isa

Lake

Lewis

Falls

Hot Springs Basin Hebgen Lake fault zone

Elephant Back flow

and fault zone

Pelican Cone

CALDERA RIM

Yellowstone 

Lake

Lewis 

Lake

Shoshone

Lake

Heart 

Lake

YELLOWSTONE NATIONAL PARK BOUNDARY

To 

Jackson Wyoming

To 

Cody

Wyoming

East

Entrance

North Entrance

West

Entrance

Mammoth Hot

Springs Junction Tower-Roosevelt

Junction

Madison

Junction

Fishing Bridge 

Junction Steamboat 

Point

Lake Butte

Flat 

Mountain

Mount 

Sheridan

Canyon

Junction

Mount Washburn

West 

Thumb

Junction

Norris

Junction

West 

Yellowstone

Mallard Lake

Dome

Sour Creek

Dome

0 5 10 20 KILOMETERS

0 5 10 MILES

Figure 1. Map of Yellowstone National Park showing the Yellowstone caldera and its two resurgent 

domes, the park’s five entrances and public road system with major junctions, large lakes, and other 

place names used in the text. See figure 4 for the relation between the Yellowstone caldera shown here 

and two older calderas on the Yellowstone Plateau. Figures 5 and 6 show the network of level lines 

in and around Yellowstone National Park, which mostly follows the public road system shown here. 

The Yellowstone River flows northwestward from the outlet of Yellowstone Lake near Fishing Bridge 

Junction by way of Le Hardys Rapids to near Canyon Junction and beyond. The Grand Loop Road 

parallels the river in this area, and one of the main level lines discussed in the text follows the road.

system have been deduced from observational, experimental, 

and theoretical studies (Fournier, 1989, 1999, 2007; Fournier 

and Thompson, 1993) and by scientific drilling (Allen and 

Day, 1935; Fenner, 1936; White and others, 1971, 1975). This 

section includes brief summaries and key references for each 

of these major lines of investigation to provide necessary 

context for the main subject of this report—the remarkable 

record of surface displacements at Yellowstone. The report 

deals mainly with leveling and InSAR results for 1983–2008, 

because the authors have first-hand experience with those 

datasets. For completeness, we also discuss (1) results from 

earlier leveling surveys, including those pertinent to coseismic 

and postseismic motions associated with two large historical 
Previous Work 5

earthquakes in the Yellowstone region; (2) surficial-geology 

studies that extend the deformation record back in time for 

several thousand years; and (3) CGPS observations that bring 

the record up to date since the most recent leveling and InSAR 

observations in 2007–8. Readers who are familiar with the 

scientific literature dealing with Yellowstone’s eruptive history, magmatic system, and hydrothermal system might prefer 

to skip this section. Others might find it helpful as context for 

what follows. Additional information is available in numerous 

publications cited in the text.

Christiansen (2001) provides an excellent summary of 

the development of scientific knowledge about Yellowstone, 

beginning with F.V. Hayden’s Geological and Geographical Survey of the Territories (Hayden, 1872; 1873; Bradley, 

1873; Peale, 1873). According to Christiansen (2001, p. G3), 

“Hayden’s perceptive first description of the Yellowstone 

Plateau, viewed from the summit of Mount Washburn, even 

anticipates the concept of the Yellowstone caldera.” Christiansen quotes Hayden (1872, p. 81):

From the summit of Mount Washburn, a bird’s-eye 

view of the entire basin may be obtained, with the 

mountains surrounding it on every side without any 

apparent break in the rim * * *. It is probable that 

during the Pliocene period the entire country drained 

by the sources of the Yellowstone and the Columbia 

was the scene of as great volcanic activity as that of 

any portion of the globe. It might be called one vast 

crater, made up of thousands of smaller volcanic 

vents and fissures out of which the fluid interior of 

the earth, fragments of rock, and volcanic dust were 

poured in unlimited quantities * * *.

The view across the Yellowstone caldera from the summit 

of Mount Washburn, and the opposing view from the summit 

of Lake Butte toward Mount Washburn, are mostly unchanged 

since Hayden’s time (figs. 2 and 3).4

 What has changed dramatically is our understanding of Yellowstone’s tumultuous 

volcanic history, due in large part to four decades of meticulous study by Christiansen and others. This monumental body 

of work, documented in Christiansen (2001) and several other 

seminal papers on the subject (for example, Christiansen, 

1984; Christiansen and others, 2002, and references therein; 

Morgan and others, 2009), defines the current state of knowledge concerning the Yellowstone caldera, the eruptive history 

of the Yellowstone Plateau volcanic field, and the evolution of 

the Yellowstone magmatic system.

We know from Christiansen’s (2001) work that the Yellowstone caldera is the youngest of three collapse calderas in 

the Yellowstone Plateau volcanic field of Wyoming, Idaho, and 

Montana (fig. 4). The first caldera formed 2.1 Ma during eruption of the Huckleberry Ridge Tuff, a rhyolitic ash-flow sheet of 

more than 2,450 km3

 that was emplaced over an area of 15,500 

km2 (Christiansen, 2001, p. G53–G63). The more than 280-km3

Mesa Falls Tuff was emplaced 1.3 Ma over nearly 2,700 km2, 

resulting in formation of the approximately 16-km-diameter 

second-cycle caldera (Christiansen, 2001, p. G64–G66). The 

third Yellowstone Plateau volcanic cycle, which began about 

1.2 Ma, culminated at 0.64 Ma with eruption of the >1,000-km3

Lava Creek Tuff over more than 7,500 km2 and collapse of the 

85-by-45 km Yellowstone caldera.5 The Sour Creek structural 

dome in the northeast part of the caldera became resurgent soon 

after caldera collapse, and the present Mallard Lake dome in 

the southwestern part formed about 160 ka (thousands of years 

before the present) in response to renewed influx of magma 

beneath the caldera.6 In a series of eruptive episodes at 164 ka, 

152 ka, 114 ka, 102 ka, and 72 ka, intracaldera rhyolitic lava 

flows with an aggregate volume of more than 500 km3 partly 

filled the caldera and in places overtopped its rim (Christiansen, 

2001. p. G16–G48). As recently as 59.15 ± 1.97 ka (40Ar/39Ar 

age), the Obsidian Creek rhyolite flows erupted just north of 

the Yellowstone caldera near the south end of the Norris–Mammoth corridor (Wooton and Spell, 2007).7 Since Yellowstone’s 

most recent magmatic eruption, an extensive hydrothermal 

system has developed within the caldera (Fournier, 1989). At 

least 18 large hydrothermal explosions have formed craters 

more than 100 m wide since the Yellowstone Plateau was 

last glaciated about 16 ka, and at least 26 smaller hydrothermal explosions have been documented since the park was 

established in 1872; others undoubtedly escaped observation 

(Muffler and others, 1971; Christiansen and others, 2007). 

Considering the prospects for future volcanism at Yellowstone, Christiansen (2001, p. G131) concluded:

4

On March 1, 1872, President Ulysses S. Grant signed a declaration by 

the United States Congress that established Yellowstone National Park “* 

* * as a public park or pleasuring ground for the benefit and enjoyment of 

the people.” The newly formed National Park Service was given jurisdiction over the country’s first national park and charged with preserving its 

natural wonders for future generations. National Park Service stewardship 

of Yellowstone has endured numerous challenges over the years, none more 

daunting than those faced by park managers today. For a fascinating personal 

account of the pressures involved in preserving the United States’ premier 

national park from the perspective of a former Chief Ranger at Yellowstone, 

see Sholly and Newman (1991). 

5

Christiansen (2001) refers to the three calderas of the Yellowstone Plateau 

volcanic field as the first-cycle caldera (2.1 Ma), the second-cycle caldera (1.3 

Ma), and the Yellowstone caldera (0.64 Ma). The associated ash-flow sheets 

he designates are the Huckleberry Ridge Tuff, Mesa Falls Tuff, and Lava 

Creek Tuff, respectively. Other authors have used the names Huckleberry 

Ridge caldera, Island Park caldera, and Lava Creek caldera.

6

Christiansen (2001, p. G39) states: “The present [Mallard Lake] dome, on 

which only the Mallard Lake flow is exposed, must be about 160,000 years 

old * * *. Nevertheless, some evidence suggests that there may also have been 

early resurgent doming of the Mallard Lake block * * *.”

7

 The Norris–Mammoth corridor is a zone of faults, volcanic vents, and 

thermal activity that strikes north from the north rim of the Yellowstone caldera 

near Norris Geyser Basin to Mammoth Hot Springs near the northern boundary 

of Yellowstone National Park. It contains most of the young rhyolitic domes 

and basaltic lava vents of the past few hundred thousand years that are located 

outside the caldera margin (White and others, 1988, p. 3–4, 73–77).
6 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

A p p r o x i m a t e e x t e n t o f t h e Y e l l o w s t o n e C a l d e r a

East Elephant Back West

Yellowstone Lake

Lake Butte

Hayden Valley

Grand Canyon, Yellowstone River

Sour Creek Resurgent Dome

Steamboat Point Teton Range

Yellowstone River

Teton

Range Mount Sheridan

Frank Island

West

Thumb

Sour Creek

Resurgent Dome Yellowstone River

Steamboat Point

Gallatin Range

Flat Mountain

Elephant Back Mount

Washburn

A p p r o x i m a t e e x t e n t o f t h e Y e l l o w s t o n e C a l d e r a

Y e l l o w s t o n e L a k e

South North

Figure 3.

Figure 2. Panoramic photograph looking south across the Yellowstone caldera, Yellowstone National Park, from Mount Washburn, a 

Tertiary volcanic assemblage that forms part of the northeast caldera rim. The Grand Canyon of the Yellowstone River is visible in the 

middle ground in the left part of the image. Lake Butte, the vantage point for figure 3, is visible in the distance along the east (left) shore 

of Yellowstone Lake. Features visible in both figure 2 and figure 3 include the Sour Creek resurgent dome, Steamboat Point, Elephant 

Back lava flow, and Teton Range. Visible in the left center of this image is the route taken by a level line that was measured in 1923 and 

repeatedly measured from 1976 to 2007. The line extends from the south flank of Mount Washburn to Canyon Village Junction, through 

Hayden Valley, past Le Hardys Rapids and the nose of the Elephant Back lava flow and through Fishing Bridge Junction southeastward, 

past Steamboat Point, to the vicinity of Lake Butte—a road distance of 50 km. The route was chosen because it included Le Hardys 

Rapids, the site of maximum uplift measured by leveling surveys in 1923 and 1975–77, and because it is the shortest road route across 

the 85-by-45-km caldera. (USGS composite photograph by D. Dzurisin.)

Figure 3. Panoramic photograph looking west across the Yellowstone caldera, Yellowstone National Park, from Lake Butte, an intrusive 

mass of Tertiary age that forms part of the east caldera rim. The topographic basin occupied by the caldera is approximately 85 km 

southwest-northeast by 45 km northwest-southeast. The northern part of Yellowstone Lake, in the foreground, occupies about 10 

percent of the basin, which is largely buried by rhyolite lava flows that erupted from vents within the caldera between 151.4 ka and 

71.2 ka (Christiansen, 2001). The Elephant Back lava flow, visible on the skyline in the right third of the image, is among the oldest 

of the intracaldera flows. Flat Mountain, 20 km distant and in the left third of the image, forms part of the south caldera rim. Mount 

Washburn, the vantage point for figure 2, is 34 km distant and near the right edge; it forms part of the northeast rim. The Teton Range, 

100 km distant and near the left edge, is one of several ranges trending north-south to northwest-southeast that predate the caldera 

and remain seismically active. Other examples are the Madison Range and Gallatin Range to the north; the latter is visible in the right 

third of the image. A 50-km-long level line across the caldera floor from Lake Butte to Mount Washburn by way of Steamboat Point, Le 

Hardys Rapids (Yellowstone River near nose of Elephant Back flow), and Hayden Valley was measured in 1923 and repeatedly measured 

from 1976 to 2007 to study uplift and subsidence of the caldera floor. (USGS composite photograph by D. Dzurisin.)
Previous Work 7

1.5 2 2.5 3 3.5 4

Elevation, in kilometers

111.5o 

W

44.0o 

N

44.5o 

N

45.0o 

N

111.0o 

W 110.5o 

W 110.0o 

W

Mallard Lake Dome

1959

M 7.5

Sour Creek

 Dome

1975

M 6.1

Hebgen Lake

Figure 4. Oblique map diagram of eastern Snake River Plain and Yellowstone Plateau with outlines of three youngest 

calderas: (I) 2.1 Ma stage 1; (II) 1.3 Ma stage 2; and (III) 0.64 Ma Yellowstone caldera (Christiansen, 2001). The stage-1 

caldera is largest; the stage-2 caldera and the Yellowstone caldera lie entirely within the stage-1 caldera. Also shown, 

outlines of Mallard Lake and Sour Creek resurgent domes. White stars indicate epicenters of 1959 Ms

 (surface-wave 

magnitude) 7.5 Hebgen Lake earthquake and 1975 M 6.1 Yellowstone Park earthquake.

No rhyolitic eruptions have occurred for about 

70,000 years within the third-cycle source area of 

the volcanic field, and rhyolitic volcanism related 

to this cycle may have waned. The very high heat 

flow represented by the Yellowstone hydrothermal 

system, however, as well as individually inconclusive geologic, gravity, aeromagnetic, geodetic, and 

seismic evidence, together suggest that rhyolitic 

magma is still crystallizing and perhaps still rising 

below the plateau region.

The seismic and other geophysical evidence for magma 

beneath Yellowstone to which Christiansen (2001) refers was 

summarized by Eaton and others (1975), who wrote 

(p. 795–6):

Doming 150,000 years ago, followed by voluminous 

rhyolitic extrusions as recently as 70,000 years ago, 

and high convective heat flow at present indicate 

that the latest phase of volcanism may represent 

a new magmatic insurgence. These observations, 

coupled with (i) localized postglacial arcuate faulting beyond the northeast margin of the Yellowstone 

caldera, (ii) a major gravity low with steep bounding 

gradients and an amplitude regionally atypical for 

the elevation of the plateau, (iii) an aeromagnetic 

low reflecting extensive hydrothermal alteration and 

possibly indicating the presence of shallow material 

above its Curie temperature, (iv) only minor shallow 

seismicity within the caldera (in contrast to a high 

level of activity in some areas immediately outside), 

(v) attenuation and change of character of seismic 

waves crossing the caldera area, and (vi) a strong 

azimuthal pattern of teleseismic P-wave delays, 

strongly suggest that a body of magma underlies the 

region of the rhyolite plateau, including the Tertiary 

volcanics immediately to its northeast.
8 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Eaton and others’ (1975) conclusion that magma exists 

today in the crust beneath Yellowstone has been corroborated by more than three decades of subsequent geophysical investigations, which are documented or summarized in 

numerous publications (for example, Smith and Braile, 1984, 

1994; Smith and Siegel, 2000; Waite and Smith, 2002; Husen 

and others 2004; Waite and others, 2005, 2006; Lowenstern 

and others, 2006; Lowenstern and Hurwitz, 2008). The fact 

that post-caldera basaltic volcanism has occurred around the 

periphery of, but not inside, the Yellowstone caldera leads to 

the same conclusion—the presence of a body of partly molten 

silicic magma in the crust beneath the caldera that prevents 

denser, basaltic magmas from rising to the surface (Hildreth, 

1981; Hildreth and others, 1991; Christiansen, 2001).

There is general agreement that Yellowstone is underlain by 

a persistent source of mantle-derived, basaltic magma—the Yellowstone hotspot—that produced a linear track of silicic, calderaforming volcanism stretching from the vicinity of the OregonNevada-Idaho border at 16 Ma to the Yellowstone Plateau today.8

Whether the Yellowstone hotspot (1) fits the deep-mantle plume 

model with convection originating at the core-mantle boundary 

(for example, Morgan, 1971, 1972a,b; Pierce and Morgan, 1992; 

Smith and Braile, 1994), (2) fits an upper-mantle model with convective plumes rising from a phase-change discontinuity at 410 or 

660 km depth (for example, Waite and others, 2006), or (3) instead 

reflects feedback between upper-mantle convection and regional 

lithospheric tectonics (Christiansen and others, 2002), remains 

unresolved. However, we suspect that Waite and others (2006, p. 

B04303) are correct in their assessment that all three models are 

relevant at Yellowstone:

Upper mantle convection models are not contradicted by a [deep] plume model. Rather, convection, 

lithosphere convection, and upwelling from below 

likely work together at Yellowstone. Small-scale 

convection helps explain the strong low-velocity 

anomaly beneath Yellowstone and the Snake River 

Plain to ~200 km depth. The high topography 

on both sides of the ESRP [Eastern Snake River 

Plain] may be supported by melt residuum that 

has been pushed away from the upwelling zone 

under the ESRP. The possible eastward migration 

of the basin-range extensional regime is partly a 

consequence of the active system moving in the 

direction opposite plate motion. Without all three 

mechanisms, Yellowstone volcanism may not have 

persisted for ~16 million years.

Germane to our discussion of ground deformation at 

Yellowstone, all three models for the origin of the Yellowstone hotspot include high intrusion rates of basaltic magma, 

comparable to the Hawaiian hotspot and as much as 100 times 

greater than the long-term eruption rate of rhyolite in Yellowstone (0.003 km3

/yr; Christiansen, 2001).

Lowenstern and Hurwitz (2008) make a compelling case 

for a high intrusion rate of basalt into the lower to mid-crust 

beneath Yellowstone, although they do not exclude an alternative scenario involving convection of a large volume of low 

viscosity basalt in the lower crust with episodic release of gases. 

The Yellowstone system emits at least 45,000±16,000 metric 

tons per day of CO2, which is about 5 percent of the global 

volcanogenic CO2 flux (Werner and Brantley, 2003). Werner 

and Brantley (2003) estimate that only 30–50 percent of that 

total is derived from pre-Tertiary basement rocks; the remainder 

must have exsolved from magma. Lowenstern and Hurwitz 

(2008) argue that the only plausible source for that much CO2

is approximately 0.3 km3/yr of basalt degassing beneath the 

caldera. The volume of silicic magma that is estimated to exist 

beneath the caldera (~1.5×104 km3; Lowenstern and others, 

2006) is an insufficient source, because Yellowstone silicic 

magmas contain less than 500 parts per million dissolved CO2. 

At the observed CO2 degassing rate, the silicic reservoir would 

be entirely purged of dissolved CO2 in about 1,000 years. 

Lowenstern and Hurwitz (2008, p. 37) point out: “Assigning the 

annual CO2 flux to a plausible amount of annually degassed and 

crystallized rhyolitic magma (0.1 km3; Fournier 1989) requires 

5.5 wt percent CO2, about 20 times that which can be dissolved 

at 400 MPa (~16 km depth) in rhyolitic liquid (Lowenstern, 

2001).” Simply stated, the remarkably high CO2 flux at Yellowstone can be explained only if the caldera system is underlain by 

a large reservoir of basaltic magma that is persistently replenished—by intrusion, convection, or both—at a rate comparable 

to that attributed to the Hawaiian hotspot.

The same conclusion is supported by the relative proportions of magmatic gases emitted at Yellowstone. Lowenstern 

and Hurwitz (2008, p. 37) note that “* * * orders of magnitude 

more CO2

 is emitted through the Yellowstone hydrothermal 

system than Cl, F, or S * * *.” Silicate melt inclusions in Yellowstone rhyolitic lavas contain abundant dissolved Cl and F, 

but scarce CO2 and S. Basaltic magmas, in contrast, generally 

contain lower concentrations of dissolved Cl and F, and 10–30 

times more dissolved CO2 than mid-crustal rhyolites. This is 

due in part to the fact that basalt becomes supersaturated with 

CO2 at lower crustal depths. Lowenstern and Hurwitz (2008, p. 

37) conclude: “* * * basaltic intrusions could contain 50 to 100 

times more available CO2 than the silicic magmas themselves. 

The basaltic input means that magma continues to accumulate 

and sustains the overlying silicic magma reservoir * * *.”

Heat from magmatic intrusions in the crust, both fresh 

influxes of basalt and older, cooling bodies of rhyolite, powers 

Yellowstone’s extensive hydrothermal system. Thermal output 

8

Several authors, including Pierce and Morgan (1992) and Smith and Braile 

(1994), have noted the following age progression in silicic volcanic fields and 

associated calderas: McDermitt (Nevada, 16.1 Ma), Owyhee (Idaho, 13.8 Ma), 

Bruneau–Jarbridge, Twin Falls, and Picabo (Idaho, 11.3–10.3 Ma, respectively), Heise (Idaho, 6.5–4.3 Ma), and Yellowstone (Wyoming, 2.1–0.6 Ma). 

The pattern is attributed to southwest motion of the crust relative to a mantle 

hotspot at an average rate of 4.5 cm/yr—a combination of North American 

Plate motion (~2.5 cm/yr) and crustal extension (1–2 cm/yr). At the present 

time, the Yellowstone hotspot is beneath the northeast part of the Yellowstone 

Plateau, about 700 km northeast of the McDermitt volcanic field.
Historical Context 9

of the Yellowstone system is calculated to be in the range 4.5–

6.0 gigawatts (Fournier 1989; Friedman and Norton, 2007), 

which corresponds to 1,550–2,100 milliwatts per square meter 

over the 2,900-km2 Yellowstone caldera. This heat flux density 

is 30–40 times that of the neighboring Rocky Mountains. 

Lowenstern and Hurwitz (2008, p. 36) point out:

If this high surface heat flux were transferred by 

conduction alone, it would imply a temperature gradient of 700 to 1000°C per kilometer (Fournier 1989), 

suggesting the presence of very shallow magma; 

however, research drilling in the 1960s (White et al. 

1971) revealed temperature–depth relations inconsistent with temperatures exceeding ~310°C at 1000 m. 

Clearly, the enormous hydrothermal system is able 

to act as a wick, transferring heat advectively from a 

deep but very large magma source.

Various interactions among basaltic intrusions, a large 

rhyolitic body containing an unknown proportion of partial 

melt9, and the overlying hydrothermal system are key to 

understanding the possible causes of ground deformation at 

Yellowstone (see, for example, Dzurisin and others, 1990). 

This subject is explored below in the Discussion section, 

following a description of leveling, InSAR, and other ground 

deformation observations.

Nature of This Report

This report (1) summarizes knowledge about Yellowstone’s eruptive history, magmatic system, and hydrothermal 

system, particularly as those topics pertain to the interpretation 

of ground deformation measurements; (2) describes the history 

of leveling and InSAR observations at Yellowstone, as well as 

other evidence for paleo and contemporary ground deformation; 

(3) discusses the evolution of thought concerning the mechanisms of ground deformation at Yellowstone; and (4) provides 

an archive of information concerning more than 500 geodetic 

benchmarks in Yellowstone National Park (see data files that 

accompany this report; in the printed version, these files are 

provided on a CD–ROM). We begin with an overview of leveling and InSAR results for 1983–2008 as an introduction to the 

dynamic nature of the Yellowstone system.

Historical Context

In 1980, the attention of many volcanologists was 

focused on Mount St. Helens in the aftermath of its historic 

landslide and eruption on May 18 (Lipman and Mullineaux, 

1981). Starting one week after that monumental event, the tiny 

town of Mammoth Lakes, California, was rattled by an intense 

swarm of earthquakes, including four magnitude (M) 6 shocks 

on May 25–27, 1980. Mammoth Lakes is nestled in the south 

moat of the Long Valley caldera, between Mammoth Mountain 

and the caldera’s resurgent dome, along the eastern front of the 

tectonically active Sierra Nevada range. In September 1980, 

first-order releveling of a line along U.S. Highway 395 across 

the western parts of the caldera floor and resurgent dome 

revealed about 250 mm of surface uplift centered on the dome 

since the previous survey in 1975.10 Resurveys showed that the 

dome rose an additional 100 mm during 1980–82 and 70 mm 

during 1982–83, accompanied by more earthquake swarms. 

The latter period encompassed a particularly strong swarm in 

January 1983, this one with characteristics indicative of dike 

intrusion to within a few kilometers of the surface (Savage 

and Cockerham, 1984; Hill, 1984; Hill and others, 1985a,b). 

The unusual seismicity and ground deformation underscored 

the possibility of renewed volcanic activity in the area, and in 

May 1982 the United States Geological Survey (USGS) issued 

a Notice of Potential Volcanic Hazard (Miller and others, 

1982; Miller, 1985; 1989).

In autumn 1983, concern over the developing situation at 

Long Valley by managers of the USGS Volcano Hazards Program led them to initiate a project, headquartered at the USGS 

Cascades Volcano Observatory (CVO) in Vancouver, Washington, to study processes of unrest at large silicic calderas. A 

primary objective of the Physical Processes in Large Silicic 

Magma Systems project was to establish a context for interpretation of the activity at Long Valley through comparisons 

with other restless calderas around the world (Newhall and 

Dzurisin, 1988). Two notable examples were the Phlegraean 

Fields caldera near Naples, Italy, and the Rabaul caldera in 

Papua New Guinea, where sporadic earthquake swarms and 

rapid ground uplift were generating concerns similar to those 

expressed for Long Valley. It was known from seismic monitoring, repeated leveling surveys, and water-level records from 

Yellowstone Lake that the Yellowstone caldera also had been 

restless, having risen at an average rate of 14 mm/yr between 

leveling surveys in 1923 and 1975–77 (Pelton and Smith, 

1979; 1982; Hamilton, 1987). To determine whether the uplift 

at Yellowstone was continuing and, if so, to study it together 

with the uplift at Long Valley, CVO staff assigned to the 

Physical Processes in Large Silicic Magma Systems project 

decided to relevel a line across the Yellowstone caldera where 

the greatest uplift had been measured from 1923 to 1975–77.

9

A recent modeling study by Chu and others (2010) of first teleseismic 

P-wave arrivals recorded by broadband seismic stations in Yellowstone 

National Park led those authors to identify a magma body at 5–15 km depth 

beneath the Yellowstone caldera with a volume greater than 4,300 km3 and a 

porosity of about 32 percent—the latter filled with at least 90 percent rhyolite 

melt and 8 percent water-CO2 by volume. That proportion of partial melt is 

considerably higher than indicated by earlier seismic tomography studies (for 

example, Miller and Smith, 1999; Husen and others, 2004), which place the 

number at a few to perhaps 10 percent. The discrepancy is sure to receive 

considerable attention from volcano seismologists following an upgrade of the 

Yellowstone broadband network in 2010–11.

 10See section Results of Leveling Surveys and InSAR Observations, 

1923–2008, for description of first-order and second-order leveling surveys.
10 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Thus it was that Ken Yamashita, a recent transplant to 

CVO from the USGS Hawaiian Volcano Observatory (HVO), 

found himself in Yellowstone leveling through wind-driven 

snow in October 1983. He managed to complete 17 km of 

double-run, first-order leveling in three weeks before winter 

storms forced an end to the survey (see next section for discussion of differences between first-order and second-order leveling surveys). The plan had been to survey a 44-km traverse 

across the caldera floor where the greatest uplift had been 

measured by surveys in 1923 and 1975–77. Failure to finish 

the survey as planned caused some initial disappointment, 

but the reason soon became clear. Back at CVO, Yamashita 

discovered that the allowable discrepancy between two height 

differences measured at each setup of the leveling instrument 

had been incorrectly set at 0.03 mm, instead of the 0.30 mm 

value specified by the Federal Geodetic Control Committee 

(1984) for first-order leveling surveys. He had double-run 17 

km of level line with a tolerance comparable to the diameter 

of a human hair and an order of magnitude more stringent than 

that specified for the most precise surveys conducted by the 

National Geodetic Survey (NGS)!

The mistake was corrected before the next Yellowstone leveling survey in September 1984, when the line was 

extended from Lake Butte on the southeast rim of the caldera 

by way of Fishing Bridge, Le Hardys Rapids, and Hayden 

Valley to Canyon Junction near the northeast rim—a distance 

along the line (stadia distance) of 44 km (figs. 1 and 6). The 

line was extended 5 km farther north to Mount Washburn 

in 1987 and, from 1983 to 2007, CVO staff members and 

dozens of volunteers measured part or all of the line 15 times 

(figs. 1, 6, 8, and 16). Results of the 1983 and 1984 surveys 

indicated that uplift of Le Hardys Rapids had continued since 

the 1975–77 survey, but thereafter the pattern changed. No 

additional uplift was detected by the 1985 survey and, starting in 1986 and continuing through 1995, Le Hardys Rapids 

and the adjacent Sour Creek resurgent dome subsided at rates 

comparable to or greater than the uplift rates measured previously. The discovery triggered renewed interest in the cause(s) 

of surface deformation at Yellowstone and other large silicic 

calderas, including Long Valley (Dzurisin and Yamashita, 

1987; Dzurisin and others, 1990; 1994; 1999; Hill and others, 

1985a,b). Soon, with the introduction of InSAR (then a new 

remote-sensing technique for measuring deformation), the 

puzzle became even more complicated.

The first useful InSAR image of the Yellowstone region 

showed subsidence in the vicinity of the Sour Creek dome 

during August 1992–June 1993, which was consistent with 

leveling results for the same period (Wicks and others, 

1998).11 No leveling was done along the Lake Butte–Mount 

Washburn line in 1994, but surveys in 1993 and 1995 indicated continued subsidence centered near Le Hardys Rapids. 

An interferogram for June 1993–August 1995 also showed 

subsidence of the Sour Creek dome, but even more subsidence of the Mallard Lake dome in the southwest part of 

the caldera. Uplift of the Sour Creek dome resumed during 

August 1995–September 1996, while subsidence of the Mallard Lake dome continued. By July 1995–June 1997, InSAR 

observations indicated that both domes were rising—a result 

that was corroborated by leveling along the Lake Butte–

Mount Washburn line in September 1998 (Wicks and others, 

1998; Dzurisin and others, 1999).

A third deformation source, this one beneath the north 

caldera rim near Norris Geyser Basin, was identified from 

InSAR images that collectively span 1996–2002 (Wicks and 

others, 2006). Localized uplift in that area was first detected 

by GPS surveys made in 1995 and 2000 (Puskas and others, 

2007). As the north rim source inflated from 1995 to 1997, 

motion of the central part of the caldera, including both 

resurgent domes, changed from uplift to subsidence between 

late 1997 and early 1998. The north rim source continued 

to inflate, while the caldera floor subsided until 2002, when 

both motions paused (Wicks and others, 2006). Then, starting 

in mid-2004, CGPS observations and interferograms showed 

strong uplift of the entire caldera floor, at rates as high as 70 

mm/yr, and lesser subsidence along the north caldera rim. 

Leveling surveys along the Lake Butte–Mount Washburn 

line in 2005 and 2007 likewise indicated historically high 

rates of uplift near Le Hardys Rapids. Uplift of the caldera 

floor slowed during 2007–8, and by September 2009, it had 

slowed to the point that it may have stopped (Yellowstone 

Volcano Observatory (YVO) Monthly Update issued January 

4, 2010; http://volcanoes.usgs.gov/yvo/activity/archive/2009.

php). An undertaking that started 25 years earlier, using 

a century-old surveying technique to monitor a pattern of 

ground motion that was thought to be simple in space and 

time, had evolved into a multifaceted investigation of a complex and unexpectedly dynamic system.

Results of Leveling Surveys and InSAR 

Observations, 1923–2008

Initial Leveling Surveys, 1923 and 1934

The history of geodetic measurements in the Yellowstone 

region begins in 1923, when the U.S. Coast and Geodetic 

Survey (USC&GS) conducted initial leveling surveys along 

first-order lines from Livingston, Montana, to Mammoth, 

Wyoming; from West Yellowstone, Montana, to Marysville, 

Idaho; and along second-order lines in Yellowstone National 

Park. Within the park the leveling network follows major roads, 

including the figure-eight-shaped Grand Loop Road and spurs 

to each of five park entrances (figs. 1, 5, and 6).12 In 1934, the 

USC&GS did additional second-order surveys on lines from 

 11From 1984 to 2007, all of the CVO leveling surveys along the line from 

Lake Butte to Canyon Junction by way of Le Hardys Rapids were conducted 

in the month of September. The 1983 survey was conducted in October.
Results of Leveling Surveys and InSAR Observations, 1923–2008 11

Bozeman, Sappington, and Monida, Montana, each of which 

terminate at the town of West Yellowstone, Montana (Holdahl 

and Dzurisin, 1991). West Yellowstone is along the western 

boundary of Yellowstone National Park, adjacent to the park’s 

West Entrance (fig. 1). All of these lines are part of a nationwide 

vertical control network with a main purpose of establishing 

benchmark elevations, but they also provide a means to measure 

vertical surface displacements by repeated surveys. The primary 

purpose of the early USC&GS surveys was to upgrade the 

vertical control network in the Yellowstone region, not to serve 

as a baseline for future caldera deformation studies. In 1923, 

Yellowstone National Park had been in existence for more than 

50 years, and it was time to tie its road system into the nation’s 

vertical control network. At the time, no one could have imagined that some of the benchmark elevations established by the 

1923 survey would change by several decimeters by the time 

they were resurveyed during 1975–77.

Procedures and Accuracy for Leveling Surveys

A report by the Federal Geodetic Control Committee 

(1984) establishes standards and specifications for geodetic 

control networks in the United States, including leveling networks and surveys. The FGCC report divides leveling surveys 

into five types based on the procedures and equipment used 

and on the expected accuracy. The most stringent specifications apply to first-order, class I networks and surveys, which 

provide the greatest precision. Next in order of decreasing 

precision are first-order, class II; second-order, class I and 

II; and third-order leveling surveys. The report specifies, for 

each order and class of leveling, the requirements for network 

geometry, instrumentation, calibration procedures, field procedures, and office procedures. All of the Yellowstone leveling 

surveys conducted by CVO between 1983 and 2007 conformed to first-order, class II standards or better. First-order, 

class I surveys are double-run; first-order, class II surveys can 

be single run (Dzurisin, 2007, p. 53–61).

The equipment and procedures used for leveling surveys at 

Yellowstone were described by Dzurisin and Yamashita (1987) 

and by Dzurisin and others (1990, 1994, and 1999). Uncertainties associated with leveling and InSAR observations are 

addressed in appendix A. Throughout this report, uncertainties 

in vertical displacement measurements from leveling surveys 

are expressed as one standard deviation from the combination of 

random leveling error and long-term benchmark instability. The 

one-sigma uncertainty in range changes (that is, line-of-sight 

surface displacements) from InSAR observations is difficult 

to specify, so we adopt what we believe to be a conservative 

estimate of ±1 cm (see appendix A).

1959 Ms 7.5 Hebgen Lake Earthquake, Coseismic 

and Postseismic Displacements

On August 18, 1959, the largest earthquake ever recorded 

in the Yellowstone region, the Ms (surface-wave magnitude) 

7.5 Hebgen Lake earthquake, occurred about 20 km northwest 

of the northwest rim of the Yellowstone caldera and about 15 

km north of West Yellowstone.13 Within 24 hours, six aftershocks of M 5.5–6.3 had occurred in an east-west-trending 

zone nearly 100 km long, centered on the main shock. The 

Hebgen Lake earthquake caused 28 fatalities and extensive 

damage to highways and timber, produced a 26-km-long zone 

of surface faulting with scarps as high as 6 m, and triggered 

a catastrophic rockslide that dammed the Madison River and 

created Quake Lake (Smith and Arabasz, 1991). To measure 

coseismic and postseismic vertical surface displacements associated with the earthquake and its aftershocks, the USC&GS 

conducted first-order relevelings in 1959, 1960, 1964, and 

1967 along lines in Idaho and Montana that converge at West 

Yellowstone (fig. 5). All lines in Yellowstone National Park, 

plus the line from Sappington, Montana, to West Yellowstone, 

Montana, were releveled to first-order standards in 1987. Holdahl and Dzurisin (1991) summarized results of those surveys, 

including 6 m of coseismic subsidence along the Sappington 

to West Yellowstone profile, based on surveys in 1934 and 

1959, and 32 cm of postseismic uplift that occurred between 

surveys in 1959 and 1987.14

Reilinger (1986) fit an exponential curve to the changing height relationship between two benchmarks 1.2 km apart 

on the Norris to West Yellowstone profile, deduced a 10-year 

decay period for postseismic uplift, and attributed postseismic 

motion to viscoelastic response of the asthenosphere. Holdahl and Dzurisin (1991) fit the changing height relationships 

among more marks along the same profile and deduced a 

2-year decay period for the motion, which according to their 

preferred model essentially had stopped by 1970. They noted 

that the lack of any additional uplift at West Yellowstone, relative to Norris, after 1975 (based on surveys in 1983 and 1987), 

precludes a decay period longer than 4 years. The discovery by 

Wicks and others (2006) of a deformation source beneath the 

north caldera rim near Norris complicates the interpretation 

of postseismic motion associated with the 1959 Hebgen Lake 

earthquake, because the north rim source could have caused 

undocumented height changes at Norris, relative to West 

 12The public road system within Yellowstone National Park has changed 

since the 1923 leveling survey such that several segments of the leveling network now lie along closed service roads or trails. A notable example is a line 

along an abandoned road—now the Mary Mountain trail—across the central 

part of the Yellowstone caldera by way of Mary Lake (figs. 1 and 6).

 13Ms

 refers to surface-wave magnitude, which is based on the amplitude 

of Rayleigh surface waves measured at a period near 20 s. Depending on 

how they were determined, earthquake magnitudes are expressed as local 

magnitude (ML), body wave magnitude (Mb, based on the amplitude of P 

body-waves), or moment magnitude (Mw, based on the seismic moment of the 

earthquake). Unless specified otherwise, earthquake magnitude M refers to 

local magnitude.

14The 1934–59 result includes any preseismic displacements that might 

have occurred but were not documented.
12 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

SL

Monida

Marysville

LF 

WTJ

CJ

ML

IL

OF

NJ

TJ

FBJ LB

MJ

Mammoth

Livingston

Bozeman

Sappington

WYOMING

MONTANA

IDAHO

Yellowstone National Park

CALDERA RIM

111 ̊30' 

44 ̊

44 ̊30'

45 ̊30'

45 ̊

46 ̊

111 ̊ 110 ̊30' 110 ̊W

South Entrance

East

Entrance

North Entrance

Northeast Entrance

West Yellowstone

(West Entrance)

Level line

Benchmark

1050 20 KILOMETERS

0 5 1O MILES

Figure 5. Map of level lines in the Yellowstone-Hebgen Lake region, after Holdahl and Dzurisin (1991). See text for 

dates of leveling surveys along various lines. TJ, Tower-Roosevelt Junction; CJ, Canyon Junction; NJ, Norris Junction; 

MJ, Madison Junction; ML, Mary Lake; OF, Old Faithful; IL, Isa Lake; WTJ, West Thumb Junction; FBJ, Fishing Bridge 

Junction; LB, Lake Butte; SL Sylvan Lake; LF, Lewis Falls.
Results of Leveling Surveys and InSAR Observations, 1923–2008 13

Yellowstone, during the period of postseismic deformation. 

The line between West Yellowstone and Norris was measured 

in 1923, 1975–77, 1983, 1987, and 2004—not frequently 

enough to distinguish among relative height changes associated with the 1959 Hebgen Lake earthquake, the 1975 Yellowstone Park earthquake (see below), or the north rim source. 

This ambiguity cannot be resolved in hindsight, so we are left 

with the less than satisfying necessity of neglecting tectonic 

stress changes as a primary cause of surface displacements, 

except perhaps indirectly by enhancing or restricting fluid flow 

out of the caldera (Wicks and others, 2006). Our concern over 

this shortcoming is lessened for two reasons. First, Pelton and 

Smith (1982) rejected tectonic stress as a likely cause of uplift 

that occurred in the caldera between leveling surveys in 1923 

and 1975–77. They cited the predominance of normal Quaternary faults and earthquake fault-plane solutions indicating 

near vertical P axes (inferred axes of maximum compressive 

stress) as evidence that the dominant horizontal stress in the 

Yellowstone region is tensional, which is inconsistent with 

surface uplift. In addition, InSAR observations (Wicks and 

others, 1998, 2006; Chang and others, 2007, 2010) require 

a forcing function other than stress changes associated with 

major earthquakes to explain the complexity and rapidity of 

changing deformation patterns at the Yellowstone caldera.

1975 M 6.1 Yellowstone Park Earthquake

Another major earthquake, the M 6.1 Yellowstone Park 

event, occurred on June 30, 1975, beneath the north rim of 

the caldera near the Norris Geyser Basin (Pitt and others, 

1979) (fig. 4). In August 1975, the USGS performed a firstorder resurvey of the line from Madison Junction to Mammoth Hot Springs Junction by way of Norris Junction (figs. 

1 and 6). Comparison of the 1923, 1960, and 1975 surveys 

indicated large relative height changes between Madison and 

Norris—more than 200 mm during 1923–60 and about 120 

mm during and 1960–75—with Norris down with respect 

to Madison in both cases. Pitt and others (1979) urged that 

the results be interpreted with caution, because they could 

not separate effects of the 1959 and 1975 earthquakes. In an 

attempt to resolve that ambiguity, the USGS and University 

of Utah releveled the rest of the 1923 network of lines within 

the park, except for the line across Mary Mountain by way of 

Mary Lake (figs. 1, 5, and 6), to first-order standards during 

the summer of 1976. The 1976 survey produced a major surprise—parts of the caldera floor had risen more than 700 mm 

since the 1923 survey—and more uncertainty. In this case, 

lack of data from the Mary Lake line, between the Mallard 

Lake and Sour Creek resurgent domes, left open the possibility that an even greater amount of uplift might have occurred 

there. That question was addressed by releveling the line 

along the Mary Mountain trail during the summer of 1977, 

which completed the 1975–77 resurvey of the entire 1923 

network. Those results are discussed in the following section.

It is noteworthy that Hamilton (1987) saw no evidence in 

data from USGS water-level gages along the shore of Yellowstone Lake for any significant change in the long-term trend of 

ground tilt at the time of the 1959 Ms

 7.5 Hebgen Lake earthquake; that is, 0.4 microradians per year (μradian/yr) up to the 

northeast. That trend is consistent with uplift of the Sour Creek 

dome relative to the south caldera rim, as shown by the 1923 

and 1975–77 leveling surveys. On the other hand, Hamilton 

(1987) noted apparent subsidence of the Sour Creek dome and 

concurrent uplift of the Mallard Lake dome at the time of the 

1975 M 6.1 Yellowstone Park earthquake. In hindsight this may 

have been the earliest indication of simultaneous uplift and subsidence at different locations within the caldera, as demonstrated 

clearly by InSAR images for the period August 1995–September 1996 (see section on Variable Subsidence and Resumption 

of Uplift from InSAR Observations, 1992–1997).

Discovery of Caldera Floor Uplift, 1923 to 1975–

1977

Pelton and Smith (1979, 1982) compared the 1923 and 

1975–77 leveling results, which indicated that the central part 

of the caldera floor had risen substantially during the interval between surveys. Maximum uplift of 726±22 mm, which 

corresponds to an average rate of 14±1 mm/yr, was measured 

at benchmark B11 1923 near Le Hardys Rapids.15 Differential 

vertical motions along the level line across Mary Mountain were 

small, indicating that the zone of maximum uplift was elongate 

and coincided with the long axis of the caldera (figs. 1 and 6; 

table 1). The deformation field mimicked the shape of the elliptical caldera, except on the north and west sides, where measurable amounts of surface displacement extended well outside the 

caldera rim. Those areas had risen 200–300 mm with respect to 

K12 1923, the designated reference benchmark near the southeast caldera rim (fig. 6). The Norris area had subsided a similar 

amount with respect to its immediate surroundings, such that its 

net displacement relative to K12 1923 was near zero.

Pelton and Smith (1982) noted the possibility of coseismic 

and postseismic displacements associated with the 1959 Ms

 7.5 

Hebgen Lake earthquake and 1975 M 6.1 Yellowstone Park 

earthquake, as had been suggested by Myers and Hamilton 

(1964) and Reilinger and others (1977). Referring to the 

newly discovered uplift inside the caldera, Pelton and Smith 

(1979, p. 1179) wrote: “The most likely cause of this rapid 

and unusually large surface deformation is a recent influx of 

partially molten material to a location within the crust beneath 

15Pelton and Smith (1979, 1982) calculated vertical displacements with 

respect to K12 1923, a reference benchmark about 8 km outside the southeast 

caldera rim (fig. 6). Their choice makes sense, because the leveling results 

show that any differential vertical ground motions outside the caldera were 

small compared to those in the central part of the caldera, and because K12 

1923 is far removed from any effects of the 1959 Ms 7.5 Hebgen Lake earthquake or 1975 M 6.1 Yellowstone Park earthquake.
14 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

A

Ring fracture

Leveling line

Benchmark

NorrisMammoth

Corridor

Mary Mountain Trail

Trout Creek

Mount

Washburn

Canyon

Junction

Lake Butte

Indian Pond

Fishing Bridge

Pelican

Cone

Madison Junction

Lewis

Falls

Norris

Junction

West Thumb

Junction

Old

Faithful

Isa

Lake

LeHardys Rapids

C10

Y9

DA3

B11

A11

F12

36MDC K12

Z9

M10

P10

H13

N13

L13

M13

F10

51MDC

R366

J11

G11

11MDC

B

Mallard

 Lake

Dome

Sour

Creek

Dome

CALDERA

RIM

0 5 10 20 KILOMETERS

0 5 10 MILES

Figure 6. A, Map of Yellowstone National Park leveling network and vertical surface displacements measured by comparison 

of leveling surveys in 1923 and 1975–77. Uplift contours in millimeters are from Pelton and Smith (1982). Salmon-colored 

highlighting indicates Norris–Mammoth corridor. B, Enlargement of Yellowstone caldera area shown in A. Blue line shows level 

line between Lake Butte and Mount Washburn, which was measured in 1923, 1975–77, in whole or part each year from 1983 to 

1993, and in 1995, 1998, 2005, and 2007. The line crosses the maximum-uplift axis, which is located within the 700-mm contour 

line in A, near DA3 and Le Hardys Rapids. Green line shows the level line between Lewis Falls and Madison Junction, which 

was measured in 1923, 1975–77, 1986, and 1987. It crosses the maximum-uplift axis near 51 MDC and Isa Lake. A level line that 

follows the Mary Mountain trail was measured in 1923, 1977, and 1987.
Results of Leveling Surveys and InSAR Observations, 1923–2008 15

Table 1. Average vertical displacement rates near Le Hardys Rapids, Yellowstone National Park, from leveling surveys.

[The 1923–1976 result is for benchmark B11 1923 relative to K12 1923. The 1998–2005 and 2005–2007 results are for DA3 1934 relative to CVO 

84-24. All other results are for DA3 1934 relative to 36 MDC 1976. Benchmarks B11 1923 and DA3 1934 are 2 km apart and near Le Hardys Rapids. 

Reference benchmarks K12 1923, CVO 84-24, and 36 MDC 1976 are within 11 km of one another and near Lake Butte on the southeast rim of the Yellowstone caldera. Uplift and subsidence profiles in the northeast part of the caldera (fig. 10) show that maximum displacements consistently occur near 

Le Hardys Rapids and that relative displacements among marks near Lake Butte are small. Therefore, these values are a reasonably consistent measure 

of the displacement rate at Le Hardys Rapids relative to Lake Butte. See text and references cited in the table for details]

Time period Average uplift (+) or subsidence (-) rate Source

1923–1976 +14±1 mm/yr Pelton and Smith (1979, 1982)

1976–1983 +23±1 mm/yr (see caption) Dzurisin and Yamashita (1987)

1983–1984 +16±5 mm/yr (see caption) Dzurisin and Yamashita (1987)

1976–1984 +22±1 mm/yr Dzurisin and Yamashita (1987)

Dzurisin and others (1990)

1984–1985 −2±5 mm/yr Dzurisin and others (1990)

1985–1986 −25±5 mm/yr Dzurisin and Yamashita (1987)

Dzurisin and others (1990)

Dzurisin and others (1994)

1986–1987 −35±5 mm/yr Dzurisin and others (1994)

1987–1988 −9±5 mm/yr Dzurisin and others (1994)

1988–1989 −11±5 mm/yr Dzurisin and others (1994)

1989–1990 −13±5 mm/yr Dzurisin and others (1994)

1990–1991 −11±5 mm/yr Dzurisin and others (1994)

1991–1992 −32±5 mm/yr Dzurisin and others (1994)

1992–1993 −13±5 mm/yr Dzurisin and others (1994)

1993–1995 −20±3 mm/yr Dzurisin and others (1999)

1995–1998 −8±3 mm/yr Dzurisin and others (1999)

1998–2005 +4±1 mm/yr This paper

2005–2007 +53±3 mm/yr This paper

Yellowstone National Park.” That conclusion has withstood 

the test of time—but not without considerable ongoing debate 

and elaboration (see Discussion section).

One of the authors of this report (Dzurisin) started working 

in Yellowstone National Park in 1983, about 4 years after Pelton 

and Smith (1979) published their discovery of rapid crustal 

uplift in the prestigious journal Science. Personal experience 

during 15 leveling surveys across the caldera from 1983 to 2007 

provides some insight into what must have been a puzzling and 

exciting experience for the surveyors and scientists involved in 

that discovery. During summer 1975, the line between Madison Junction and Mammoth Hot Springs Junction by way of 

Norris Junction was surveyed in response to the June 30, 1975, 

M 6.1 Yellowstone Park earthquake near Norris (figs. 1 and 6). 

Results showed that there had been substantial changes since 

the previous survey in 1923, which prompted leveling of two 

additional lines during summer 1976—(1) between Norris Junction and reference mark K12 1923 by way of Canyon Junction 

and Fishing Bridge Junction and (2) between Madison Junction 

and West Thumb Junction by way of the Old Faithful area. The 

first line includes benchmark B11 1923 near Le Hardys Rapids, 

where the 1976 results indicated more than 700 mm of uplift 

relative to K12 1923 since the 1923 survey. Mark B11 1923 is 

about 30 km along the level line from K12 1923, so 700 mm of 

uplift corresponds to an average gradient of more than 20 mm/

km—a number that might have seemed implausibly large at the 

time. One can imagine excitement mounting among members 

of the survey party as the discrepancy between their results 

and the 1923 survey continued to accumulate—100 mm, 300 

mm, and 700 mm—and increasing away from the epicenter of 

the 1975 earthquake! Could those huge differences possibly be 

real? After all, the level lines surveyed in 1975 and 1976 did not 

form a closed loop, so leveling error could not be eliminated as 

a possible explanation of the differences. The definitive answer 

was not found until summer 1977, when three additional lines 

were surveyed—(1) between Canyon Junction and the Northeast Entrance by way of Tower-Roosevelt Junction, (2) between 

Fishing Bridge Junction and the South Entrance by way of West 

Thumb Junction, and (3) across Mary Mountain. Together, the 

1975, 1976, and 1977 surveys form a closure loop around the 

southern part of the Grand Loop Road; that is, from Madison 

Junction to Norris Junction (surveyed in 1975), Norris Junction 
16 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

to Fishing Bridge Junction by way of Canyon Junction (1976), 

Canyon Junction to West Thumb Junction (1977), and back to 

Madison Junction by way of the Old Faithful area (1976). Even 

though the surveys spanned three successive summers, closure 

around the loop was within the acceptable limit for first-order 

surveys. This meant two things: (1) any movement that might 

have occurred during the course of the surveys was small, so the 

results could be combined and analyzed as a single large survey, 

and (2) more importantly, it was clear that something remarkable 

had happened since the 1923 survey—parts of the caldera floor 

had risen more than 700 mm at an average rate of 14 mm/yr.

Pelton and Smith (1982) published a thorough analysis of 

the Yellowstone leveling results in 1982, not long after two other 

remarkable happenings in the world of volcanology—the May 

18, 1980, eruption of Mount St. Helens and a strong earthquake 

swarm that same month at the Long Valley caldera, California. 

Until that time, the discovery of crustal uplift in Yellowstone 

was regarded by many volcanologists as important but not pressing. The 1980 eruption at Mount St. Helens, followed a week 

later by the intense earthquake swarm at Long Valley, caused 

some volcanologists to connect the dots and wonder about the 

possibility of an eruption at Long Valley in our lifetimes—an 

idea that would have seemed far-fetched just a few months 

earlier. The Long Valley caldera is similar in many respects to its 

somewhat larger and younger cousin, the Yellowstone caldera. 

Pelton and Smith’s (1982) paper appeared in the Journal of 

Geophysical Research amidst a buzz in the volcanological community over the ongoing eruption at Mount St. Helens, where a 

dacite lava dome was growing episodically on the floor of the 

1980 crater, and persistent unrest at Long Valley that led to the 

USGS Notice of Potential Volcanic Hazard in May 1982. The 

stage was set for a quarter century of intensive study at Yellowstone—the exciting period of discovery initiated by Pelton and 

Smith (1979, 1982) and discussed in this report.

Continuing Caldera Floor Uplift, 1976–1984

As mentioned earlier, the 1983 leveling survey across the 

northeast part of the caldera was not completed owing to the 

onset of winter weather. Even so, the partial survey showed 

that uplift had continued since the 1976 survey. That conclusion is based on the shape of the relative displacement profile 

for 1976–83, which shows that the central part of the level 

line near Le Hardys Rapids rose 78±4 mm (benchmark DA3 

1934) with respect to the south end (32 MDC 1976) and 46±4 

mm with respect to the north end (F11 A) (fig. 7). Although 

the total uplift at Le Hardys Rapids relative to K12 1923, the 

reference mark for the 1976 survey, was not measured by the 

1983 survey, we are confident that an appreciable amount of 

uplift did occur in the central part of the caldera between 1976 

and 1983 (see below).

That confidence is reinforced by results of the 1984 

survey, which indicate that DA3 1934 near Le Hardys Rapids 

rose 177±6 mm with respect to 36 MDC 1976 near Lake Butte 

during 1976–84—at an average rate of 22±1 mm/yr (fig. 7 and 

16Benchmark 36 MDC 1976 is along the East Entrance Road, 3 km 

southeast of the summit of Lake Butte and 6 km northwest of K12 1923. 

Lake Butte is an intrusive mass of Tertiary age along the southeast rim of the 

Yellowstone caldera. Mark 36 MDC 1976 was used as a reference mark for 

leveling surveys from 1984 to 1998. It was not recovered for the 2005 survey 

and is presumed destroyed by roadwork. Nearby mark CVO 84-24 was used 

as reference for the 2005 and 2007 surveys. Results of the 1923 and 1975–77 

surveys (fig. 5) indicate that any relative motions among K12 1923, 36 MDC 

1976, and CVO 84-24 were probably small relative to motions in the central 

part of the caldera.

table 1) ( Dzurisin and Yamashita, 1987).16 DA3 1934 is only 

about 2 km north of B11 1923 and 36 MDC 1976 is only about 

6 km northwest of K12 1923, the reference mark for the 1923 

and 1975–77 surveys. So the average uplift rate at DA3 1934 

during 1976–84 (22±1 mm/yr with respect to 36 MDC 1976) 

can be compared in a general way to the average uplift rate at 

B11 1923 during 1923–76 (14±1 mm/yr with respect to K12 

1923). Even with the ambiguity introduced by using different 

reference marks, it seems clear that the average uplift rate was 

greater during 1976–84 than during 1923–76. On the basis of 

results of subsequent leveling surveys and InSAR observations 

(see below), we suspect that this is due in part to one or more 

unrecorded episodes of subsidence during 1923–76.

The shapes of the 1976–83 and 1976–84 displacement 

profiles are very similar where they overlap for 17 km in the 

central part of the caldera. Both profiles are reminiscent of 

the uplift pattern for the caldera as a whole that Pelton and 

Smith (1979, 1982) deduced from the 1923 and 1975–77 

surveys. This observation led Dzurisin and Yamashita (1987) 

to extrapolate the 1976–83 vertical displacement profile to 36 

MDC 1976, using the 1976–84 profile as a guide. The resulting maximum uplift value, 161±6 mm at DA3 1934 with 

respect to 36 MDC 1976 from 1976 to 1983, corresponds to an 

average uplift rate of 23±1 mm/yr during 1976–83 and 16±5 

mm/yr during 1983–4. Owing to uncertainty inherent in the 

extrapolation, we are not confident that the apparent difference 

in average uplift rates near Le Hardys Rapids for 1976–83 

and 1983–84 is real. On the other hand, the average uplift rate 

for 1976–84 (22±1 mm/yr) is more than 50 percent greater 

than for 1923–76 (14±1 mm/yr), and that difference is almost 

surely real. The difference might indicate that (1) uplift began 

sometime after 1923 and proceeded at a relatively constant 

rate, (2) the uplift rate increased over time, or (3) some more 

complicated history, which could have included pauses or 

periods of subsidence. The latter possibility was underlined by 

results of the 1985 survey.

A Pause in Vertical Surface Motion, 1984–1985

The 1985 leveling survey between Lake Butte and 

Canyon Junction produced a surprise—that is, the vertical displacement measured near Le Hardys Rapids (DA3 1934) with 

respect to the south caldera rim (36 MDC 1976) from September 1984 to September 1985 was essentially zero (−2±5 mm) 
Results of Leveling Surveys and InSAR Observations, 1923–2008 17

-50

-25

0

25

50

75

100

125

150

175

200 3,300

3,200

3,100

3,000

2,900

2,800

2,700

2,600

2,500

2,400

2,300

Elevation, in meters

Vertical displacement, in millimeters

11 MDC 1976

R366 1987

32 MDC 1976

36 MDC 1976

Not leveled in 1983

Alternate routes in 

1976 and 1984

DA3 1934

28 MDC 1976

LeHardys

Rapids

Ring Fracture Ring Fracture Topography

1976-1984

1976-1983

Topography

F11 A

Projected stadia distance, in kilometers

0 5 10 15 20 25 30 35

Figure 7. Diagram showing profiles of topography (black diamonds, bottom) and vertical displacements during 1976–83 (green 

circles) and 1976–84 (red squares) along the level line between Lake Butte and Canyon Junction by way of Le Hardys Rapids 

in Yellowstone National Park (Dzurisin and Yamashita, 1987). Error bars are one standard deviation from the combination of 

random leveling error and random-walk benchmark instability, as discussed in the text. Results of the smaller 1983 survey 

were extrapolated from 32 MDC 1976 to 36 MDC 1976, assuming that 36 MDC 1976 did not move up or down and using the 

shape of the 1923–76 uplift profile as a guide. Descriptions and photographs of benchmarks, which are indicated by arrows 

with labels, are included in appendix B in the interactive data files that accompany this report. The benchmark information is 

available online at http://pubs.usgs.gov/pp/1788/pp1788_benchmarks/.

(fig. 8 and table 1). None of the marks along the line moved 

by an amount that is statistically significant at two standard 

deviations. This was the first documented occurrence of a 

pause in caldera-floor uplift since the initial leveling survey in 

1923, although such occurrences could have gone unnoticed as 

a result of the sporadic nature of the surveys. Another surprise 

came just one year later.

Caldera-Floor Subsidence, 1986–1995

In 1986, leveling surveys were conducted along two lines 

that cross the long axis of the Yellowstone caldera—(1) the 

44-km line from Lake Butte to Canyon Junction in the northeast part of the caldera and (2) the 76-km line from Lewis 

Falls to Madison Junction, by way of West Thumb Junction 

and the Old Faithful area, in the southwest part (fig. 6). The 

lines skirt the southwest margins of the Sour Creek and Mallard Lake resurgent domes, respectively, and cross the axis of 

maximum uplift deduced by Pelton and Smith (1979, 1982). 

We were surprised to discover that since the 1985 survey 

mark DA3 1934 near Le Hardys Rapids had moved down 

by 25±5 mm with respect to reference mark 36 MDC 1976 

near the south caldera rim—the first documented occurrence 

of caldera-floor subsidence (fig. 8) (Dzurisin and Yamashita, 

1987; Dzurisin and others, 1990). In the southwest part of the 

caldera, mark 51 MDC 1976 near Isa Lake moved up 107±7 

mm with respect to reference mark N13 1923 near Lewis 

Falls, along the southwest caldera rim, between the surveys in 

1976 and 1986 (fig. 9).17 The average uplift rate for 1976–86 

at 51 MDC 1976, 11±1 mm/yr, was less than the rate at DA3 

1934 near Le Hardys Rapids for the same period. However, 

the rate was comparable to the maximum value of 14±1 mm/yr 
18 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

at B11 1923 (also near Le Hardys Rapids) for 1923–76 from 

Pelton and Smith (1979, 1982). Whether or not any subsidence 

occurred in the southwest part of the caldera during 1985–6, as 

it had in the northeast part, could not be determined until the 

next survey in 1987.

The discovery of subsidence in the northeast part of the 

caldera during 1985–86 led the USGS to contract to the National 

Geodetic Survey (NGS) for a complete resurvey of the Yellowstone network during summer 1987. Results for the level line 

between Lewis Falls and Madison Junction indicated that the 

area near Isa Lake in the west-central part of the caldera subsided 

13±6 mm (51 MDC 1976) with respect to the south caldera 

rim (N13 1923) from 1986 to 1987. Marks several kilometers 

farther east (that is, closer to the Mallard Lake resurgent dome), 

which were not surveyed in 1976, subsided as much as 25±6 mm 

during the same period (fig. 9). Along the other level line, Le 

Hardys Rapids moved down 35±5 mm (DA3 1934) with respect 

to the south rim (36 MDC 1976) during 1986–7. Within the 

uncertainty in the measurements, and allowing for the possibility of relative motion between 36 MDC 1976 and N13 1923, 

the maximum subsidence rates measured in the northeast and 

southwest parts of the caldera were essentially the same. In 1996, 

a partial survey across the southwest part of the caldera indicated 

that mark MEERTENS2 near Old Faithful subsided 149±8 mm 

with respect to B144 near West Yellowstone during 1987–96, 

at an average rate of 17±1 mm/yr. For the same period, the 

subsidence rate at MEERTENS2 with respect to USBPR 6802 at 

Madison Junction was 15±1 mm/yr (fig. 9).

Taken together, the leveling results for 1986–96 indicate that the uplift pattern reported by Pelton and Smith 

0

1984-1985

1984-1986

1984-1987

1984-1988

1984-1989

1984-1990

1984-1991

1984-1992

1984-1993

1984-1995

Topography

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❋

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-75

-50

-25

25

0 5 10 15 20 25 30 35

Lake

Butte

Canyon

Junction

LeHardys

Rapids

11MDC 1976 11MDC 1976

DA3 

1934

28 MDC 

1976

Ring Fracture Ring Fracture CVO 84-24 CVO 84-24 R366 1987

Vertical

Displacements

Topography

36 MDC 1976 

3,300

3,200

3,100

3,000

2,900

2,800

2,700

2,600

2,500

2,400

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Elevation, in meters

Projected stadia distance, in kilometers

Elevation, in meters

Vertical displacement, in millimeters

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❋❋

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❋❋

❋❋ ❋❋

❋

❋ ❋❋ ❋ ❋

❋

❋

❋

❋

❋

❋❋

❋

❋

❋

Figure 8. Diagram showing profiles of topography (black diamonds, bottom) and cumulative vertical displacements from 

1984 to 1995 along the level line between Lake Butte and Canyon Junction by way of Le Hardys Rapids in Yellowstone 

National Park. Data are from leveling surveys conducted in September of each year, except 1994. Error bars are one standard 

deviation from the combination of random leveling error and random-walk benchmark instability, as discussed in the text. 

Descriptions and photographs of benchmarks, which are indicated by arrows with labels, are included in appendix B in the 

interactive data files that accompany this report. The benchmark information is available online at http://pubs.usgs.gov/

pp/1788/pp1788_benchmarks/.

17Level lines between 36 MDC 1976 and N13 1923 were not measured in 

1986, so each of those marks is a local reference only for a line across the 

caldera floor (Lake Butte to Mount Washburn and Lewis Falls to Madison 

Junction, respectively). A tie between the cross-caldera lines was made next 

during a complete survey of the Yellowstone leveling network in 1987.
Results of Leveling Surveys and InSAR Observations, 1923–2008 19

-200

-150

-100

-50

0

50

100

150

0 10 20 30 40 50 60

Distance, in kilometers

Displacement, in millimeters

2,000

2,500

3,000

3,500

4,000

4,500

5,000

5,500

6,000

6,500

7,000

Elevation, in meters

Old

Faithful

Madison

Junction

Isa

Lake

West Thumb

Junction

1976-86

1986-87

1987-96

Topography

N13 1923

USBPR 6802

51 MDC 1976

MEERTENS2

Topography

Ring 

Fracture

Ring 

Fracture

Lewis

Falls

Figure 9. Diagram showing profiles of topography (black diamonds, bottom) and vertical displacements along the level 

line between Lewis Falls and Madison Junction by way of West Thumb Junction and the Old Faithful area in Yellowstone 

National Park from 1976 to 1986 (red squares), 1986 to 1987 (blue circles), and 1987 to 1996 (green triangles). Descriptions 

and photographs of benchmarks, which are indicated by arrows with labels, are included in appendix B in the interactive 

data files that accompany this report. The benchmark information is available online at http://pubs.usgs.gov/pp/1788/pp1788_

benchmarks/.

(1979, 1982) for 1923–76 reversed to one of subsidence, with 

approximately the same shape and average rate, sometime during 1984–86. During that interval, beginning in October 1985 

and continuing for more than 3 months, the largest earthquake 

swarm ever recorded in the Yellowstone region (excluding 

mainshock-aftershock sequences) occurred near Madison 

Junction just outside the northwest caldera rim (Waite and 

Smith, 2002). In the Discussion section of this report, we infer 

a causal relationship between that swarm and the onset of 

caldera-floor subsidence during 1984–86.

Leveling surveys between Lake Butte and Mount Washburn carried out each September from 1988 to 1995, except 

1994, showed progressive subsidence near Le Hardys Rapids at 

rates that ranged from 11±5 mm/yr to 32±5 mm/yr (fig. 8 and 

table 1). At DA3 1934, the average uplift rate during 1976–84 

was 22±1 mm/yr, and the average subsidence rate during 

1985–93 was 19±1 mm/yr—nearly identical displacement rates, 

but in opposite directions, during successive 8-year periods. The 

shapes of the uplift and subsidence profiles are nearly mirror 

images of each other (fig. 10), suggesting a common source—a 

topic to be addressed in the Discussion section.

Variable Subsidence and Resumption of Uplift 

from InSAR Observations, 1992–1997

No leveling surveys were conducted in Yellowstone 

National Park in 1996 or 1997, but InSAR observations that 

collectively span 1992–97 provided two additional surprises. 

Wicks and others (1998) presented interferograms produced 

from ERS-1 and ERS-2 C-band radar images that show a subsidence bowl in the northeast part of the caldera from August 

1992 to June 1993, an occurrence that is consistent with the 

results of leveling surveys across that part of the caldera in 

September 1992 and September 1993 (fig. 11). An interferogram for the period from June 1993 to August 1995 showed 

additional subsidence in the same area, as corroborated by a 

leveling survey in September 1995. However, the 1993–95 
20 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

interferogram showed an even greater amount of subsidence in 

the southwest part of the caldera, where no leveling had been 

done since 1987. This was a surprise because the pattern of 

uplift revealed by the complete surveys in 1923, 1975–77, and 

1987 was nearly symmetric; that is, the northeast and southwest 

parts of the caldera floor moved similar amounts. The 1993–95 

interferogram was the first clear indication of asymmetric uplift 

or subsidence within the caldera, although such a pattern could 

have escaped detection by sporadic leveling surveys.

InSAR images for the period August 1995–September 

1996 produced the second surprise: the northeast part of the 

caldera floor rose about 20 mm, while the southwest part continued subsiding. Such a bimodal pattern had not been detected 

by any previous leveling surveys, although Hamilton (1987) 

had reported a similar pattern based on differential water-level 

records from gages along the shore of Yellowstone Lake. His 

analysis showed that at the time of the 1975 M 6.1 Yellowstone 

Park earthquake there was subsidence of the northeast part of 

the caldera and concurrent uplift of the southwest part (that is, 

the inverse of the pattern revealed by InSAR in the mid-1990s). 

By June 1997, InSAR observations showed that more than 30 

mm of uplift had occurred throughout the central part of the 

caldera, in a pattern reminiscent of the 1923 to 1975–77 uplift 

as contoured by Pelton and Smith (1979, 1982; figs. 6 and 11). 

The onset of uplift was coincident with, or followed shortly 

after, a strong earthquake swarm during June–July 1995—the 

largest since the 1985 swarm that occurred in the same general 

area near the northwest caldera rim (USGS–University of Utah 

earthquake catalog). The InSAR results clearly showed that 

the caldera’s two resurgent domes sometimes move in opposite 

directions and, in this instance, uplift began at the Sour Creek 

dome and spread southwestward to the Mallard Lake dome.

-200

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-80

-60

-40

-20

0

20

40

60

80

100

120

140

160

180

200

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Projected stadia distance, in kilometers

Vertical displacement, in millimeters

2,300

2,400

2,500

2,600

2,700

2,800

2,900

3,000

3,100

3,200

3,300

Elevation, in meters

36 MDC 1976

11 MDC 1976 

32 MDC 1976

Alternate routes

in 1976 and 1984

CVO 84-24

R366 1987

28 MDC 1976

DA3 1934

27 MDC 1976

29 MDC 1976 

Lake

Butte

Canyon

Junction

LeHardys

Rapids

Ring Fracture Ring Fracture Topography

1976-1984

1984-1985

1985-1993

Topography

Figure 10. Diagram showing profiles of topography (black diamonds, bottom) and vertical displacements along the level line between 

Lake Butte and Canyon Junction by way of Le Hardys Rapids in Yellowstone National Park from 1976 to 1984 (red circles), from 1984 

to 1985 (green triangles), and from 1985 to 1993 (blue squares). Descriptions and photographs of benchmarks, which are indicated by 

arrows with labels, are included in appendix B in the data files that accompany this report. The benchmark information is available 

online at http://pubs.usgs.gov/pp/1788/pp1788_benchmarks/. The subsidence profile for 1985–93 is nearly a mirror image of the uplift 

profile for 1976–84. No significant deformation occurred during the intervening year, 1984–85.
Results of Leveling Surveys and InSAR Observations, 1923–2008 21

A B

C D

June 1993--August 1995

July 1995-June 1997

August 1992--June 1993

August 1995-September 1996

0 28 mm

0 2π

0 5 10 20 30 KILOMETERS

0 5 10 20 MILES

Figure 11. Four interferograms of the Yellowstone National Park region produced by Wicks and others (1998) from ERS-1 and 

ERS-2 radar images. The range of colors from violet to red, shown in the color bar, corresponds to one cycle of phase from 0 to 

2p (one fringe). Each fringe represents 28.3 mm of range change (surface displacement) between a point on the ground and the 

satellite. Solid white lines, caldera rim and two resurgent domes (A only). Dashed white line, boundary of Yellowstone National 

Park. Figure modified slightly from Dzurisin and others (1999). The first three interferograms are sequential, whereas the time 

periods covered by the third and fourth overlap. Note the pixilated nature of these early Yellowstone interferograms compared 

to more recent ones elsewhere in this report (figs. 12–15)—a result of improved data processing techniques. A, August 1992 

to June 1993. This image shows more than 30 mm of subsidence centered in the northeast half of Yellowstone caldera in the 

vicinity of the Sour Creek resurgent dome. B, June 1993 to August 1995. The center of subsidence (~40 mm maximum) has 

shifted in this image to the southwest half of the caldera near the Mallard Lake resurgent dome. C, August 1995 to September 

1996. The fringe pattern in the northeast half of the caldera in this image corresponds to ~20 mm of uplift in the vicinity of the 

Sour Creek dome (note the reversed color sequence toward the center of the fringe pattern relative to A and B). D, July 1995 to 

June 1997. In this image, uplift extends throughout the central part of the caldera (maximum ~30 mm). 
22 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

North Rim Uplift and Caldera-Floor Subsidence, 

1997–2002

InSAR observations contributed another piece to the 

Yellowstone deformation puzzle when an ERS-2 interferogram for the period from September 1996 to September 2000 

revealed a bulls-eye pattern of uplift centered along the north 

caldera rim in the vicinity of Norris Geyser Basin (Wicks and 

others, 2006, fig. 12). The uplifted area also had been identified by differencing results of GPS surveys made in 1995 and 

2000 (Puskas and others, 2007). The maximum range change 

from the interferogram was about 80 mm during 1996–2000. 

Wicks and others (2006) noted that the 1996–2000 interferogram indicated only slight subsidence in the central part of the 

caldera, where ~30 mm of uplift had occurred during 1996–97 

(Wicks and others, 1998). Therefore, the caldera floor between 

the two resurgent domes must have subsided more than 30 mm 

from 1997 to 2000, while the north rim was rising. If uplift 

of the north rim began in 1997, the average uplift rate there 

during 1997–2000 was about 27 mm/yr. The north rim source 

continued to inflate at a declining rate, while the caldera floor 

subsided, until 2002. An interferogram for the period from 

October 2002 to August 2003 showed slight subsidence of 

the north rim area, extending northward along the Norris–

Mammoth corridor, and no measurable motion of the caldera 

floor. The second known pause in caldera deformation was 

underway.

Seeing the Yellowstone InSAR results for the first time 

was akin to switching from black-and-white to color television or from two-dimensional (2 D) to three-dimensional (3 D) 

movies—the effect was dramatic. Our repeated leveling surveys 

across the northeast part of the caldera had yielded a series of 

vertical displacement profiles (fig. 8) that we hoped were a good 

representation of the caldera deformation field as a whole. The 

August 1995–September 1996 interferogram showed in the 

blink of an eye that this was not the case, that is, the northeast 

part of the caldera had gone up, while the southwest part had 

gone down. Even more striking was the bulls-eye pattern of 

uplift along the north caldera rim that appeared in the September 1996–September 2000 interferogram. Suddenly the situation 

was much more complicated and intriguing than we had thought 

since the discovery of uplift in the mid-1970s. The InSAR 

results heralded the start of a new chapter in the Yellowstone 

deformation story, and this one was in color!

Creation of the Yellowstone Volcano 

Observatory, May 2001

It was during a brief respite in caldera deformation that 

scientists and managers of the USGS, University of Utah, and 

Yellowstone National Park came together to establish the Yellowstone Volcano Observatory (YVO) in May 2001 (http://

volcanoes.usgs.gov/yvo/). YVO is an instrument-based monitoring facility designed for observing volcanic, hydrothermal, and 

earthquake activity in the Yellowstone National Park region. 

The monitoring primarily involves a seismic network operated 

by the University of Utah Seismograph Stations under a cooperative funding agreement with the USGS and with ancillary 

support from the National Science Foundation (NSF). YVO also 

operates a network of CGPS stations in the park that complements a network of CGPS stations and borehole strainmeters 

operated by the EarthScope Plate Boundary Observatory (PBO, 

http://pboweb.unavco.org/). In addition, campaign GPS, leveling, microgravity, and InSAR observations have been carried 

out by scientists of both the University of Utah and the USGS 

since YVO was created in 2001. Among these activities has 

been the study of ground deformation that we discuss here, a 

study which has evolved over time and is expected to continue 

under YVO auspices for the foreseeable future.

Rapid Uplift of the Sour Creek Dome and North 

Rim Subsidence, 2004–2008

The period of relative quiescence that began in 2002 

ended in remarkable fashion less than two years later. Chang 

and others (2007) reported that, during the first 6 months of 

2004, data from CGPS stations indicated 10–20 mm of subsidence of the Sour Creek dome and a similar amount of uplift of 

the north rim near Norris. Then, starting in July 2004, the caldera motion suddenly reversed to uplift at rates as high as ~70 

mm/yr (an historically unprecedented value) at CGPS station 

WLWY near the eastern margin of the Sour Creek dome. The 

Norris area began to subside about 3 months later. Chang and 

others (2007) noted that the entire caldera floor began to rise 

nearly simultaneously in mid-2004. This is in contrast to the 

observation by Wicks and others (1998) that during 1995–96 

uplift started at the Sour Creek dome—at a time when the 

Mallard Lake dome was subsiding—then spread to the entire 

caldera floor by 1996–97.

A stack of two ENVISAT IS2-mode18 interferograms that 

collectively span the period from September 2004 to August 

2006 depict more clearly the spatial pattern of caldera uplift 

and north rim subsidence (fig. 13). With reference to the 

stacked interferogram, Chang and others (2007, p. 953) wrote:

 The inflation increases symmetrically toward the 

caldera center about the long axis (northeast-southwest), with the highest rate of ~7 cm/year at the Sour 

Creek dome being three to five times faster than uplift 

rates in 1923–1984 and 1995–1997 * * *. The Norris 

18The ASAR instrument aboard ENVISAT is capable of multiple, interleaved imaging modes. IS2 refers to one such mode, which is characterized 

by an incidence angle of 21.5 degrees from vertical and either HH or VV 

polarization. HH refers to a data acquisition mode in which radar pulses with 

horizontal polarization are transmitted and received by ASAR; VV refers 

to the case in which the pulses are transmitted and received with vertical 

polarization. Mode IS2, VV is suitable for interferometry and similar to the 

imaging mode of ERS-1 and ERS-2.
Results of Leveling Surveys and InSAR Observations, 1923–2008 23

0

0 2π Range Change

28.3 mm

NGB

A

C D

1996-2000

2001-2002 2002-2003

B 2000-2001

0 5 10 20 30 KILOMETERS

0 5 10 20 MILES

Figure 12. Four interferograms superimposed on digital terrain showing range changes (ground deformation) at the 

Yellowstone caldera, Yellowstone National Park, between 1996 and 2003. A color change from violet to blue to green to 

yellow to red, shown in the color bar, marks an increase in the range (distance from the satellite to points on the ground) of 

28.3 mm. White dots represent epicenters of earthquakes recorded during the time interval spanned by each interferogram. 

White outline shows caldera rim. The interferograms were generated using European Space Agency ERS-2 data and the 

two-pass method of interferometry (Massonnet and Feigl, 1998). A, Summer 1996 to summer 2000 interferogram. Concentric 

pattern of fringes indicates uplift along the north caldera rim. Although the caldera floor appears to have subsided only 

slightly, this period includes about 30 mm of caldera-wide uplift from 1995 to 1997 (Wicks and others, 1998). Therefore, more 

than 30 mm of subsidence of the caldera floor occurred between the Mallard Lake and Sour Creek resurgent domes from 

1997 to 2000, while the north rim rose about 80 mm. B, Summer 2000 to summer 2001 interferogram showing continuing uplift 

along the north caldera rim and subsidence of the caldera floor. C, Summer 2001 to summer 2002 interferogram. Uplift along 

the north rim continued and extended northward along the Norris–Mammoth corridor, while the caldera floor continued to 

subside at a declining rate. Arrow labeled NGB marks the location of Norris Geyser Basin. D, Summer 2002 to summer 2003 

interferogram showing slight subsidence of the north rim area, extending northward along the Norris–Mammoth corridor, 

and no measurable motion of the caldera floor. Figure from Wicks and others (2006).
24 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

–111 ̊30' –111 ̊00' –110 ̊30' –110 ̊00'

44 ̊00'

44 ̊30'

45 ̊00'

0

0 2π Range Change

28.3 mm

1050 20 KILOMETERS

0 5 1O MILES

Figure 13. Stack of two deformation interferograms of the Yellowstone caldera, Yellowstone National Park, superimposed on digital 

terrain map, that collectively span the period from September 22, 2004, to August 23, 2006. The constituent interferograms were 

generated using European Space Agency Envisat IS2 mode data and the two-pass method of interferometry (Massonnet and Feigl, 

1998). The range of colors from violet to red, shown in the color bar, corresponds to one cycle of phase from 0 to 2p (one fringe). Each 

fringe represents 28 mm of range change (surface displacement) between a point on the ground and the satellite. The line-of-sight 

(LOS) displacements represented by fringes in the interferogram indicate total uplift of about 110 mm in the west part of the caldera 

(Mallard Lake dome), uplift of about 150 mm in the east part of the caldera (Sour Creek dome), and subsidence of 60–70 mm near the 

Norris Geyser Basin. Yellow lines, major roads.
Results of Leveling Surveys and InSAR Observations, 1923–2008 25

–111 ̊30' –111 ̊00' –110 ̊30' –110 ̊00'

44 ̊00'

44 ̊30'

45 ̊00'

0 10 20

Kilometers 0

0 2π Range Change

28.3 mm PBO CGPS

YVO CGPS

1050 20 KILOMETERS

0 5 1O MILES

Figure 14. ENVISAT IS2 mode interferogram of the Yellowstone caldera, Yellowstone National Park, superimposed on digital terrain 

map. Interferogram was produced from images acquired on August 3, 2005, and July 4, 2007, showing uplift of the Yellowstone 

caldera (Sour Creek and Mallard Lake resurgent domes) and subsidence of the north rim in the vicinity of Norris Geyser Basin. The 

range of colors from violet to red, shown in the color bar, corresponds to one cycle of phase from 0 to 2p (one fringe). Each fringe 

represents 28 mm of range change (surface displacement) between a point on the ground and the satellite. Red triangles, continuous 

GPS (CGPS) stations operated by the Yellowstone Volcano Observatory (YVO). White triangles, continuous CGPS stations operated by 

the Plate Boundary Observatory (PBO). Yellow lines, major roads. Compare this figure to the 2004–6 interferogram in figure 13 and to 

the 2007–8 interferogram in figure 15. Caldera-floor uplift and north-rim subsidence continued at a slightly lower average rate during 

2005–7 and 2007–8 than during 2004–6.
26 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

–111 ̊30' –111 ̊00' –110 ̊30' –110 ̊00'

44 ̊00'

44 ̊30'

45 ̊00'

0

0 2π Range Change

28.3 mm 0 10 20 KILOMETERS

0 5 1O MILES

Figure 15. ENVISAT IS2 mode interferogram of the Yellowstone caldera, Yellowstone National Park, superimposed on digital terrain 

map. Interferogram was produced from images acquired on September 12, 2007, and August 27, 2008. The range of colors from 

violet to red, shown in the color bar, corresponds to one cycle of phase from 0 to 2p (one fringe). Each fringe represents 28 mm of 

range change (surface displacement) between a point on the ground and the satellite. Uplift during the period extended across 

most of the caldera floor, including the Mallard Lake and Sour Creek resurgent domes, and reached a peak value of about 50 mm at 

the Sour Creek dome. At the same time, an area along the north caldera rim in the vicinity of Norris Geyser Basin subsided about 8 

mm. Yellow lines, major roads.
Results of Leveling Surveys and InSAR Observations, 1923–2008 27

-30

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-10

0

10

20

30

40

50

60

70

80

90

100

110

120

130

0 5 10 15 20 25 30 35

Projected stadia distance, in kilometers 

Vertical displacement, in millimeters

2,300

2,400

2,500

2,600

2,700

2,800

2,900

3,000

3,100

3,200

3,300

Elevation, in meters

36 MDC 1976

R366 1987

DA3 1934

29 MDC 1976 

LEHARDY

11 MDC 1976

CVO 84-24

Ring Fracture Ring Fracture Topography

Lake

Butte

Canyon

Junction

LeHardys

Rapids

1995-1998

2005-2007

1998-2005

Topography

Figure 16. Diagram showing profiles of topography (black diamonds, bottom) and vertical displacements along the level line between 

Lake Butte and Canyon Junction by way of Le Hardys Rapids in Yellowstone National Park from 1995 to 1998 (red circles), from 1998 

to 2005 (green triangles), and from 2005 to 2007 (blue squares). Descriptions and photographs of benchmarks, which are indicated by 

arrows with labels, are included in appendix B in the interactive data files that accompany this report. The benchmark information is 

available online at http://pubs.usgs.gov/pp/1788/pp1788_benchmarks/.

subsidence is ~3 cm/year, more than two times greater 

than the 1996–2002 uplift rate in this area.

The pattern of caldera-floor uplift and north-rim subsidence persisted at lower rates during 2006–7, as shown by 

comparison of figures 13 and 14 for 2004–6 and 2005–7, 

respectively. An Envisat IS2 interferogram for the interval 

September 12, 2007, to August 27, 2008, indicated ~50 mm 

of peak uplift centered on the Sour Creek dome and ~8 mm 

of subsidence along the north caldera rim (fig. 15). The uplift 

rate at CGPS station WLWY on the Sour Creek dome was 

40–50 mm/yr during 2007–8; by September 2009, uplift there 

had slowed to the point that it may have stopped. The station 

had undergone a total of about 230 mm of uplift since mid2004 (http://volcanoes.usgs.gov/yvo/activity/archive/2009.

php). Minor subsidence in the Norris area also seemed to have 

stopped by the third quarter of 2009, as indicated by CGPS 

station NRWY. Up-to-date time-series data from CGPS stations 

in the Yellowstone network, including WLWY and NRWY, are 

available at http://www.uusatrg.utah.edu/ts_ysrp.html.

Results of leveling surveys conducted along the line 

between Lake Butte and Mount Washburn in 1998, 2005, and 

2007 are consistent with the InSAR and CGPS observations 

described above (fig. 16). During 1995–98, DA3 1934 near Le 

Hardys Rapids rose 24±5 mm with respect to 36 MDC 1976. 

Recall that InSAR observations indicated more than 30 mm 

of uplift of the Sour Creek dome during 1995–7 and a slightly 

greater amount of subsidence during 1997–2000. The 1995–8 

leveling results indicate that most of the subsidence near Le 

Hardys Rapids occurred after September 1998. Note from figure 16 that the shape of the 1995–98 uplift profile differs from 

all previous ones, in that the greatest amount of uplift occurred 

at the north end of the line near Canyon Junction—29±7 

mm at 11 MDC 1976. About half of that amount reasonably 

could be attributed to leveling error, but the InSAR image for 

1996–2000 (fig. 12) suggests another explanation. The northrim uplift feature that was centered near Norris was extensive enough to have affected the northern half of the level 

line, which would account for some of the uplift at Canyon 
28 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Junction. We suspect that the shape of the 1995–98 uplift profile reflects the combined effects of two inflation sources, one 

beneath the Sour Creek dome and the other beneath the north 

caldera rim.

The shape of the 1998–2005 uplift profile is nearly identical to that of the 1995–98 profile (fig. 16). InSAR observations 

indicate that the north-rim area rose as much as 80 mm from 

1996 to 2000, while the Sour Creek dome subsided slightly. 

Uplift of the north rim area reached a maximum of about 120 

mm during 2000–2002, while the area between both resurgent domes subsided as much as 40 mm (Wicks and others, 

2006). The pattern changed to one of strong uplift of the Sour 

Creek dome starting in mid-2004, accompanied by subsidence 

of the north-rim area starting a few months later. From the 

1998–2005 leveling results, we conclude that (1) the amount 

of uplift of the Sour Creek dome from mid-2004 to September 

2005 exceeded the amount of subsidence there from September 1998 to mid-2004, which accounts for net uplift at Le 

Hardys Rapids during 1998–2005, and (2) net uplift in the 

Canyon Junction area during 1998–2005 reflects the combined 

effects of north-rim uplift during 1998–2002 and Sour Creek 

dome uplift starting in mid-2004.

The 1995–98 and 1998–2005 leveling results for the 

Canyon Junction area are reminiscent of the results for 1923 to 

1975–77 as contoured by Pelton and Smith (1979, 1982). Note 

from figure 6 that Lake Butte and Canyon Junction are near the 

0-mm and 200-mm uplift contours, respectively, and that uplift 

by 100–200 mm is shown extending well north of the north 

caldera rim. The only exception is a relative subsidence bowl 

centered near Norris Junction that might have been associated 

with the 1959 Hebgen Lake earthquake, the 1975 Yellowstone 

Park earthquake, or the north-rim deformation source identified 

by InSAR. Why such a broad area between Canyon Junction, 

Madison Junction, and Mammoth Hot Springs Junction should 

experience uplift relative to the south caldera rim remains an 

enigma and a topic worthy of further investigation.

The leveling results for 2005–7 are more straightforward, 

because they can be explained almost entirely by inflation of 

the Sour Creek dome source starting in mid-2004 (fig. 16). 

Uplift at DA3 1934 occurred at an average rate of 53±3 mm/yr 

during 2005–7, which is 3.8 times the average rate at nearby 

B11 1923 during 1923–76 (14±1 mm/yr) and 2.4 times the 

average rate at DA3 1934 during 1976–84 (22±1 mm/yr). 

Minor subsidence at the north end of the line (11±7 mm at 

11 MDC 1976) could be a result of deflation of the north-rim 

source or leveling error.

It is interesting to note that the average uplift rate at B11 

1923 near Le Hardys Rapids for 1976–2007, including the 

period of subsidence from 1985 to 1995, was 4±1 mm/yr. If 

caldera subsidence during 1985–95 had gone undocumented, 

we would know only that the average rate of caldera uplift 

had decreased markedly since the 1975–77 leveling survey. 

We actually know much more, thanks to frequent leveling 

surveys, CGPS monitoring, and repeated InSAR observations. 

Now the challenge is to translate that knowledge into better 

understanding of the causes and implications of ground motion 

at Yellowstone and other restless calderas.

Discussion

Considerable progress has been made in documenting the 

spatial and temporal patterns of ground deformation at the Yellowstone caldera since the first repeat leveling survey in 1975–

77. The 1923 and 1975–77 surveys called attention to rapid 

crustal uplift at Yellowstone, but the surveys were separated 

by too many years and the leveling network was too sparse to 

provide a detailed picture of the deformation pattern. The first 

limitation was addressed by 15 leveling surveys between 1983 

and 2007, which provided the first indication of caldera-floor 

subsidence and showed that uplift can reverse to subsidence 

in 2 years or less. However, those surveys were limited to two 

unconnected lines across the northeast and southwest parts 

of the caldera, so they could not shed much additional light 

on the spatial pattern of deformation. InSAR observations 

resulted in a breakthrough in that regard, first by identifying 

distinct deformation sources beneath the two resurgent domes 

and then by revealing temporal changes across the caldera 

over time scales too short to be captured by leveling. Another 

milestone was the discovery by InSAR of a third deformation 

source beneath the north rim of the caldera, which seems to 

interact with the other sources over time scales as short as a 

few months. Finally, CGPS stations have measured surface 

velocities two to three times greater than those captured by 

annual leveling surveys and have shown that changes in the 

deformation mode can occur across the entire caldera almost 

simultaneously (Chang and others, 2007, 2010).

Persistent snow cover at Yellowstone precludes leveling 

surveys or useful InSAR observations for more than 6 months 

each year, so the temporal resolution of these techniques is 

about 1 year. CGPS observations, on the other hand, provide 

temporal resolution on the order of 24 hours or less and represent a major step forward in monitoring ground deformation.

There has been parallel progress in understanding the 

causes of ground deformation at Yellowstone, but it has been 

less dramatic than the observational breakthroughs provided 

by leveling, CGPS, and InSAR. In a general way, Pelton and 

Smith (1979, 1982) anticipated the current view that the dominant uplift mechanism is basaltic intrusion into the mid-crust. 

They speculated that the intrusion responsible for historical 

uplift must have been “* * * recent (not more than 500 years 

ago) * * *” (Pelton and Smith, 1982, p. 2758). They also 

considered—and rejected as highly unlikely uplift mechanisms—glacio-isostatic rebound, tectonic horizontal compressive stress, and aseismic creep on an unmapped reverse fault. 

Rock dilatancy and coseismic deformation associated with 

the 1959 Hebgen Lake earthquake or the 1975 Yellowstone 

Park earthquake were regarded as speculative or impossible to 

confirm with available data.

The wealth of additional information concerning Yellowstone deformation that has been acquired by leveling, GPS, 
Discussion 29

19InSAR measures range changes in the satellite-to-target direction and cannot separate the effects of vertical and horizontal surface displacements. The 

interferograms discussed here are about three times more sensitive to vertical 

displacements than to horizontal displacements, owing to the relatively steep 

incidence angles of the ERS-1, ERS-2, and ENVISAT radars (23 degrees at 

swath center for ERS-1 and ERS-2; 19–27 degrees for ENVISAT IS2 mode).

20 An exception to this pattern occurred from August 1995 to September 

1996. An InSAR image for that period showed uplift in the northeast part of 

the caldera and subsidence in the southwest part (Wicks and others, 1998).

and InSAR observations since the 1975–77 leveling survey 

generally supports the views expressed by Pelton and Smith 

(1979, 1982). However, we now know that the deformation 

field is more complicated and changeable than could be known 

from the 1923 and 1975–77 leveling surveys alone. These 

additional complexities call for some mechanism other than a 

single intrusive event to explain the situation fully. The next 

section outlines the salient facts about deformation at Yellowstone that must be accommodated in any successful model.

Observational Constraints on Deformation Models

The 1923–2007 leveling results and 1992–2008 InSAR 

observations impose several constraints on the mechanism(s) 

of uplift and subsidence at Yellowstone. The net elevation 

change at benchmark B11 1923 near Le Hardys Rapids from 

1923 to 2007 was 828±23 mm, which corresponds to an 

average uplift rate of 10±1 mm/yr. This is the best estimate 

available for the long-term (century-scale) deformation rate 

and therefore is the appropriate value to use for models that 

seek to explain the caldera’s background deformation mode. 

From the 1983 and 1984 leveling surveys, we know that uplift 

of the Sour Creek dome continued in the same basic pattern (that is, centered along the long axis of the caldera and 

extending approximately from rim to rim along the short axis) 

and at a somewhat greater average rate during 1976–84 (22±1 

mm/yr) than during 1923–76 (14±1 mm/yr). We also know 

that uplift continued during 1983–84 at roughly the same rate 

(16±5 mm/yr), essentially stopped during 1984–85 (−2±5 mm/

yr), and reversed to subsidence during 1985–86 (25±5 mm/yr). 

Therefore, if the 1923–84 uplift occurred as a single episode 

with a simple time history (1) the uplift probably started late in 

the 1923–76 interval, and (2) the uplift rate must have decayed 

rapidly from a value much higher than 22 mm/yr before 1976 

to essentially zero in 1985. For example, if the uplift at Le 

Hardys Rapids began at the time of the 1959 Hebgen Lake 

earthquake, the average uplift rates at DA3 1934 were 43 mm/

yr during 1959–76, 23 mm/yr during 1976–83, 16 mm/yr during 1983–84, and 0 mm/yr during 1984–85. The hypothetical 

uplift rate for 1959–76 is based on the assumption that all of 

the uplift at DA3 1934 that was measured by leveling surveys 

in 1923 and 1975–77 actually occurred during 1959–76 (726 

mm/17 yr=43 mm/yr). Such a rate is plausible, given that the 

uplift rate from mid-2004 to 2008 was that high or higher. 

However, (1) no tectonic mechanism has been proposed to 

account for the shape of the 1923–84 uplift, (2) there is no 

evidence to support high uplift rates in the few years following the Hebgen Lake earthquake (Hamilton, 1987), and 

(3) the deformation pattern since 1985 requires a different 

explanation. Alternatively, if uplift began sometime before 

1923, the higher average uplift rate during 1976–84, relative to 

1923–76, requires a separate explanation. Neither scenario can 

be supported with available information, and neither can easily 

account for the abrupt onset of subsidence during 1985–86.

We know from the 1985–95 leveling surveys that the 

Sour Creek dome subsided progressively and without any 

pause lasting a full year during that entire interval. The 1-year 

subsidence rate at DA3 1934 varied between 35±5 mm/yr during 1986–87 and 9±5 mm/yr during 1987–88, but the average 

subsidence rate during 1985–93 (19±1 mm/yr) was nearly the 

same as the average uplift rate during 1976–84 (22±1 mm/yr). 

What’s more, the 1976–84 uplift profile and 1985–93 subsidence profile are very similar in shape, which suggests a common 

source. The similarity in profile shapes is not definitive in this 

regard, because different sources can produce the same vertical or horizontal surface-deformation pattern. Distinguishing 

between such sources requires information about both vertical 

and horizontal surface displacements (Dieterich and Decker, 

1975). Information about horizontal surface displacements at 

Yellowstone is available from two sources—trilateration and 

GPS surveys.19 The Yellowstone trilateration network was 

established in 1984 and remeasured with a geodolite, a 1-mmresolution electro-optical distance-measuring instrument, in 

1985 and 1987 (Dzurisin and others, 1990). A regional network 

of GPS stations was established in 1987 and remeasured several 

times thereafter, most recently in 2010 (Smith and others, 1989; 

Puskas and others, 2007; Smith and others, 2009; R.B. Smith, 

University of Utah, written commun., November 2010). Neither 

dataset predates the change from uplift to subsidence during 

1984–6, so it is not possible to distinguish between uplift and 

subsidence sources, if any distinction exists, on that basis.

The Occam’s razor principle leads us to speculate that the 

similarity between uplift and subsidence profiles for 1976–84 

and 1985–93, respectively, reflects a simple reversal in some 

process that can cause both uplift and subsidence at different 

times, rather than separate explanations for two opposite but 

otherwise similar outcomes. Furthermore, lacking horizontal 

surface-displacement data to indicate otherwise, we assume that 

uplift and subsidence are sourced in the same place beneath the 

caldera. From the 1976, 1986, 1987, and 1996 leveling surveys 

in the southwest part of the caldera, we know that the Mallard 

Lake dome moves mostly in the same direction as the Sour 

Creek dome during any given interval and at approximately the 

same rate.20 The same inference was drawn from the 1923 and 

1975–77 surveys, which showed similar amounts of net uplift 

at both domes. However, the greater than 50-year interval 

between those surveys left open the possibility of differential 

motion over shorter time scales. We conclude that whatever 

the deformation mechanism is, its reach includes both domes, 
30 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

and its effects are sometimes felt at both domes over time

scales of 2 years or less. For the case of rapid caldera uplift 

starting in mid-2004, the onset was nearly simultaneous across 

the entire caldera (Chang and others, 2007, 2010). If lateral 

transport of magma or hydrothermal fluid is involved, this 

implies a high degree of connectivity within the magmatic or 

hydrothermal system beneath the caldera.

The discovery of subsidence during 1985–93 that mirrored 

the uplift during 1976–84 seemed to make a purely magmatic 

model for Yellowstone deformation less tenable. If surface uplift is 

caused by magma intrusion, then mirror-image subsidence could 

reasonably be attributed to magma withdrawal—but to where? 

Unlike many basaltic magma systems (for example, Kīlauea and 

Mauna Loa, Hawai‘i; Krafla, Iceland; Mount Etna, Italy; Piton de 

la Fournaise, and Réunion Island, Indian Ocean), the Yellowstone 

caldera does not have radiating rift zones to accommodate intrusions of magma from a deflating central reservoir. Conceivably, 

basaltic magma rises to the base of the silicic magma system, 

where it cools, densifies, and then sinks—a process that could 

account for both uplift and subsidence. This mechanism is akin to 

that proposed by Shaw and others (1971, p. 880–882) on a much 

larger scale to account for thermal circulation in the upper mantle. 

However, the time scale required for densification by cooling, 

crystallization, or volatile loss from basaltic magma beneath Yellowstone likely is too long to account for sudden reversals from 

uplift to subsidence, such as the one that occurred between 1984 

and 1986. We revisit the magma-intrusion scenario in the Deformation Mechanisms and Models section of this paper, where we 

propose a variation of the concept that satisfies all of the observational constraints on deformation mechanisms.

An alternative explanation of mirror-image uplift and 

subsidence is that surface uplift-subsidence cycles reflect 

pressurization-depressurization episodes in Yellowstone’s 

deep hydrothermal system (Dzurisin and others, 1990). We 

know from research drilling at Yellowstone and elsewhere, 

laboratory experiments, and theoretical considerations that the 

hydrothermal system consists of a deep, ductile zone in which 

pressure is near lithostatic and a shallow zone in which pressure is hydrostatic (Fournier and Pitt, 1985; Fournier, 1991, 

1999). The two zones are separated by an impermeable, selfsealing layer created by mineral deposition and quasi-plastic 

flow, which at Yellowstone is thought to occur at depths of 

about 5 km (Fournier, 1999). The deep hydrothermal system 

beneath the self-sealing layer is near lithostatic pressure, so 

any change in pressure there has the potential to cause uplift or 

subsidence. Magmatic brine that forms during phase separation following cooling and crystallization of rhyolitic magma 

beneath the caldera would become trapped in horizontal lenses 

beneath the self-sealing layer, pressurizing it and potentially 

causing uplift (Fournier, 2007). Rupture of the self-sealing 

layer, likely during an earthquake swarm, would trigger fluid 

loss by upward transport, depressurization, and subsidence. If, 

as suggested by Fournier and Pitt (1985), a state of hydraulic 

equilibrium prevails throughout the deep hydrothermal system, 

subsidence would extend far beyond the epicentral area. Thus, 

rupturing of the self-sealing layer anywhere might cause the 

entire caldera floor to subside, as was the case during 1986–95.

Waite and Smith (2002) proposed such a mechanism to 

explain the October 1985 to early 1986 earthquake swarm near 

Madison Junction, which occurred between leveling surveys in 

September 1985 and September 1986 that bracketed the onset 

of subsidence inside the caldera. The swarm was noteworthy for 

several reasons: (1) it was the most energetic earthquake swarm 

ever recorded at Yellowstone; (2) seismic activity migrated at 

an average rate of 113 m/day to the northwest, away from the 

caldera rim, during the first month of the swarm, and thereafter 

it gradually migrated downward from depths of 2–5 km to more 

than 10 km; (3) the dominant focal mechanisms for the early 

part of the swarm were oblique-normal strike-slip instead of 

more typical normal-faulting mechanisms; (4) the maximum 

principal stress rotated from vertical to horizontal and subparallel to the axis of the swarm; and (5) the swarm was accompanied by two small steam explosions, increasing ground temperatures, and formation of new fumaroles and a mud volcano at 

three widely separated sites near the caldera rim (Dzurisin and 

others, 1994). Waite and Smith (2002, p. 2190 or ESE 1–14) 

concluded: “The most likely scenario [to account for the swarm 

and caldera subsidence] involves the rupture of a self-sealed 

hydrothermal layer and subsequent migration of hydrothermal 

fluid through a preexisting fracture zone out of the caldera.” 

Before exploring this idea further, we move on to additional 

observational constraints on the role that hydrothermal fluid 

might play in Yellowstone deformation.

Starting in 1992, the most revealing information about 

ground deformation at Yellowstone came, not from leveling 

surveys, but from sequential InSAR observations. We learned 

from interferograms that spanned 1992–93 and 1993–95 that, 

although both domes subsided during both of those intervals, 

they did so differentially. That is, the area of maximum subsidence switched from the Sour Creek dome during 1992–93 

to the Mallard Lake dome during 1993–95. This occurrence 

suggested either a lateral connection beneath the domes or 

separate deformation sources that might be connected at 

greater depth. An interferogram for 1995–96 captured what 

seems to be a relatively rare occurrence; that is, the Sour 

Creek dome rose while the Mallard Lake dome subsided. By 

1997, uplift had spread southwestward across the caldera to 

include the Mallard Lake dome. These observations led Wicks 

and others (1998, p. 461) to conclude (1) that “interacting fluid 

reservoirs exist beneath the SC [Sour Creek] and ML [Mallard 

Lake] domes, with the inflation of each being regulated by 

flow through two conduits, one beneath SC * * * and the other 

connecting SC and ML* * *”, and (2) that “* * * the driving 

pressure source may lie beneath the SC dome.” The key bit of 

evidence for a lateral connection between the domes was the 

migration of the deformation front from one dome to the other 

within 2 years or less.

The Yellowstone deformation puzzle became even more 

complicated when InSAR images for 1996–2000 revealed a 

third deformation source beneath the north caldera rim (Wicks 
Discussion 31

and others, 2006). Recall that an earlier InSAR study and 

leveling surveys (Wicks and others, 1998; Dzurisin and others, 

1999) indicated that following the period of caldera subsidence from 1985 to 1995 uplift began at the Sour Creek dome 

during 1995–96 and spread southwestward to the Mallard 

Lake dome during 1996–97. North-rim uplift began in 1997 

and continued into 2002, while the caldera floor subsided. In 

this case, nearly simultaneous onset of north-rim uplift and 

caldera-wide subsidence suggested a connection among three 

sources—two beneath the resurgent domes and one beneath 

the north rim. The latter was centered near the structural intersection of the caldera ring fault, north-trending Norris–Mammoth corridor, and a west-northwest striking seismic belt east 

of the Hebgen Lake fault zone.

The most recent chapter in the Yellowstone deformation 

story is the near-simultaneous onset of uplift across the entire 

caldera in mid-2004, accompanied by north-rim subsidence. 

Chang and others (2007) jointly inverted GPS and InSAR 

observations, and wrote: “We thus suggest that the 2004–2006 

episode of accelerated inflation occurred in response to a 

caldera-wide magma recharge of the Yellowstone volcanic 

system.” Acknowledging that models of earlier deformation 

episodes are consistent with fluid pressurization in the upper 

part of the magmatic system, they concluded, “* * * magma 

intrusion and fluid pressurization should be considered as 

jointly operating processes to explain the accelerated caldera 

uplift reported here, although our estimate of large volume 

increase implies the former as a preferred source model.” 

With regard to north-rim subsidence, Chang and others (2007) 

wrote: “We therefore propose that the 2004–2006 deflation 

near Norris Geyser Basin was in response to a redistribution of 

hydrothermal fluids as a consequence of caldera inflation.”

To the extent that the model of Chang and others (2007, 

2010) reflects current consensus, we can say that progress 

since the work of Pelton and Smith (1979, 1982) has been 

mainly in two areas: (1) documentation of the complexity—in 

both space and time—of the deformation field and (2) recognition of the likely role played by fluids in causing uplift and, 

especially, subsidence. Two remaining challenges are (1) to 

identify the fluids involved and (2) to improve understanding 

of the physical interactions among magma, fluids, heat, the 

hydrothermal system, and country rock that produce seismicity and surface deformation.

These challenges are being addressed in several ways. 

In addition to improved monitoring of deformation, gravity, 

seismic, and geochemical parameters, more realistic models 

are being developed to investigate the likely role of fluids and 

heat in causing ground deformation. For example, Hurwitz 

and others (2007a) carried out numerical simulations of 

hydrothermal-fluid flow and deformation in large calderas by 

coupling two numerical computer codes—TOUGH2 (Pruess 

and others, 1999) and BIOT2 (Hsieh, 1996). TOUGH2 is a 

three-dimensional (3-D) integrated finite difference simulator 

for nonisothermal, multicomponent (for example, H2

O-CO2-

NaCl), and multiphase flow in porous or fractured media. 

BIOT2 simulates axisymmetric or plane-strain deformation 

and fluid flow in a linearly elastic porous medium. To simulate the effects of fluids exsolving from a crystallizing magma 

body beneath a Yellowstone-size caldera, Hurwitz and others 

(2007a) used TOUGH2 to calculate the pressure and temperature distributions resulting from injection and flow of hightemperature (350oC) water within a cylindrical volume with a 

radius of 50 km and a height (thickness) of 3–5 km. They used 

the output from TOUGH2 as input to BIOT2 to calculate the 

resulting elastic deformation field as a function of time. Their 

simulations showed that small differences in the assumed values of permeability and its anisotropy, shear modulus, and the 

depth and rate of fluid injection lead to large variations in the 

magnitude, rate, and geometry of surface deformation. Nonetheless, they were able to simulate the uplift rates observed 

at the Yellowstone, Long Valley, and Campi Flegrei calderas 

by using reasonable suites of parameter values. Todesco and 

others (2004) used a similar approach and showed that the 

temporal pattern of deformation at Campi Flegrei (that is, 

rapid uplift followed by slower subsidence) could be produced 

by poroelastic response of the shallow hydrothermal system 

to a pulse of magmatic fluid from depth. Hurwitz and others 

(2007a, p. 1) concluded that “* * * the injection of aqueous 

fluids into the shallow crust may explain some of the deformation observed in calderas.”

Using a similar approach, Hutnak and others (2009) 

addressed an acknowledged limitation of the study by Hurwitz and others (2007a), that is, the assumption of single 

phase (liquid) and single component (pure water) fluid flow. 

In their simulations, Hutnak and others (2009) included two 

components (hot water and carbon dioxide, which are sourced 

into the overlying hydrothermal system from a crystallizing 

magma body) and two water phases (liquid and steam). For a 

range of plausible hydrologic parameters (permeability, injection rates, depths, and fluid compositions), they investigated 

effects of (1) high versus low permeability in rock hosting the 

hydrothermal system, (2) fluid injection from a point source 

versus a distributed source, (3) fluid saturation in water versus 

water + carbon dioxide systems, and (4) pulsed injection to 

simulate repeated rupturing of a self-sealing layer. Hutnak and 

others (2009) concluded:

(1) Numerical models are capable of generating ground 

surface deformation rates, magnitudes, and geometries similar 

to those observed in some large calderas,

(2) Simulated deformation (including subsidence) can 

be complex in time and space, even for homogenous property 

distributions and steady fluid sourcing,

(3) Gas formation in the shallow subsurface leads to 

broader radial extents and magnitudes of uplift, as well as 

increased uplift rates, and

(4) Hydrothermal circulation may explain periods of 

rapid surface deformation not followed by volcanic eruption.

Hutnak and others (2009, p. 10) summarized their findings as follows: “The potential for both rapid and gradual GSD 
32 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

[ground-surface displacement] resulting from magma derived 

aqueous fluids, as demonstrated in this study, suggests that 

hydrothermal fluid circulation may help explain some occurrences of GSD that have not culminated in magmatic eruptions.” In the future, this study could be extended to include 

other volatile components (sulfur, chlorine), hypersaline brine 

(a byproduct of crystallization of silicic magmas), and heterogeneous permeability (layers, fractures).

Complications that include heterogeneous and poorly 

known crustal properties are a problem for any type of deformation model, whether its approach is analytical or numerical 

and whether the assumed crustal rheology is elastic, plastic, or 

other. Some of these effects have been explored in recent publications by Masterlark (2007), Manconi and others (2007), 

Hurwitz and others (2007a), and Hutnak and others (2009). 

Such complexities might preclude a definitive result from 

modeling for the foreseeable future, but we believe nonetheless that modeling holds considerable promise for providing 

insights into caldera deformation processes and for guiding 

future observational and experimental studies.

More than three decades after the discovery of surface 

uplift at Yellowstone (Pelton and Smith, 1979; 1982), the 

understanding of its causes is still a work in progress. Consideration of poroelastic and thermoelastic effects, and of nonelastic rheology—which were ignored in most previous studies 

of Yellowstone but now are amenable to investigation—calls 

into question any conclusion about the causes of deformation 

(including our preferred model, see below). Future work, both 

experimental and theoretical, will reveal to what extent our current thinking reflects the real situation at Yellowstone.

Constraints from Deformed Shorelines and 

Microgravity Measurements

Two sets of observations that bear directly on the nature 

and causes of ground deformation at Yellowstone have yet to 

be discussed: (1) studies of deformed shorelines that extend 

the record of vertical surface displacements backward in time 

for several thousand years (Pierce and others, 2002), which 

provides important context for conceptual models; and (2) 

repeated microgravity observations, which when combined 

with height-change information from leveling and GPS, provide constraints on mass changes beneath the caldera and have 

the potential to distinguish between relatively low-density 

hydrothermal or magmatic fluids and higher-density magma as 

the primary deformation agent.

Evidence from Deformed Shorelines

 The shorelines of Yellowstone Lake and its outlet stream, 

the Yellowstone River, record a rich Holocene deformation 

history, including at least two major inflation-deflation cycles 

that were thousands of years long and ~30 m in amplitude at 

Le Hardys Rapids relative to the caldera margin (Pierce and 

others, 2002, p. 2, 13; Locke and Meyer, 1994; Meyer and 

Locke, 1986). Between the lake outlet at Fishing Bridge and 

the bedrock knickpoint at Le Hardys Rapids, 5 km downstream, the river’s surface drops only 0.25 m. As a consequence of this extremely low gradient, vertical motion of the 

knickpoint affects mean lake level. Uplift of Le Hardys Rapids 

causes mean lake level to rise, and subsidence causes the mean 

level to fall. These changes are superimposed on a long-term 

trend of falling lake level, as a result of erosion at the outlet (Pierce and others, 2002, p. 16). During relatively stable 

stands of the lake, terraces form in horizontal position along 

the shoreline. Subsequent vertical motion at Le Hardys Rapids 

affects preexisting terraces in two ways: (1) uplift or subsidence of the knickpoint cause terraces to submerge or emerge, 

respectively, as the lake level rises or falls, and (2) differential uplift or subsidence warps terrace surfaces that were 

initially planar and horizontal. Thus, tilting of the lake basin 

has resulted in terraces that range from about 30 m below the 

present shoreline to about 20 m above it. Each terrace records 

the cumulative effects of all uplift or subsidence episodes 

since the terrace formed. By measuring the tilts of a sequence 

of remnant terrace surfaces that surround the modern lake and 

determining the ages of terraces, it is possible to unravel the 

postglacial deformation history of the lake basin.21

Pierce and others (2002) deduced from their study of 

deformed terraces at Yellowstone Lake that twice during the 

past ~15,000 years Le Hardys Rapids has been uplifted ~8 m 

relative to the current lake outlet—most recently from 3,000–

4,000 years ago to the present. Because those locations span 

only ~25 percent of the uplift measured by leveling surveys, 

they projected ~32 m of total uplift relative to the caldera 

margin. Nonetheless, subsidence has balanced or slightly 

exceeded uplift over the entire postglacial period, as shown by 

older shorelines that descend toward the caldera axis. Pierce and 

others (2002) inferred average uplift rates at Le Hardys Rapids 

over millennial time scales in the range 12–16 mm/yr, which is 

close to the average historical uplift rate from leveling surveys.

Although the average uplift rates during parts of postglacial 

time and historical time are similar, the sheer magnitude of 

postglacial motions adds another dimension to considerations 

of deformation mechanisms. Pierce and others (2002) considered magma intrusion, crustal stretching, batholithic cooling, 

and hydrothermal pressure changes to be likely contributors. 

They concluded (Pierce and others, 2002, p. 2): 

We favor a hydrothermal mechanism for inflation 

and deflation because it provides for both inflation and deflation with little overall change. Other 

mechanisms such as inflation by magma intrusion 

and deflation by extensional stretching require two 

separate processes to alternate and yet result in no 

net elevation change. 

21Terrace ages were established by radiocarbon dating of incorporated 

organic material, projectile-point ages obtained in archeological surveys, and 

geological studies (Pierce and others, 2002).
Discussion 33

They clarified their use of the term “hydrothermal 

mechanism” as follows (Pierce and others, 2002, p. 24): “The 

mechanism of uplift by hydrothermal pressure buildup beneath 

a hydrothermal seal followed by subsidence due to rupture 

of the seal, and release of fluids * * *.” It is important to note 

that the fluids to which they refer are most likely of magmatic 

origin, that is, released during intrusion or subsequent crystallization of basaltic or rhyolitic magma beneath the caldera 

(Fournier, 1989, 1999). Fluids of meteoric origin are limited to 

the hydrostatically pressured part of the hydrothermal system, 

above the self-sealing layer, where fracture permeability is 

high. Magmatic fluids that are trapped beneath, and eventually 

escape through, the self-sealing layer that caps the lithostatically pressured part of the hydrothermal system are the more 

likely deformation agent.

Evidence from Repeated Microgravity Surveys

Fluid connections among three separate but interacting 

deformation sources beneath the Yellowstone caldera are a 

plausible explanation for (1) migration of subsidence from the 

Sour Creek dome to the Mallard Lake dome during 1993–95, 

(2) migration of uplift from the Sour Creek dome to the Mallard Lake dome during 1995–97, (3) the near-simultaneous 

onset of north-rim uplift and caldera-wide subsidence in 1997, 

and (4) the reversal from north-rim uplift and caldera-wide 

subsidence during 1997–2002 to north-rim subsidence and 

caldera-wide uplift in mid-2004. A key question—What is the 

nature of the fluid?—is best addressed by repeat microgravity 

observations, which are sensitive to subsurface mass changes 

and to the density of the material added or lost. In the absence 

of any subsurface mass change, the ratio of gravity change 

to height change, ∆g/∆h, is the average free-air gradient of 

−3.086 microgal per centimeter (μGal/cm). If height changes 

are accompanied by a change in mass (density) within a subsurface body of arbitrary shape, the value of ∆g/∆h varies with 

the body’s shape and its density contrast with the surrounding 

rock (Telford and others, 1990; Turcotte and Schubert, 2002).

The Yellowstone gravity network consists of 160 stations along the level lines shown in figure 6, plus 40 additional 

stations in remote areas where height changes are measured 

by GPS observations (Hollis and others, 1987; Smith and 

others, 1989; Arnet and others, 1997). Each station comprises 

a benchmark and some means to precisely position a gravity 

meter either directly over the mark or adjacent to it (for example, three small, shallow holes (“dimples”) in bedrock that 

align with the feet of a gravity-meter base plate). The gravity 

network was observed in 1977, 1983, 1987, 1991, and 1993, 

each time during late summer or early autumn to reduce the 

effect of water-table changes. Arnet and others (1997) reported 

that the gravity field decreased across the caldera from 1977 to 

1983 by as much as 60±12 μGal along the caldera axis, then 

increased by the same amount from 1986 to 1993. The sense 

of the changes is consistent with leveling observations that 

indicated uplift from 1975–77 to 1984 and subsidence from 

1985 to 1993. For stations along the level line between Lake 

Butte and Mount Washburn, Arnet and others (1997) found 

the weighted mean value of ∆g/∆h=−1.7±0.7 μGal/cm for the 

1977–83 period of uplift and ∆g/∆h=−3.3±1.0 μGal/cm for 

the 1986–93 period of subsidence. Arnet and others (1997, p. 

2743) wrote:

Based on these results, we conclude that a mass 

increase occurred during the uplift episode. During 

the subsidence the gravity change followed almost 

the free-air gradient, thereby indicating that probably no mass change took place during this period.

To account for contemporaneous surface uplift and mass 

increase, Arnet and others (1997, p. 2744) called upon redistribution of hydrothermal fluid and magma input:

We suggest that the most likely source of the gravity decrease, concomitant with the crustal uplift, is 

related to widespread hydrothermal fluid movement, 

which furthermore is related to input by magma. 

Basaltic intrusions into the mid or upper crust, 

pressurization of a deep hydrothermal system by 

magmatic gas, or brine released by crystallization 

of a rhyolitic melt, are also plausible sources for the 

uplift. It cannot have been caused solely by pressurization of the deep hydrothermal system, without 

any significant mass increase.

Regarding the mechanism for subsidence, which occurred 

without a detectable change in subsurface mass, Arnet and others (1997, p. 2744) wrote:

For the episode of gravity increase – ground subsidence, the spatial synchroneity between the area 

of uplift (1923–1984) and the area of subsidence 

(1984–1994) suggests a common causative mechanism related to magmatic/hydrothermal intrusion 

followed by degassing from the same source. The 

fact that subsidence has not been accompanied by 

a measurable mass change is best explained by 

depressurization of the deep hydrothermal system as 

a result of fracturing and volatile loss to the shallow 

hydrothermal system or, less likely, reduced input 

of brine to the deep system. The first interpretation is preferred because it more easily explains the 

relatively abrupt change from uplift to subsidence 

during 1985–1986 (Dzurisin et al., 1990), which 

was also accompanied by sizable earthquake swarms 

originating in the upper crust.

Three factors weigh on the side of caution in interpreting 

the gravity results. First, the ∆g/∆h values for periods of uplift 

and subsidence are essentially the same within two standard 

deviations. Second, the ∆g/∆h value for a single station north 

of Fishing Bridge, where the gravity and height changes were 

greatest, was −3.3±0.7 μGal/cm for the entire period from 

1977 to 1993, which included roughly equal amounts of uplift 

and subsidence (Arnet and others, 1997, p. 2743). Within 

uncertainty, this value is the same as the average free-air 
34 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

gradient (−3.086 μGal/cm). Third, owing to a lack of suitable 

water-well data, no corrections were made to the gravity data 

for water-table differences between surveys. Although all of 

the surveys were conducted during the same time of year to 

reduce water-table effects, some unknown contribution to the 

gravity measurements is still possible. Uncorrected water-table 

variations would introduce noise of unknown magnitude to the 

gravity measurements and possibly bias ∆g/∆h values in an 

unpredictable way.

These reservations notwithstanding, the microgravity 

observations reported by Arnet and others (1997) provide support for the idea that surface uplift is caused, at least in part, 

by intrusion of magma—presumably basalt—in the mid-crust. 

Likewise, the evidence from microgravity observations that 

subsidence can occur with little or no subsurface mass change 

is consistent with the idea that fluid loss from magma—either 

as brine or gas and either from crystallizing rhyolite or degassing basalt—is a viable deformation mechanism. Additional 

microgravity observations, during periods of both uplift and 

subsidence and in areas affected by all three deformation 

sources, should provide additional constraints on deformation mechanisms. Scientists from the University of Utah now 

conduct microgravity surveys throughout Yellowstone National 

Park each autumn, most recently in 2010 (R.B. Smith, University of Utah, written commun., November 2010).

Deformation Mechanisms and Models

Previous Work

The discovery of caldera-floor subsidence starting 

in 1986, following a period of uplift, triggered a search 

for a deformation mechanism that could account for both 

of those deformation modes and also the relatively rapid 

switch between them. Subsequent observations—namely 

(1) deformation-front migration from the Sour Creek dome 

to the Mallard Lake dome, (2) alternating inflation-deflation 

cycles between the north-rim source and both caldera-floor 

sources, and (3) large but nearly offsetting amounts of uplift 

and subsidence within the caldera during postglacial time—led 

to general agreement that fluids of some sort are involved in 

the deformation mechanism (Dzurisin and others, 1990; Wicks 

and others, 1998, 2006; Pierce and others, 2002; Waite and 

Smith, 2002; Chang and others, 2007, 2010).

Two general classes of models have been proposed to 

explain surface deformation at Yellowstone since its discovery more than three decades ago, and both remain viable 

in various forms today. One class of models, commonly 

referred to as “hydrothermal” models, relies mainly on pressurization/depressurization of fluids other than magma to 

produce surface uplift/subsidence. Three variations on this 

theme are (1) supercritical fluid forms by phase separation 

from crystallizing magma and is trapped beneath a selfsealing layer at lithostatic pressure in ductile rock to produce 

uplift; episodic rupturing of the seal produces subsidence 

(for example, Dzurisin and others, 1990); (2) poroelastic 

deformation in the shallow hydrothermal system is induced 

by magmatic gas influx at the base of the hydrothermal 

system (for example, Todesco and others, 2004; Hurwitz and 

others, 2007a; Hutnak and others, 2009); and (3) precipitation recharge induces a poroelastic response of the shallow 

hydrothermal system (for example, Howle and others, 2003; 

Jahr and others, 2008; Westerhaus and others, 2005). In the 

first case, one might expect an increase in the surface flux of 

magmatic fluid associated with the onset of subsidence. This 

has not been observed, but there are reasons to think that the 

effect could be masked by the shallow hydrothermal system 

(Fournier, 2004) or so long-delayed as to be unrecognizable 

(Hurwitz and others, 2007b). The second type of hydrothermal model, involving poroelastic deformation in response to 

magmatic gas influx at the base of the hydrothermal system, 

requires a high rate of basaltic intrusions or convection of a 

large volume of basalt in the lower to mid-crust (Lowenstern 

and Hurwitz, 2008). Both processes would produce episodic 

releases of magmatic gases including CO2

—which could 

account for the very high surface flux of CO2 observed at 

Yellowstone (Werner and Brantley, 2003). A third type of 

hydrothermal model, in which precipitation recharge induces 

a poroelastic response of the shallow hydrothermal system, 

has been invoked to explain surface motions at the Casa 

Diablo geothermal area in California (Howle and others, 

2003) and at Merapi volcano, Indonesia, but as yet has not 

been applied to Yellowstone. In our opinion, all three of 

these “hydrothermal” mechanisms probably contribute in 

some measure and over various time scales to deformation at 

Yellowstone, and they merit further investigation through the 

use of numerical simulations or other modeling techniques.

The second class of models proposed to explain surface 

deformation at Yellowstone is generally referred to as “magmatic” for brevity. We will be explicit about our usage of the 

terms magma, magmatic fluid, and hydrothermal fluid later in 

this section, but for now the following distinction is sufficient. 

The primary deformation mechanism in magmatic models is 

injection of magma or fluid exsolved from magma at or below 

the base of the hydrothermal system, whereas in hydrothermal models the primary deformation mechanism is a pressure 

change within the hydrothermal system in response to some 

perturbation such as a gas pulse or precipitation recharge. Our 

preferred model is an amalgamation of ideas from recent studies by several authors, as summarized below.

Chang and others (2007) used joint inversion of InSAR 

and GPS data for 2004–6 to identify two deformation sources 

at Yellowstone: (1) an expanding, subhorizontal, sill-like 

source (that is, a tabular body) at a depth of 7–12 km beneath 

the caldera, and (2) a contracting volume at a depth of 8–16 

km under the Norris area. These results are consistent with the 

earlier model of Wicks and others (2006) for 1996–2002, a 

period of north-rim uplift and caldera-floor subsidence. Their 

model included two subcaldera, contracting tabular bodies, 

one beneath the Sour Creek dome and the other beneath the 

central part of the caldera and Mallard Lake dome, both at 
Discussion 35

depths of 6–14 km. These correspond to the single, expanding tabular source at 7–12 km depth in the model of Chang 

and others (2007). For the north-rim source, Wicks and others 

(2006) proposed an expanding tabular source at 9–16 km 

depth, which corresponds to the contracting volume at 8–16 

km depth in the model of Chang and others (2007). Both sets 

of authors called on fluid interactions between/among sources 

and regarded magma as the “fluid” most likely to be involved. 

However, both groups also mentioned the alternative possibility that hydrothermal or magma-derived fluids were the 

main deformation agent. In supplementary material for their 

manuscript, Wicks and others (2006) noted that a distributed, 

shallower source could be constructed to fit the north-rim 

uplift data for 1996–2002 equally well. A shallower source 

would be more permissive of the idea that deformation originates in the deep part of the hydrothermal system rather than 

in the underlying magmatic system. We note that the source 

depths indicated by these models are below the brittle-ductile 

transition, which is thought to occur at ~5 km depth beneath 

the Yellowstone caldera (Smith and Bruhn, 1984; Smith and 

Braile, 1994). Therefore, the models’ assumption of linear 

elastic behavior is likely to be violated in rock surrounding the 

deformation sources, which adds uncertainty to any inferences 

drawn from the results.

Parallel studies by Puskas and others (2007) and Vasco 

and others (2007) produced results that are consistent with 

those of Wicks and others (2006) and Chang and others (2007, 

2010) and offer additional insight into the configuration of 

deformation sources beneath the caldera. Puskas and others 

(2007) presented results of campaign-mode GPS and CGPS 

observations of the Yellowstone–Snake River Plain region for 

1987–2004. For the Yellowstone caldera, including the northwest caldera boundary (equivalent to the north-rim source 

area discussed here), they divided the observations into three 

time windows of uniform ground motion. During 1987–95, 

the GPS observations indicate that the center of the caldera 

subsided 14±3 mm/yr. Leveling in the vicinity of Le Hardys 

Rapids gives essentially the same result for the same time 

period, that is, subsidence at an average rate of 16±1 mm/yr. 

For 1995–2000, Puskas and others (2007) reported from their 

analysis of GPS data that the north-rim area rose 15±4 mm/

yr. That result is consistent in a qualitative sense with InSAR 

observations reported by Wicks and others (2006), which 

indicate about 100 mm of range decrease (mostly uplift) along 

the north-rim during 1996–2000. For 2000–3, the GPS and 

InSAR observations both indicate caldera-floor subsidence 

and north-rim uplift. Puskas and others (2007) found that the 

caldera floor subsided at an average rate of 9±6 mm/yr during 

that period, while the north-rim area rose at an average rate of 

12±4 mm/yr. Wicks and others (2006) reported about 50 mm 

of range increase (mostly subsidence) in the central part of the 

caldera and about 40 mm of range decrease (mostly uplift) in 

the north-rim area during 2000–2002, and they also reported 

that both motions paused during 2002–3. Thus, the average 

displacement rates for 2000–3 from InSAR are about 17 mm/

yr subsidence in the caldera and about 13 mm/yr uplift along 

the north rim. The GPS and InSAR results for this period are 

consistent qualitatively, although the measured caldera-floor 

subsidence rates differ by about a factor of 2.

Vasco and others (2007) used InSAR, GPS, and leveling 

observations to constrain deformation source models for the 

Yellowstone caldera for 1992–95, 1996–2000, 2000–2001, 

and 2001–2. Their modeling approach differed from those 

of Wicks and others (2006) and Chang and others (2007) 

in that Vasco and others (2007) allowed for an arbitrary, 

three-dimensional distribution of subsurface volume change 

rather than prescribing a specific source shape, such as a 

sphere, ellipsoid, or sill. To make the resulting mathematical 

problem tractable, they included penalty terms and inequality constraints in their formulation of the problem. Interested 

readers should consult Vasco and others (2002, 2007) for 

details. For 1992–95, Vasco and others (2007) concluded 

that caldera-floor subsidence was associated with decreases 

in source volume at depths of 6–10 km beneath the Elephant 

Back fault zone, that is, beneath a southwest-trending zone 

between the Sour Creek and Mallard Lake resurgent domes 

and extending northeast to Hot Springs Basin. The Elephant 

Back fault zone is approximately parallel to, and offset 

slightly to the southeast of, the long axis of the caldera (fig. 

1). Both features are coincident with the area of maximum 

caldera-floor uplift from 1923 to 1975–77 as contoured by 

Pelton and Smith (1979, 1982). Another intriguing aspect of 

the source model for caldera-floor subsidence during 1992–5 

is a linear trend of volume decrease extending north from the 

center of the caldera toward the north rim in the Norris area 

(fig. 17, reproduced from Vasco and others, 2007, fig. 5, p. 

9). Vasco and others (2007) noted that the linear subsidence 

feature is located beneath an alignment of volcanic vents that 

fed post-caldera collapse (that is, younger than 0.64 Ma) lava 

flows, as mapped by Christiansen (2001).

We are struck by the similarity between the pattern of 

volume decreases inferred by Vasco and others (2007) and 

the “magma migration paths” interpreted by Wicks and others (2006) (compare figs. 17 and 18). Granted that both studies 

analyzed the same InSAR data, the similarity is nonetheless 

remarkable when one considers that Vasco and others (2007) 

also included GPS and leveling data in their analysis and used a 

different modeling approach. Wicks and others (2006) modeled InSAR data only, which do not provide the independent 

measurements of vertical and horizontal surface displacements that would greatly improve the constraints on source 

depth and shape (Dieterich and Decker, 1975). Accordingly, 

they considered only simple source shapes (prolate ellipsoid, 

sill) and showed that slightly dipping sills beneath the caldera 

(depth about 11 km, contracting) and north-rim (depth 9–16 

km, expanding) provide the best fit to the 1996–2002 InSAR 

data. Their hypothesized magma migration paths (fig. 18) do 

not derive directly from modeling but instead derive from 

consideration of the relative timing of uplift and subsidence at 

different locations. Uplift of the Sour Creek resurgent dome 

started during 1995–96 and spread to the Mallard Lake dome 

during 1996–97. North-rim uplift began in 1997 and continued 
36 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

into 2002, while the resurgent domes subsided. Wicks and others (2006) inferred from that pattern of deformation that a pulse 

of basaltic magma rose beneath the Sour Creek dome during 

1995–96 and then migrated laterally to beneath the Mallard 

Lake dome and north rim area during 1996–97. They sketched 

their magma migration pathways accordingly.

Vasco and others (2007), on the other hand, included 

GPS and leveling data in their analysis, which together with 

the InSAR data provide a stronger constraint on the depth and 

shape of the source (although still subject to the limitation, 

noted above, imposed by assuming linear elastic rheology in 

the source region). Their result for the depth of the subcaldera 

source (6–10 km) differs slightly from that of Wicks and others (2006, about 11 km), but both results can be interpreted to 

coincide with the upper part of a crystallizing rhyolite body 

that has been hypothesized to exist beneath the caldera (for 

example, Eaton and others, 1975; Miller and Smith, 1999; 

Christiansen, 2001). Consistent with that idea, Husen and 

others (2004, p. 397) identified a low velocity zone for seismic 

P waves below 8 km depth that they interpreted as “* * * possibly representing hot, crystallizing magma.” More important, 

in our opinion, than the exact source depth is the fact that 

Vasco and others (2007) were able to discern a connection 

among the deformation sources that underlie the Sour Creek 

dome, Mallard Lake dome, and north rim area. Their depiction 

of subsurface volume decreases during 1992–95 is reminiscent 

of the magma plumbing system sketched by Wicks and others 

(2006) on the basis of the chronology of caldera-floor subsidence and north-rim uplift during 1995–2002. Furthermore, 

Vasco and others (2007) claim to have detected an episode of 

lateral fluid intrusion from beneath the center of the caldera to 

the north-rim area during that period, which is consonant with 

the conceptual model of Wicks and others (2006). Vasco and 

others (2007, p. 15–18 of 19) concluded:

Based upon contemporary Yellowstone deformation, we find that subsurface volume changes correlate with the resurgent domes and the Elephant 

Back fault zone, north-trending extensional faults 

related to a line of volcanic vents, and the extensive 

magma body beneath the caldera. These correlations suggest that such features control or at least 

influence deformation within the caldera, either as 

zones of mechanical weakness or as pathways for 

fluid flow or both. * * * We thus hypothesize that the 

observed surface deformation within and adjacent 

to the Yellowstone Caldera is due to the interaction 

of an underlying, large-scale crystallizing magmatic 

system and zones of weakness associated with 

crustal faults. In particular, large-scale pressure and 

mass changes within the magma body are focused 

into faults that act as narrow conduits or pathways 

for flow. The focused flow and pressure changes 

give rise to observable surface deformation. * * * 

We envision that the Yellowstone volcanic system 

involves the interaction of at least three subsystems: the shallow crustal hydrothermal system and 

seismogenic faults, the semisolid magmatic body, 

and a deeper underlying basaltic magma system. 

As yet, we cannot specify the exact nature of the 

interaction between these major components, and we 

can only draw very general conclusions about the 

influence of one component upon another. Specifically, the results here imply that crustal faults act as 

conduits for flow from within the Earth. Thus we 

are not proposing an encompassing model of the 

Yellowstone system, as given recently by Wicks et 

al. [2006].

Notwithstanding the qualification at the end of the preceding quotation, the general agreement between the conceptual model of Wicks and others (2006) and the numerical 

modeling results of Vasco and others (2007) is encouraging 

and, in our view, lends support to both studies.

Our Preferred Model

We are cognizant of the fact that geodetic data such as 

those presented here provide information only about displacements of the ground surface, which are but one manifestation 

of complex subsurface processes. To infer what those processes 

might be from such information, modelers generally must incorporate additional datasets and rely on assumptions or simplifications to make the modeling problem tractable. This introduces 

an additional level of uncertainty, which is best addressed by 

quantitative models that can test specific assumptions or hypotheses. The conceptual model outlined below does not meet that 

standard, nor is it intended to exclude other types of models that 

ultimately might prove to be correct. Instead, it describes our 

current thinking based on more than 25 years of deformation 

studies at Yellowstone and it offers some testable hypotheses 

for others to confirm or reject. Our preferred model is not an 

original contribution but rather an amalgam of ideas from earlier 

studies by several authors.

Taken together, the ideas put forward by Wicks and others (2006), Chang and others (2007, 2010), and Vasco and others (2007) can explain both the spatial and temporal patterns 

of deformation observed at Yellowstone. The leveling and 

InSAR results for 1983–2008 suggest an oscillatory pattern of 

uplift and subsidence within the caldera and along the north 

caldera rim (fig. 19). The northeast floor of the caldera rose 

during 1976–84, subsided during 1975–95, rose during 1995–

97, subsided during 1997–2002, rose rapidly but at a declining 

rate during 2004–8, and had paused by late 2009. Meanwhile, 

the north rim area rose during 1997–2002, subsided during 

2002–8, and also had paused by late 2009 (both 2009 results 

from CGPS observations). Generally speaking, the caldera 

floor rose while the north rim subsided and vice versa. Data 

for the Mallard Lake dome are not included in figure 19, 

owing to the paucity of leveling data there, but we know from 

leveling surveys in 1986 and 1987, and from InSAR observations since 1992, that the two domes tend to move up or 

down in unison. The only known exception occurred during 
Discussion 37

25

0 15

0.00 Volume Change (x0.0001) -0.35

30 45 60

Distance North (km)

Distance East (km)

C. 2000-2001 D. 2001-2002

75

41

57

73

89

105

25

Distance North (km)

A. 1992-1995

41

57

73

89

105

B. 1996-2000

0 15 30 5 60

Distance East (km)

75

Figure 17. Diagrams of estimates of subsurface volume decrease at 6–8 km depth within the Yellowstone caldera, Yellowstone National 

Park, based on a best-fit model of GPS and InSAR observations. Time intervals shown are (A) 1992–1995, (B) 1996–2000, (C) 2000–2001, 

and (D) 2001–2002. For each time interval, the model layer at 6–8 km depth was the uppermost one containing volume decreases that 

the authors regarded as significant. Color scale is in terms of fractional volume change, which is unitless. Compare to figure 18 in this 

paper, which is based on an independent study by Wicks and others (2006). Figure reproduced from Vasco and others (2007, fig. 5).
38 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

CANADA

Pacific 

Ocean

0 10 20 KILOMETERS

0 5 1O MILES

Figure 18. Map of structural, thermal and volcanic features in and around the Yellowstone caldera, Yellowstone National Park. Red circles 

mark locations of known volcanic vents that erupted after the caldera-forming event at 0.64 Ma (Christiansen, 2001). Areas of known past 

or present thermal activity are colored yellow. The part of the caldera bounded by ring-fractures is shown green, and the slumped zone 

between the ring-fracture zone and the best estimate of the caldera rim is shown in salmon. The National Park boundary is the dashed 

black line. Faults active in the Quaternary are marked with black lines. The labeled features are Norris Geyser basin (NGB), Mammoth 

Hot Springs (M), Sour Creek resurgent dome (SC), Mallard Lake resurgent dome (ML), Hebgen Lake (HL) and Yellowstone Lake (YL). White 

arrows show interpreted magma migration paths. Red square in the inset map (bottom right) shows the location of the study area. Compare 

to figure 17, which is based on an independent study by Vasco and others (2007). Figure from Wicks and others (2006).
Discussion 39

1976

-50

-40

-30

-20

-10

0

10

20

30

40

50

60

70

80

90

100

1980 1985 1990 1995 2000 2005 2010

Year

Surface displacement rate, north caldera, in millimeters per year

Surface displacement rate, Sour Creek dome, in millimeters per year

-50

-40

-30

-20

-10

0

10

20

30

40

50

60

70

80

90

100

Sour Creek Dome, InSAR

Sour Creek Dome, Leveling

North Rim Source, InSAR

Subsidence Subsidence Subsidence

Uplift

Uplift Uplift

Figure 19. Diagram showing recent history of vertical surface displacements from leveling and line-of-sight surface displacements 

from InSAR at Yellowstone’s Sour Creek resurgent dome (black and blue diamonds) and north caldera rim area (red squares). 

Inset, 2004–6 interferogram from figure 13, showing locations of north rim and Sour Creek dome deformation sources. For the Sour 

Creek dome, black diamonds represent leveling data and blue diamonds represent InSAR data. For both types of data, the average 

surface displacement rate during a time period spanned by two sequential observations is represented by a pair of symbols, with 

corresponding error bars, connected by a horizontal line that spans the time period. Since 1997–2000, when uplift of the north rim area 

was first detected using InSAR, there has been a tendency for uplift of the north rim to be accompanied by subsidence of the Sour 

Creek dome and vice versa—a pattern denoted here by blue and red arrows indicating the sense of motion of the Sour Creek dome 

and north caldera rim, respectively. Before 1997, there is insufficient information to reconstruct the displacement history of the north 

rim area. See text for discussion.

1995–96, when the Sour Creek dome started to rise while the 

Mallard Lake dome was still subsiding slightly. However, by 

the following year, both domes were moving upward. We caution that the uplift and subsidence rates shown in figure 19 are 

not strictly comparable, for three reasons: (1) the periods of 

observation for leveling and InSAR differ in some cases; (2) 

the InSAR rates shown are average maximum values for the 

Sour Creek dome as a whole, whereas the leveling rates are for 

specific benchmarks near Le Hardys Rapids; and (3) InSAR 

measures line-of-sight range changes, whereas leveling measures vertical surface displacements. Nonetheless, the pattern 

described above generally has held since 1992, when InSAR 

data first became available.

The pattern of oscillatory, out-of-phase uplift and subsidence at the Sour Creek dome and north caldera rim can be 

explained by the models of Wicks and others (2006), Chang 

and others (2007, 2010), and Vasco and others (2007) as 

follows. A period of increased magma supply from a source 

beneath the northeast part of the caldera is expected to cause 

uplift first of the Sour Creek dome and then, as magma or 

an associated pressure pulse migrates southwestward across 

the caldera along subhorizontal pathways (sills or fluid-filled 

lenses), cause uplift of the Mallard Lake dome as well. For 

example, this scenario explains InSAR observations for the 

period 1995–97. If the migration rate is high enough, uplift 

of both domes might begin nearly simultaneously, as shown 
40 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

by CGPS observations in mid-2004. During periods of lower 

magma supply, exsolution of volatiles from magma beneath 

the caldera results in subsidence of the caldera floor if those 

volatiles are lost to the shallow hydrothermal system or to the 

surface. If volatiles are trapped temporarily beneath a selfsealing layer, resulting pressurization contributes to surface 

uplift; subsidence ensues when the seal eventually ruptures 

(Dzurisin and others, 1994; Fournier, 1999, 2004, 2007). 

An example of the latter process might be the decade-long 

subsidence episode that followed an intense earthquake swarm 

near the caldera rim starting in October 1985. Accumulation 

of magma and volatiles beneath the north caldera rim causes 

uplift there until the resulting pressure increase initiates 

an injection into the Norris–Mammoth corridor or into the 

southeastward extension of the Hebgen Lake fault zone. The 

sequence begins again with the next episode of increased 

magma supply beneath the Sour Creek dome. We caution that 

the same patterns might be explained equally well or better 

by other models. Clearly, additional investigation into the 

cause(s) of surface deformation at Yellowstone is needed.

Before moving on to a discussion of current and future 

research topics, we first assess and discount the possibility that thermal contraction of solidified intrusions beneath 

the caldera—rather than depressurization as a result of fluid 

loss—is a primary subsidence mechanism at Yellowstone. 

Dzurisin and others (1990) based their assessment of potential subsidence mechanisms on the thermal energy output 

from Yellowstone’s hydrothermal system, which they took to 

be about 4×1016 cal/yr. Subsequent estimates are in the range 

4.5–6.0 gigawatts (Fournier 1989; Friedman and Norton, 

2007), which corresponds to 3.4×1016−4.5×1016 cal/yr, so the 

calculations by Dzurisin and others (1990) remain valid. A 

typical value for the specific heat capacity, C, of rock is 800 

Joules/kg⋅oC or 190 cal/kg⋅oC. Following Dzurisin and others 

(1990), we attribute Yellowstone’s entire thermal energy flux 

to cooling rock from T=800oC (near the liquidus temperature 

of rhyolite) to T=450oC (maximum estimated temperature 

within the deep hydrothermal system). Using 2,500 kg/m3 for 

the average density  of intrusive rock and ∆T=350oC, we 

can compute the volume of rock V needed to supply Yellowstone’s thermal energy flux dQ/dT:

V = dQ / dt

C� p� ∆T =2×108

 m3=0.2 km3/yr. (1)

 What is the volume change ∆V associated with 

cooling 0.2 km3 of rock by 350oC? The coefficient of 

thermal expansion  for granite is about 8×10-6/oC, so 

the answer is:

∆V = α� ∆T� V =-0.0006 km3

/yr. (2)

Dzurisin and others (1990, p. 260–263) characterized this 

result as “negligible” compared to their best-fit model result 

based on leveling and trilateration data for the period 1985–7. 

No additional trilateration data are available for the period of 

caldera subsidence from 1985 to 1995, but leveling surveys 

across the northeast part of the caldera each year from 1985 

to 1995, except 1994, show that the average 1985–87 subsidence rate of 30±2 mm/yr is representative (within a factor of 

2) of the entire period (19±3 mm/yr for 1985–95). The model 

result for 1985–87 mentioned above was contraction of a 

tabular body located 10±5 km beneath the caldera floor by 

0.019±0.002 km3

, which corresponds to about -0.01 km3/yr. 

So, even allowing for substantial uncertainties in the values 

used for the calculation above and in the model result, thermal 

contraction of crystalline rock beneath the caldera alone is 

inadequate by more than an order of magnitude to explain 

historical subsidence rates observed at Yellowstone.

Dzurisin and others (1990, p. 263–264) went on to 

consider magmatic fluid loss during crystallization of rhyolitic 

magma as a potential subsidence mechanism and came to a 

much different conclusion. In this case, they assumed that the 

entire heat flux at the surface is extracted by crystallizing without cooling rhyolitic magma near the top of the magmatic system. The latent heat of crystallization L for rhyolitic magma 

is about 75 cal/g (75×103 cal/kg), so the volume of magma 

required to supply the known heat flux dQ/dt=4×1016 cal/yr is:

V = dQ / dt

L� ρ

=2×108 m3=0.2 km3/yr. (3)

What is the volume change ∆V associated with crystallizing 0.2 km3

 of rhyolitic magma, assuming that all the 

exsolved fluid produced escapes into the shallow, hydrostatically pressured part of the hydrothermal system? Dzurisin 

and others (1990, p. 263–264) assumed a 7-percent decrease 

in volume associated with crystallization of the magma, 

which corresponds to 0.014 km3 in this case. This is of the 

same order as the best-fit model result of Dzurisin and others 

(1990) for the decrease in source volume during 1985–87 

(0.019±0.002 km3). To verify this result, we checked the 

assumption of a 7-percent volume decrease during crystallization in the following way.

A water-saturated rhyolite melt having a temperature of 

about 850°C and a pressure of 250 MPa (consistent with the 

depths of deformation sources at Yellowstone) would have a 

density of about 2,200 kg/m3

, as calculated using the thermodynamic relations of Ghiorso and Sack (1995) and the program Conflow (Mastin, 2002). By contrast, the density of the 

main crystal phase expected from such a melt, an albite-rich 

plagioclase, at the same temperature and pressure calculated 

using the thermodynamic relations of Berman (1988) with 

Conflow would be about 2,520 kg/m3 (L. Mastin, written 

commun., 2009). The difference in these densities is about 

14 percent, which is twice the value assumed by Dzurisin 

and others (1990). Given this range in values (7 to 14 percent 

volume decrease), it is clear that deriving the known heat flux 

at Yellowstone from the crystallization of 0.2 km3/yr of rhyolite magma would result in a subsurface volume decrease on 
Current and Future Research Topics 41

the order of 0.01 km3/yr, which is comparable to the surface 

volume change rate measured during periods of caldera-floor 

subsidence. Conversely, deriving the heat flux by cooling the 

same volume of rhyolite magma from typical melt temperature 

to near its solidus temperature would produce a volume contraction 1–2 orders of magnitude smaller, which is insufficient 

to caused the observed subsidence rate.

Current and Future Research Topics

Among the issues raised by recent studies (for example, 

Wicks and others, 2006; Chang and others, 2007, 2010; Puskas 

ansd others, 2007; Vasco and others, 2007; Fournier, 2007) are 

(1) the nature of the fluid(s) involved in caldera deformation 

(magma, brine, gas, multiphase mixtures); (2) the nature of 

the interactions among the basaltic magma system, overlying 

rhyolitic magma body, hydrothermal system (deep and shallow 

components), and structural elements such as the Elephant 

Back fault zone, Norris–Mammoth corridor, postcaldera vent 

alignments, two resurgent domes, and caldera ring fracture 

system; and (3) the rate at which such interactions occur, 

which in turn raises the issue of the mobility of various types 

of fluids in the subcaldera environment. These frontier issues 

are being addressed by long-term, time-series measurements 

of deformation, gravity, seismicity, and magmatic-volatile 

flux (chloride, CO2), and by numerical simulations, laboratory 

experiments, and theoretical studies. A few lines of investigation that seem particularly promising are highlighted below.

Nature of Fluids Involved in Surface 

Deformation

What fluids are likely to exist at depths of 6–16 km 

beneath the three known deformation sources at Yellowstone, 

that is, beneath both resurgent domes and the north rim of 

the caldera near Norris? What characteristics of that depth 

range at those locations are conducive to the observed patterns 

of ground deformation? Are there testable hypotheses to be 

addressed with future experiments that might shed additional 

light on the causes and implications of ongoing deformation? Before addressing those questions, some clarification of 

terminology is necessary. By “magma,” we mean a mixture of 

molten or partially molten rock that generally comprises four 

phases—residual melt, hypersaline brine, gas, and crystals. At 

Yellowstone, the dominant magmas are of basaltic or rhyolitic 

composition. The basaltic magma system extends from the 

mantle upward to the mid-crust, probably to a depth of about 

12 km. A large body of partially molten, but mostly crystallized, rhyolite magma capping the basaltic magma system 

at depths of 6–12 km has been inferred from seismic results 

(Miller and Smith, 1999; Husen and others, 2004). Chu and 

others (2010) analyzed first-arrival waveforms from broadband seismometers and concluded that the melt fraction is 

much greater than previously thought; that is, they hypothesized a subcaldera magma body with a volume greater than 

4,300 km3

 and a porosity of about 32 percent—the latter filled 

with at least 90 percent rhyolite melt and 8 percent water-CO2

by volume. Their interpretation is controversial, but nonetheless there is consensus that heat from the rhyolite body, and 

ultimately from the basaltic system as well, is extracted to the 

surface through an extensive hydrothermal system (Lowenstern and Hurwitz, 2008). A self-sealing layer near the base 

of the hydrothermal system, at about 5 km depth, separates 

fluids at lithostatic pressure from fluids at hydrostatic pressure (Fournier and Pitt, 1985; Fournier, 1991, 1999). We 

use the term “magmatic fluid” to describe hypersaline brine 

and magmatic gas that exsolve from the melt as it cools and 

crystallizes. These are mostly confined to the upper part of the 

magmatic system and lower part of the hydrothermal system, 

beneath the self-sealing layer, except at times when the seal is 

ruptured. At those times, magmatic fluid mixes with fluid from 

the shallow part of the hydrothermal system, which is mostly 

of meteoric origin. We avoid using the term “hydrothermal 

fluid,” which has been applied to fluids of both magmatic and 

meteoric origin. Instead, we distinguish between magmatic 

fluid in the deep part of the hydrothermal system and mainly 

meteoric fluid in the shallow part.

All of the deformation sources proposed by Wicks and 

others (2006) and Chang and others (2007, 2010) are in the 

depth range 6–16 km, which puts them either in the crystallizing rhyolite body or near the top of the underlying basaltic 

system. The magma fluxes required to explain deformation 

rates observed in the caldera and along the north rim (0.01–0.1 

km3

/yr; Wicks and others, 2006; Chang and others, 2007) are 

a factor of 3 to 30 times less than the long-term average intrusion rate of basalt required to account for the observed flux of 

CO2 gas at the Yellowstone caldera (0.3 km3/yr, Lowenstern 

and Hurwitz, 2008). Lowenstern and Hurwitz (2008, p. 37) 

point out that an intrusion rate of 0.3 km3/yr (1) would create a 

maximum power output of 22 gigawatts (GW) (assuming full 

crystallization and further cooling of 300°C), which is 3 to 5 

times greater than the observed convective heat flow at Yellowstone (4.5–6.0 GW, Fournier, 1989; 6.4 GW, Friedman and 

Norton, 2007), and (2) is comparable to, though slightly higher 

than, those estimated for Kilauea and the Columbia River 

Basalts (Gerlach and others, 2002; Lange, 2002). We infer 

that the basaltic magma system beneath Yellowstone is both 

well situated and sufficiently dynamic to host Yellowstone’s 

three known deformation sources and that basalt is the primary 

“fluid” responsible for deformation. We suspect, but cannot 

prove, that apparent discrepancies between (1) the basaltic 

intrusion rate required to explain historical deformation rates, 

(2) the intrusion rate required to account for the observed CO2

flux, and (3) the measured convective heat flow at Yellowstone 

result from time variability of several factors. For example, 

a long-term average intrusion rate of 0.1 km3/yr of basalt is 

consistent with both the average deformation rate measured 

during historical time and the observed convective heat flow. 

The somewhat greater intrusion rate required to explain the 

observed CO2 flux can be explained by time variability in the 

intrusion rate, the CO2

 flux, or both. Changes in these rates 
42 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

would not necessarily occur in lockstep, owing to differences 

in the depth of basalt intrusion and to subsurface complexities 

that affect the transit time of CO2 through the upper crust—

complexities which themselves change with time as a result 

of sporadic earthquake swarms and processes occurring in the 

deep hydrothermal system. Within measurement uncertainties 

and allowing for time variability of processes that are occurring in a manifestly dynamic system, we conclude that the 

models proposed by Wicks and others (2006) and Chang and 

others (2007, 2010) do a reasonably good job of explaining the 

recent deformation history at Yellowstone.

The conceptual model proposed by Wicks and others 

(2006) involves nearly continuous movement of basaltic magma 

into and out of the mid-crust beneath the Yellowstone caldera, 

through fluid connections among three interacting magma-storage zones—one beneath each resurgent dome and a third beneath 

the north rim near Norris (fig. 18). Magma flux into the system 

from a deep source beneath the Sour Creek dome is allowed to 

vary with time, as is the flux out of the caldera toward two extensional “sinks”—the Norris–Mammoth corridor and a seismic belt 

east of the Hebgen Lake fault zone—by way of the north-rim 

deformation source. The seismic belt is coincident with a zone 

of crustal extension between Hebgen Lake and the Norris area, 

which was hypothesized by Savage and others (1993) to explain 

the results of trilateration measurements from 1974 to 1987. The 

model of Wicks and others (2006) is flexible enough to explain 

the full range of observed deformation patterns. For example, 

Wicks and others (1998) reported that uplift started at the Sour 

Creek dome during 1995–96 and migrated toward the Mallard 

Lake dome during 1996–97, whereas Chang and others (2007, 

2010) noted that the entire caldera began to rise nearly simultaneously in mid-2004 following a period of subsidence. The 

difference can be explained with the model of Wicks and others 

(2006) by invoking differential time histories of magma input 

from a source beneath the Sour Creek dome and magma/fluid 

output to the Norris–Mammoth corridor or Hebgen Lake area. 

Similarly, the model can explain an otherwise counterintuitive 

fact—that the hottest and most active thermal areas at Yellowstone (Norris Geyser Basin and Hot Springs Basin)—are located 

outside the caldera (Werner and others, 2008). Hot Springs Basin 

is located above the model’s proposed source of basaltic magma, 

and Norris Geyser Basin is above a proposed injection site for 

magma and exsolved fluid along the Norris–Mammoth corridor. 

A third example is the association between notable earthquake 

swarms and changes in caldera-floor deformation mode, such as 

during 1984–86 (onset of subsidence and October 1985–early 

1986 swarm) and during 1995–96 (onset of uplift and June–July 

1995 swarm). Both swarms occurred just outside the northwest 

caldera rim. Wicks and others (2006, p. 74) wrote: 

The flux of magma out of the Yellowstone system is 

controlled by extra-caldera tectonic activity acting 

on fractured rock bordering the northern caldera 

boundary. Tectonic strain can either enhance or 

restrict the flow of magma out of the caldera.

In so doing, strain along the north caldera rim might initiate 

either subsidence or uplift of the caldera floor, triggering an 

earthquake swarm in either case.

We acknowledge that the relationship between earthquake 

swarms and deformation at Yellowstone is not straightforward. 

Swarms are common occurrences and most swarms are not 

accompanied by a change in the rate or mode of deformation 

(Smith and others, 2009; Farrell and others, 2009). Nonetheless, we are struck by the fact that the two largest swarms ever 

recorded in the Yellowstone area, the 1985–86 swarm and the 

2010 Madison Plateau swarm (http://volcanoes.usgs.gov/yvo/

publications/2010/10swarm.php, accessed September 14, 2011), 

were accompanied by a change from uplift to subsidence within 

the caldera. Furthermore the June–July 1995 swarm, which at 

the time was the largest one to occur since the 1985–86 swarm, 

was accompanied by a change from caldera-floor subsidence 

to uplift during 1995–97. All three of these notable swarms 

occurred near the northwest caldera rim, which suggests to us 

that the area is an important structural element that episodically 

influences deformation modes within the caldera.

The model proposed by Wicks and others (2006) is consistent with other observations, including (1) sill-like sources 

of inflation and deflation in the depth range 6–16 km, which 

corresponds to the top of a low-velocity zone for seismic P 

waves that was identified by Husen and others (2004); (2) 

basalt flux beneath the caldera in the range 0.01–0.1 km3

/yr 

to account for observed deformation rates (see below); and 

(3) coequal amounts of uplift and subsidence during postglacial time, which would be expected if, over long time scales, 

magma influx beneath the Sour Creek dome is balanced by 

efflux to the Norris–Mammoth corridor and to the Norris–

Hebgen Lake seismic belt and extension zone. Such a balance 

does not preclude substantial heat input to the hydrothermal 

system from basalt that transits the caldera. We point out that 

the basalt-flux estimate of 0.01–0.1 km3/yr in the depth range 

6–16 km by Wicks and others (2006) to account for deformation rates does not preclude a greater flux into the lower to 

mid-crust, as proposed by Lowenstern and Hurwitz (2008) to 

account for CO2 flux and heat flow (0.3 km3/yr).

As noted by Wicks and others (2006), a viable alternative 

to the magma-transit model is a shallower, distributed deformation source that corresponds to the deep hydrothermal system. Pierce and others (2002) favored such a model, in which 

magmatic fluid alternately accumulates within and escapes 

from the deep hydrothermal system, to explain the postglacial deformation record. Referring to the offsetting amounts 

of uplift and subsidence during postglacial time, Pierce and 

others (2002, p. 24) concluded: “The inflation-deflation cycles 

seem to represent an essentially zero sum process with little 

net subsurface volume change, which seems most readily 

explained by buildup and release of hydrothermal fluids [magmatic fluids, in the nomenclature used here].” We note that 

the magma-transit model produces a zero sum by balancing 

magma influx and efflux, whereas the magmatic fluid model 

does so by alternately accumulating and dispersing brine and 
Current and Future Research Topics 43

gas. A hybrid model, in which uplift is caused primarily by 

magma intrusion and subsidence is caused by injection of 

exsolved magmatic fluid into the Norris–Mammoth corridor 

or into the eastward extension of the Hebgen Lake fault zone, 

is consistent with the isostatic gravity map for Yellowstone 

National Park and vicinity (Carle and others, 1991). Repeated 

intrusion of basalt into one or both of those extensional zones 

likely would produce an isostatic gravity high, which has not 

been observed. Intrusion of relatively low-density magmatic 

fluid into those zones would produce no such gravity high outside the caldera, and the gravity high associated with residual 

basalt intruded beneath the caldera could be masked by the 

gravitational effect of low-density caldera fill, as interpreted 

by Carle and others (1991, p. 12).

We favor the hybrid model because it explains more 

readily some of the complexities of the observed deformation 

pattern. We envision fluid exsolution from a stagnant, cooling 

rhyolitic magma body beneath the caldera as a steady, gradual 

process that is unlikely to produce the short-term changes in 

deformation mode and rates that have been observed during 

the past decade. Likewise, we would expect self-sealing of a 

layer near the base of the hydrothermal system to be a slow 

process that would result in a more gradual onset of uplift than 

was observed, for example, in mid-2004. These concerns are 

lessened if the source of magmatic fluid is basalt rising and 

accumulating beneath or within a crystallizing rhyolite body. In 

that case, any variation in the rate of magma intrusion would be 

reflected in the rate of fluid production, and a discrete intrusion 

beneath the Sour Creek dome might result in an uplift episode 

with a rapid onset (assuming rapid bubble growth and ascent). 

The rate at which basaltic magma, or fluid derived from that 

magma, migrated southwestward beneath the caldera might 

vary over time, depending on the state of tectonic stress or on 

conditions in the deep hydrothermal system that affect permeability. In this scenario, a time-varying supply of basaltic 

magma to mid-crustal levels beneath the Sour Creek dome is 

necessary, and even sufficient, to explain rapid-onset uplift 

episodes. In other words, to call on a time-varying rate of fluid 

accumulation beneath the Sour Creek dome as the primary 

uplift mechanism, we require a time-varying supply of basaltic 

magma that, by itself, could account for uplift. The situation is 

less clear elsewhere in the caldera. Presumably, lateral intrusion 

of magmatic fluid alone could account for uplift of the Mallard Lake dome and, as a result of temporary reductions in the 

rate of fluid outflow beneath the north rim, of that area as well. 

Subsequent intrusion of magmatic fluid into the Norris–Mammoth corridor or into the extension of the Hebgen Lake fault 

zone would result in subsidence. As mentioned earlier, this is 

the mechanism proposed by Waite and Smith (2002) to explain 

some unusual characteristics of the 1985–6 earthquake swarm 

near Madison Junction, a swarm which was associated with the 

onset of subsidence in the caldera.

Recent basaltic intrusions beneath the northeast part 

of the caldera are plausible for two reasons—(1) continued 

northeastward migration of the Yellowstone hotspot since the 

last caldera-forming eruption and (2) uplift that began at the 

Sour Creek dome in 1995–96, which Wicks and others (2006) 

attributed to a pulse of basaltic magma from a deep source 

beneath the dome. At an average migration rate of 4.5 cm/yr, the 

hotspot has moved approximately 30 km farther northeast since 

the Yellowstone caldera formed 0.64 Ma—toward the Sour 

Creek dome and the northeast caldera rim in the vicinity of Hot 

Springs Basin. This is the area where uplift began in 1995–96. If 

we accept that a time-variable supply of basaltic magma from a 

mantle source beneath the Sour Creek dome is a likely cause of 

uplift there, either directly or as a result of fluid-phase separation, then we can account for deformation elsewhere in the caldera through lateral migration of magma, a fluid phase, or both. 

Either could accumulate in fractured zones beneath the Mallard 

Lake dome and along the north caldera rim, causing uplift, and 

then migrate into one or more sinks outside the caldera, causing 

subsidence. Magma would be less mobile than the fluid phase, 

so we would expect the proportion of magma/fluid to decrease 

“downstream”—that is, from the source beneath the Sour Creek 

dome, to accumulation zones beneath the Mallard Lake dome 

and north caldera rim, to sinks along the Norris–Mammoth corridor and Hebgen Lake fault zone. The absence of postcaldera 

basalt flows inside the caldera is explained by the presence of a 

partly molten body of rhyolite that impedes the buoyant rise of 

basalt (Miller and Smith, 1999; Husen and others, 2004; Chu 

and others, 2010). Outside the caldera, we might expect to see 

basalt escaping to the surface along the Norris–Mammoth corridor or in the vicinity of the Hebgen Lake fault zone. Christiansen (2001) mapped the postcaldera Swan Lake Flat Basalt 

and the Madison River Basalt in those two areas, respectively. 

However, postcaldera basalts also occur elsewhere around the 

periphery of the Yellowstone Plateau (Christiansen, 2001, p. 

G48–G49). 

Structural Control of Deformation Sources

The second issue raised above (What is the nature of 

the interactions among the basaltic magma system, overlying 

rhyolitic magma body, hydrothermal system (deep and shallow 

components), and structural elements such as the Elephant 

Back fault zone, Norris–Mammoth corridor, postcaldera vent 

alignments, two resurgent domes, and caldera ring fracture 

system?) deals with the locations of deformation sources in 

structurally distinctive areas, that is, beneath each resurgent 

dome and at the three-way intersection of the caldera ring 

fault, the Norris–Mammoth corridor, and the Norris–Hebgen Lake seismic belt. How might these structural elements 

interact with Yellowstone’s magmatic and hydrothermal 

systems to produce the spatial and temporal deformation patterns described above? We hypothesize that enhanced fracture 

permeability at all three locations facilitates fluid storage and 

migration, which is a prerequisite for inflation-deflation cycles 

that produce surface deformation. This is true regardless of the 

nature of the fluid involved (magma or brine), so the observation does not favor either model. Likewise, either model can 
44 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

account for the existence of porous, gas-filled rock at depths of 

less than 2 km between Norris and Hebgen Lake, which Husen 

and others (2004) inferred from anomalously low values of Vp

and Vp/Vs in that area.22 The close spatial correlation between 

the anomaly and the epicentral area of the 1985–86 earthquake swarm, which had several characteristics that Waite 

and Smith (2002) attributed to fluid migration, led Husen and 

others (2004, p. 409) to propose “* * * a model in which CO2

as part of magmatic fluids exsolved from a large crystallizing 

magma body beneath the Yellowstone caldera and occasionally migrated outwards * * *.” We concur, but suggest that any 

CO2 that migrated out of the caldera during the 1985–86 earthquake swarm could have exsolved instead, or in part, from 

basalt that intruded beneath the northeast part of the caldera.

We acknowledge a counterpoint to this idea made by 

Evans and others (2006), who conducted a geochemical survey 

of springs outside the northwest margin of the caldera in 2003 

and 2004. The survey was designed specifically to look for 

increased CO2 flux that might be associated with the onset of 

caldera subsidence in 1985–86 or with the 1985–86 earthquake 

swarm. Evans and others (2006, p. 169) found: “* * * no evidence to associate the ten-year period of caldera deflation that 

began in 1985 with expulsion of magmatic fluids through the 

caldera rim in this area.22 They went on to say (p. 179):

Our results are compatible with the involvement 

of pressurized hydrothermal fluids in the 1985 

swarm, but they suggest that if caldera subsidence 

resulted from large-scale fluid leakage, the fluid 

mostly escaped elsewhere. An intense but short-lived 

geochemical anomaly cannot be ruled out, given the 

timing of sampling surveys, but other potential fluid 

flow paths that might be active during caldera subsidence at Yellowstone need further investigation. We 

suggest that intracaldera CO2 efflux, which is currently unmonitored apart from the one-time survey 

of Werner and Brantley (2003), might track caldera 

deformation cycles even if riverine Cl does not.

Lack of a “smoking gun” in the form of increased CO2

flux in the 1985–86 epicentral area in 2003–4 does not preclude the mechanism proposed by Waite and Smith (2002) and 

Husen and others (2004), but neither does it lend support to 

this aspect of our preferred model.

Temporal Gravity Observations as a Means to 

Distinguish Between Magma and Other Fluids

The third issue raised above deals with the relative 

mobilities of various fluids (gas, brine, magma) beneath the 

Yellowstone caldera and with testable hypotheses or observations that might distinguish between the magma-transit and 

magmatic fluid models. Both models predict the episodic release 

of magmatic fluid into the shallow part of the hydrothermal 

system or into fracture zones north of the caldera. A temporal 

correlation between earthquake swarms, periods of subsidence, 

and increased surface flux of magmatic volatiles such as chloride and CO2

 would favor the magmatic fluid model, because 

subsidence and rupture of the self-sealing layer are directly 

linked in that model but not in the magma-transit model. In the 

latter case, subsidence might occur whenever the magma flux 

is relatively low, as a result of volatile loss from crystallizing 

magma, even without rupture of a self-sealing layer.

The most convincing observations to date of a link 

between seismicity and fluid flow at Yellowstone were 

reported by Pitt and Hutchinson (1982) and Waite and Smith 

(2002) for two notable earthquake swarms. The first occurred 

during May–November 1978 beneath the Mud Volcano 

hydrothermal area, a vapor-dominated, acid-sulfate system 

located at the intersection of the Elephant Back fault zone and 

the Sour Creek resurgent dome (fig. 1). The area produced 

fewer than 10 located earthquakes from 1973 to April 1978, 

but starting in May 1978 numerous earthquakes began occurring there at 1–5 km depth. Activity culminated with intense 

swarms of about 100 events/hour on October 23 and November 7; the largest event (M 3.1) occurred on November 13 and 

at least 8 events were M 2.5 or larger. By December 1978, 

heat flux in the area had increased sufficiently to kill lodgepole 

pine trees and other vegetation. From January to May 1979, 

existing mud volcanoes became more active with increased 

turbidity, surging activity, and explosive ejection of mud; and 

new mud volcanoes and fumaroles formed. Activity began to 

decline in July 1979 and had returned to preswarm levels by 

June 1980. Pitt and Hutchinson (1982, p. 2765) concluded 

that fluids were involved in the event but declined to speculate whether they were sourced in the hydrothermal system or 

magmatic system:

The spatial and temporal association of earthquakes 

and increased heat flux at Mud Volcano is the best 

defined so far in Yellowstone. The earthquakes preceded the increase in heat flux and may have occurred 

on faults that control the convective fluid flow into 

the area. The entire hydrothermal area was affected, 

not only selected features, and the increase in heat 

flux lasted for over a year. Accumulated tectonic 

strain in the Mud Volcano area could have caused 

the earthquakes, which in turn expanded the plumbing system of the hydrothermal area, permitting a 

temporary increase in the normal convective fluid 

flow. Alternatively, if the uplift in the eastern caldera 

is due to pressure and temperature variations within 

a magma chamber, as suggested by Pelton and Smith 

[1979], hot fluids migrating upward in the crust as a 

result of the pressure in the uplift could have entered 

 22Vp

 is P-wave velocity, Vs is S-wave velocity. Husen and others (2004) proposed that the combination of low Vp

 and low Vp/Vs is indicative of a change in 

pore fluid from liquid to gas, likely CO2, at shallow crustal depths.
Current and Future Research Topics 45

the seismogenic zone and triggered the earthquake 

sequence. If the uplift and the high 3He/4He ratio at 

Mud Volcano are at least partly due to the injection of 

new magma into the eastern caldera, the seismic heat 

flux variation in the hydrothermal area may provide a 

sensitive monitor of this process.

A case can be made for similar but more widespread 

responses of Yellowstone’s hydrothermal system to the 

October 1985–early 1986 earthquake swarm near Madison 

Junction (Waite and Smith, 2002). The timing of that swarm 

was roughly coincident with the onset of caldera subsidence sometime between leveling surveys in September 1985 

and September 1986 (Dzurisin and others, 1990, 1994). The 

swarm was preceded in ~July 1985 by an explosion near Pelican Cone, about 5 km east of the east caldera rim (fig. 1), that 

formed a 20 m3 crater, killed mature trees, and gave rise to a 

new superheated fumarole. About 6 months later and 3 months 

after the start of the swarm, an explosion in mature forest 3 km 

west of Norris Junction and 10 km north of the north caldera 

rim formed a 750 m3

 crater, knocked down trees, and threw 

debris 35 m laterally. Although the swarm was over by the end 

of 1986, subsidence of the caldera floor continued for a decade 

until uplift resumed in 1995–96. During 1990–93, ground 

temperatures increased in part of Hot Springs Basin thermal 

area near the northeast caldera rim and both a new superheated 

fumarole and mud volcano emerged (Hutchinson, 1993).

If the primary cause of caldera subsidence is depressurization and loss of magmatic fluid from the deep hydrothermal system as a result of a self-sealing zone rupturing during 

an earthquake swarm, then we might expect to observe an 

increase in the surface flux of magmatic gases shortly after 

the swarm and the onset of subsidence. Concurrent changes 

in hydrothermal basins, such as those described above, might 

also occur if hot, upwardly mobile magmatic fluid were to 

encounter the shallow hydrothermal system. There is as yet 

no evidence for increased emission of magmatic gases or 

chloride flux during subsidence episodes, and the changes in 

hydrothermal activity associated with the 1978 and 1985–86 

earthquake swarms, although unusual, are not unique during 

historical time at Yellowstone. In the first case, uncertainties 

in the chloride-flux measurements combined with the buffering effect of the shallow hydrothermal system tend to downplay the significance of the negative result (Fournier, 2004). 

Furthermore, Hurwitz and others (2007b, p. 168) argue that 

any correlation between the onset of subsidence and increased 

chloride flux is highly unlikely, because transport of chloride to the surface could take thousands of years. Nonetheless, we suggest that monitoring of chloride and CO2

 fluxes 

be continued and improved, to the extent possible, through 

multiple inflation-deflation cycles. Chloride flux estimates 

have been made each year since 1983, except 1995 and 1996, 

by combining water discharge measurements at streamgaging stations with chloride concentrations in samples obtained 

at the gaging stations (Friedman and Norton, 2007). Surface 

CO2

 flux has been estimated by Werner and Brantley (2003) 

using a painstaking sampling and statistical approach. Neither 

measurement is easy to automate or obtain in real time, but 

progress in this regard would be extremely important. With 

regard to the incompletely understood relationship between 

earthquake swarms and changes in hydrothermal activity, we 

urge continued documentation of Yellowstone’s unparalleled 

hydrothermal features, in the tradition established by park 

geologist Rick Hutchinson before his untimely death in 1997. 

Fortunately, park managers and staff have seen fit to continue 

that important work in his absence.

An obvious difference between the magma-transit and 

magmatic fluid models for caldera deformation is the density 

of the fluid involved—a factor of 2–3 between hypersaline 

brine and magma. For that reason, among others, the two 

models predict different gravity signatures through time, 

suggesting another testable hypothesis. Battaglia and others 

(2003) modeled changes in gravity, surface height, and watertable level at the Long Valley caldera, California, to constrain 

the density of fluid responsible for surface inflation during 

1980–99. They interpreted their result, 1,713 kg/m3

 (best fit) or 

1,180–2,330 kg/m3 at 95-percent confidence, as evidence for 

intrusion of low-density magma or a combination of magma 

and low-density fluid. By adding additional constraints on the 

deformation model from InSAR observations, Tizzani and 

others (2009) revised the density estimate for the intrusion 

to 2,509 kg/m3 (best fit) or 2,192–3,564 kg/m3 at 95-percent 

confidence. They noted that the estimate is strongly dependent 

on the shape and location of the source model, but argued 

that these parameters were sufficiently well constrained to 

conclude: “* * * inflation of the resurgent dome from 1982 

through 1999 was the result of an intrusion of basaltic to 

silicic magma” (Tizzani and others, 2009, p. 66).

A recent microgravity study at Kīlauea Volcano, Hawai‘i, 

provides additional evidence for the technique’s unique 

capability to detect subsurface mass changes. Johnson and 

others (2010) reported a net increase in residual gravity (that 

is, after free-air correction) of 521 μGal during 1975–2008 

at a benchmark near the east rim of Halema‘uma‘u Crater in 

the volcano’s summit area. The change, which is a factor of 

about 20 times greater than the combination of measurement 

uncertainty plus any plausible water-table effect, occurred 

progressively between gravity surveys in 1975, 1981, 1998, 

2003, and 2008. Concurrent leveling surveys show that the 

mark subsided progressively by a total of nearly 2 m from 

1975–2008, a period that included nearly continuous lava 

extrusion from vents along Kīlauea’s East Rift Zone starting in 

January 1983. The combination of net subsidence and positive 

residual gravity is unambiguous evidence for an increase in 

subsurface mass, which Johnson and others (2010) interpret as 

magma accumulation at 0.5–1.0 km depth. On 19 March 2008, 

a new eruptive vent formed along the southeast margin of 

Halema‘uma‘u Crater, within 1 km of the surface projection of 
46 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

the gravity anomaly source model. Johnson and others (2010, 

p. 1142) concluded: “The eruptive activity was apparently preceded by decades of magma accumulation that gradually filled 

void space, and that was detected by gravity measurements but 

unknown from deformation monitoring alone.”

Arnet and others (1997) used a similar approach, that 

is, microgravity and leveling surveys, to infer a subsurface 

mass increase at Yellowstone during 1977–83, a period of 

net surface uplift. However, in this case, a lack of water-table 

data made it impossible to estimate the density of the intruding material. Short of monitoring water-table fluctuations in a 

network of newly drilled wells within Yellowstone National 

Park, which is impractical, the next best approach would be 

to install a small network of continuously recording gravity 

meters and GPS stations at key locations throughout the park 

and to monitor gravity and height changes during several inflation-deflation cycles. Such a network might be cost prohibitive 

and would raise serious permitting and logistics issues. A more 

viable alternative for the time being is to make repeated, very 

precise, absolute gravity measurements using a portable gravity meter, then estimate and remove the effects of water-table 

changes using a groundwater model constrained by precipitation and lake-level records. Hurwitz and others (2007a, p. 12), 

on the basis of the results of their numerical simulations of 

fluid flow, suggested: “Continuous, high-precision microgravity measurements may discriminate between magma intrusion 

and hydrothermal injection at shallow depths, because the density of magma differs by a factor of 3 or more from the density 

of superheated vapor or gas * * *.” We believe this approach, 

combined with continued improvement of four-dimensional 

(4-D) (X,Y,Z, time) subsurface imaging methods, holds the 

greatest promise for ultimately indentifying the correct deformation mechanism(s) at Yellowstone and for assessing the 

implications for associated hazards.

Four-Dimensional Numerical Simulations and 

Modeling—A Way Forward

Numerical simulations of heat and mass flow in porous, 

nonisotropic, nonelastic media (for example, Hurwitz and 

others, 2007a; Hutnak and others, 2009) are an essential 

step toward better understanding of ground deformation at 

Yellowstone and other large calderas. Only process-oriented 

models such as these can provide insight into the mechanism 

of ground deformation, in addition to insight into the geometrical parameters of the source. Knowing with confidence 

that a lens-shaped deformation source exists 8 km beneath 

a caldera’s floor is useful. Knowing whether such a source 

represents magma, brine, gas, or some combination thereof 

is essential for improved understanding and hazards assessments. Even idealized models of deformation in a homogeneous, semi-infinite, elastic half-space might give reasonable 

estimates of source depth, size, and shape in spite of the 

unrealistic assumptions involved, but full confidence in those 

results awaits verification by models that account for known 

complexities in the real Earth. The studies by Hurwitz and 

others (2007a) and Hutnak and others (2009) address such 

complexities in a theoretical sense and provide a framework 

for additional advances. The next step is to incorporate 

more detailed information about the actual 3-D distribution 

of mechanical properties in the crust beneath Yellowstone. 

Such information can be gleaned from multidisciplinary 

investigations such as one by DeNosaquo and others (2009), 

who combined gravity and heat-flow data with tomographic 

seismic velocity models, GPS-derived strain rates, earthquake locations, and chemical analyses of volcanic rocks to 

constrain models of density and lithospheric strength for the 

Yellowstone-Snake River Plain system.

Laboratory experiments to investigate processes of 

bubble nucleation, growth, and migration in magmas of 

mafic and silicic composition, under pressure and temperature conditions that prevail in the crust beneath calderas, 

are another important piece to the Yellowstone deformation 

puzzle. Also needed is a better means to incorporate experimental results into physical models of deformation processes 

at caldera scale. In numerical simulations by Hurwitz and 

others (2007a) and Hutnak and others (2009), volatiles are 

“sourced” from a crystallizing magma body to the base of a 

hydrothermal system—without consideration of processes 

by which volatiles exsolve, migrate upward, and eventually 

escape into the overlying hydrothermal system. In an ongoing series of piston-cylinder decompression experiments, 

Mangan and Sisson (2000, 2005) are exploring the relative 

importance of decompressional expansion, coalescence, and 

Ostwald ripening23 as processes of bubble growth and migration in rhyolitic magma. Working along similar lines, Lensky 

and others (2004) developed a theoretical model of bubble 

growth during decompression of supersaturated melt and 

used decompression experiments to verify essential elements 

of the model. Together with heat and mass transport through 

porous media, the mechanisms and time scales of gas escape 

from magmas are important aspects of any comprehensive 

explanation of ground deformation at Yellowstone.

Also important is better characterization of the subcaldera crust in terms of structure (faults, fractures, layering), 

temperature, permeability, pore pressure, strength (shear 

modulus), and gas content. These characteristics are amenable to study by ever-improving seismic (for example, Waite 

and others, 2005; 2006; DeNosaquo and others, 2009) and 

electromagnetic techniques.

23In this context, Ostwald ripening refers to a process in which larger 

bubbles in a magma grow at the expense of smaller ones. Large bubbles have 

a lower surface to volume ratio, which results in a lower energy state. As the 

system tries to lower its overall energy, gas molecules tend to diffuse out of 

small, energetically unfavorable bubbles and into larger ones. As a result, 

smaller bubbles continually shrink, while larger bubbles grow.
Conclusions 47

As more accurate and diverse geophysical and geochemical datasets are acquired, the need for more detailed numerical 

models will increase. Finite-element models can be used to 

account for nonhomogeneous Earth structure and properties, 

and they have the potential to incorporate all relevant information in a flexible, extensible, and updatable system. For 

example, a finite-element model for Yellowstone could include 

information—measured, inferred from modeling, or assumed—

about the 3-D distributions of such parameters as temperature, 

permeability, rheology, and gas content. The model could be 

used to investigate the system’s deformational response to various time-varying inputs, such as earthquake swarms (permeability increase), a temperature pulse (for example, heating by 

magmatic gases), or mass influx (intrusion of magma or other 

fluid). Such a model could continually be updated as new or 

better observational and experimental data are acquired. This 

“system in a box” or “virtual volcano” approach has been 

applied to subduction zones (Masterlark, 2003) and to volcanic 

centers in the Aleutian Arc, Hawaii, and Iceland (for example, 

Masterlark, 2007); a similar effort for the Long Valley caldera 

system is underway (M. Mangan, oral commun., 2009).

In the Yellowstone region, a 3-D finite-element model 

was used to explore the relationship between inflation/deflation 

cycles in the caldera and anomalous slip on the nearby Teton 

normal fault (Hampel and Hetzel, 2008). Leveling surveys 

across the fault were carried out between 1988 and 2001 and 

revealed two phases of hanging-wall uplift relative to the foot 

wall (that is, reverse slip), alternating with two phases of normal 

slip (Sylvester and others, 1991, 2001). Hampel and Hetzel 

(2008) showed that the motion is explained by modification 

of the tectonic-displacement field by inflation/deflation cycles 

in the caldera. Inflation causes enhanced horizontal extension 

across the north-trending Teton fault to the south, which favors 

accelerated normal slip on the fault. Conversely, caldera deflation causes horizontal contraction across the fault and favors 

reverse slip. Hampel and Hetzel (2008, p. 1 of 5) concluded: 

“Our findings imply a strong coupling between magmatism and 

tectonic faulting, which requires coordinated monitoring of both 

processes to improve our understanding of the resulting spatial 

and temporal strain pattern.” Such effects are another aspect of 

the Yellowstone deformation story that is amenable to theoretical and numerical investigations.

Hurwitz and others (2007a, p. 14) used numerical 

simulations to investigate the effects of fluids on groundsurface deformation, and concluded: “On the basis of the 

numerical experiments, and a growing number of observations which indicate a causal link between GSD [ground 

surface displacement] and dynamics of hydrothermal fluids 

and/or gases, we propose that future studies should focus on 

resolving the nature of the fluid causing GSD (for example, 

magma or aqueous fluid and gas).” Hutnak and others (2009) 

extended that work by including multicomponent, multiphase 

flow in their simulations, which confirmed the likely importance of thermoelastic and poroelastic effects in causing 

ground deformation at large calderas.

Conclusions

Even though crustal uplift at the Yellowstone caldera 

was discovered more than 30 years ago, the causes of rapid 

ground motions there are not fully understood. Subsequent 

work has revealed periods of subsidence that mirrors the 

pattern of uplift, identified three interacting deformation 

sources, and explored the likely roles played by magma and 

exsolved magmatic fluid in causing surface deformation. A 

definitive measurement or experiment to determine the relative importance of magma and fluid (magmatic or meteoric) 

in the deformation process as a function of time remains to 

be done. An expanded program of microgravity measurements using both relative and absolute gravity meters is 

especially promising in this regard. Also lacking at this time 

is a comprehensive, 4-D model that accounts for heterogeneous structure and physical properties of the crust beneath 

Yellowstone, characteristics of multiphase flow through 

porous media, and the effects of magmatic volatiles in addition to water and carbon dioxide (that is, sulfur, chlorine, 

hypersaline brine). In this regard, numerical simulations of 

heat and mass flow in porous, nonisotropic, nonelastic media 

(for example, Hurwitz and others, 2007a; Hutnak and others, 

2009) point the way for future studies.

Although acknowledging that many questions concerning 

deformation at Yellowstone are yet to be posed and answered, 

we suggest that the conceptual model offered by Wicks and 

others (2006)—which is generally in agreement with numerical models by Chang and others (2007), Puskas and others 

(2007), and Vasco and others (2007), among others—is a 

useful framework for future studies. The essential elements of 

the Wicks and others (2006) model that can serve as testable 

hypotheses are the following: (1) caldera-floor uplift is caused 

mainly by a variable but nearly continuous supply of basaltic 

magma beneath the northeast part of the caldera (Sour Creek 

dome); (2) basalt and exsolved fluid migrate laterally along 

a subhorizontal boundary, probably near the transition from 

ductile to brittle behavior within a mostly crystalline body of 

rhyolite, and accumulate beneath the southwest part of the 

caldera (Mallard Lake dome) and north rim (Norris area); (3) 

discrete deformation sources beneath the two domes interact 

in such a way that they mostly inflate or deflate in unison, 

but there can be brief exceptions—in such cases, uplift starts 

earlier at the Sour Creek dome because it is closer to the deep 

source of basalt; (4) basalt or fluid exsolved during ascent, 

migration, cooling, and crystallization of basalt is injected episodically into the Norris–Mammoth corridor or Hebgen Lake 

fault zone; and (5) interplay between uplift and subsidence 

at three sites within the caldera is controlled by the relative 

rates of magma input, magma migration, fluid exsolution, and 

injection of magma/fluid outside the caldera—which vary as a 

function of time. The net result of these processes, as documented here and in numerous other studies, is one of the most 

dynamic caldera systems on Earth.
48 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Acknowledgments

 The CVO leveling surveys at Yellowstone from 1983 

to 2007 would not have been possible without the efforts of 

dozens of staff members and volunteers—including crew 

chiefs John Estrem, Gene Iwatsubo, Jack Kleinman, Maurice 

Sako, Dave Wieprecht, and Ken Yamashita—who braved 

fickle weather, wildlife, wildfires, and Winnebagos to ensure 

that the data are of high quality. Numerous colleagues from 

USGS, University of Utah, and Yellowstone National Park—

notably Bob Christiansen, Bob Fournier, Hank Heasler, Jake 

Lowenstern, and Bob Smith—contributed through thoughtful 

discussions and written comments that shaped our thinking and 

corrected our mistakes. Hank Heasler, Wayne Hamilton, and 

Rick Huthchinson—current and former Yellowstone National 

Park geologists—contributed many helpful ideas and acted 

as liaisons between us and park managers to ensure that our 

research was both scientifically sound and consistent with 

park management principles. Two of those managers, former 

Chief of the Yellowstone Center for Resources Tom Olliff and 

his predecessor John Varley, served at the nexus of numerous special-interest groups with competing demands for park 

resources, including us, and successfully accommodated most, 

while adhering to the principles of stewardship that keep Yellowstone a special place for future generations. 

 Larry Mastin helped with calculations of density changes 

during crystallization of a rhyolite magma using his opensource numerical model Conflow. Megan McLay meticulously 

organized a mountain of information concerning more than 500 

benchmarks in Yellowstone National Park for appendix B. She 

created the interactive version for the data files that accompany 

this report and also an on-line version, which will ensure that 

information about this important geodetic resource is available 

to future researchers. Lisa Faust made major improvements 

to early versions of the illustrations in this report and used her 

artistic skills to portray complicated data graphs in ways that 

are both intuitive and precise. 

 Reviewers Shaul Hurwitz, Evelyn Roeloffs, and Cynthia Gardner provided thorough and constructive reviews of 

the manuscript, which significantly improved the final product. Christie Hendrix and Stacey Gunther in the Yellowstone 

Research Permit Office guided us through the permitting process 

and provided essential logistical support. Stephanie Tubman, Liz 

Juers, and Heidi Anderson helped in our search for back-country 

benchmarks. Radar images from European Space Agency (ESA) 

satellites ERS-1, ERS-2, and ENVISAT were acquired through 

ESA Category-1 Project 2765, through an ENVISAT Announcement of Opportunity, and through the WInSAR consortium (supported by NASA, the National Science Foundation, and USGS; 

administered by UNAVCO, http://winsar.unavco.org/main.php). 

 This research was supported by the USGS Volcano 

Hazards Program at the USGS David A. Johnston Cascades 

Volcano Observatory, Washington, USGS Hawaiian Volcano 

Observatory, Hawai‘i, and USGS Menlo Park, California, 

offices. In recent years, our research was conducted under the 

auspices of the Yellowstone Volcano Observatory (YVO)—

a collaboration among USGS, Yellowstone National Park, 

and the University of Utah, which was formed in 2001 to 

strengthen long-term monitoring of volcanic and earthquake 

unrest in the Yellowstone National Park region.

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Appendixes A–B
56 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Appendix A. Uncertainties in Surface Displacement Measurements from 

Repeated Leveling Surveys and InSAR

Uncertainty in Leveling Surveys

Throughout this report, uncertainties in vertical displacement measurements are expressed as one standard deviation 

from the combination of random leveling error and long-term 

benchmark instability. For two first-order, class-II leveling 

surveys, these error sources are approximated by:

√(1.0 mm.

km-1/2.√

—∆L)2 +(1.0 mm.

yr-1/2.√

—∆t)2 σ = ,

where ΔL is stadia distance in kilometers along the level line 

and Δt is time in years between surveys (Vanicek and others, 

1980; Wyatt, 1989). When a first-order, class-II survey and 

second-order, class-II survey are combined, the coefficient 

of the first term, 1.0 mm·km-1/2, is replaced by 2.2 mm·km-1/2. 

This is the case, for example, when the second-order 1923 

survey in Yellowstone National Park is combined with any 

of the subsequent, first-order surveys. The coefficient of the 

second term, 1.0 mm·yr-1/2 in this case, depends on the quality of the benchmark. For geodetic marks set in bedrock or 

on deeply anchored rods, it is generally in the range 0.2–1.0 

mm·yr-1/2 (Wyatt, 1989) or, more conservatively, 0.5–2.0 

mm·yr-1/2 (Langbein and Johnson, 1997). The marks used for 

leveling along the Lake Butte–Mount Washburn line at Yellowstone are of variable quality. Approximately two-thirds of 

the marks are set on bedrock, on substantial structures such 

as bridge abutments, or on deeply anchored rods. These types 

are considered to be high quality marks (Floyd, 1978). Most 

of the others are set on large boulders or culvert headwalls; 

they are of lower quality for long-term deformation studies, 

but useful as recoverable points between marks of higher 

quality. Our choice of 1.0 mm·yr-1/2 in the expression above is 

thought to be conservative for the high-quality marks in Yellowstone and to be generally reflective of the stability of the 

other marks as well.

The uncertainties in vertical displacement measurements 

from repeated leveling surveys are reported here in one of two 

ways, depending on the context. When the intent is to express 

the uncertainty associated with a height-change measurement 

at a given mark with respect to a reference mark, the value of 

ΔL is the cumulative stadia distance between the given mark 

and the reference mark. For example, the uncertainty in a vertical displacement measurement at DA3 1934 near Le Hardys 

Rapids relative to 36 MDC 1976 near Lake Butte based on 

comparison of two first-order, class-II surveys is ±4.9 mm if 

the surveys were separated by one year (ΔL=23.4 km, ∆t=1 

year) or ±5.8 mm if the surveys were separated by 10 years 

(ΔL=23.4 km, ∆t=10 years). Uncertainties reported in the text 

were calculated this way. To calculate the error bars for vertical-displacement profiles shown in several of the figures, we 

used a different approach. In that case, a more useful indicator 

is the uncertainty in tilt measured between adjacent marks. 

This is because the uncertainty calculated as described above 

grows with distance from the reference mark, even though 

height differences measured between adjacent marks that are 

far from the reference mark are no less precise than those 

measured between marks closer to the reference mark. To calculate the error bars shown on vertical-displacement profiles, 

we used the incremental distance between adjacent marks for 

ΔL in the expression above. It might be useful to think of these 

error bars as an uncertainty envelope for the shape of a profile, 

which can move up or down by a greater amount that corresponds to the uncertainty value reported in the text. Uncertainties reported earlier in the literature, including those by other 

authors, are replaced here by values that were calculated as 

described above; any resulting discrepancies are minor.

Uncertainty in InSAR Measurements

The measurement of surface displacements using radar 

interferometry is inherently much more complicated than the 

same measurement made by repeating leveling surveys, and 

therefore specifying the uncertainty in InSAR measurements is 

much more difficult. The fundamental measurement depicted 

in a deformation interferogram is the phase difference between 

two radar pulses backscattered to the SAR from the same 

resolution element on the ground at two different times. For a 

C-band SAR such as the ASAR instrument onboard Envisat, 

which operates at a wavelength λ=56.3 mm, each full fringe 

in a deformation interferogram represents a pulse round-trip 

travel distance of n·λ , where n=1, 2, 3... . Ignoring all other 

effects for the time being, this corresponds to a line-of-sight 

(LOS) range change (surface displacement) of n� λ

2 , or 

about 28 mm for the Envisat ASAR if n=1. Radar phase-difference measurements can be made with a precision of better 

than 1 percent, so in theory InSAR is capable of submillimeter accuracy. However, several factors make it impossible to 

achieve such high accuracy in practice.

In the preceding discussion, we ignored the effects of 

noise in SAR images on the accuracy attainable with InSAR 

measurements. For our purposes “noise” refers to any signal 

other than that resulting from displacement of Earth’s surface. 
 57

Potential noise sources include sensor noise, satellite orbital 

errors, errors in the digital elevation model (DEM) used to 

produce a deformation interferogram, and propagation delays 

caused by inhomogeneity in the troposphere and ionosphere. 

The last factor arises because the speed of radar waves 

depends on the refractive index of the medium through which 

they travel, which in this case is a function of water-vapor 

content in the troposphere and electron density in the ionosphere. In general, the net effect of various noise sources is to 

limit the range-change accuracy of C-band interferometry to 

1–10 mm under favorable conditions.

Spatial covariance in the noise structure of InSAR data is 

another factor that increases the uncertainty of surface displacement measurements made with InSAR. A typical interferogram 

comprises tens of millions of data points, but nearby data points 

are not independent of one another—they result from similar 

paths through the ionosphere and troposphere, and therefore the 

noise components in the information they represent are correlated. Lu and others (2000) have shown that propagation delays 

commonly produce as many as two to three fringes in interferograms of Aleutian volcanoes. This might be an extreme case, 

because the Aleutian Islands are subject to notoriously dynamic 

and wet weather. Even so, the effect of propagation delays 

can be mitigated if several radar images of the same area that 

span similar time intervals are available. Imagine, for example, 

that we form three coherent interferograms, AB, BC, and AC, 

from radar images A, B, and C. Assume that B includes a large 

propagation delay anomaly in part of the image. The anomaly 

will manifest itself as a closed pattern of fringes in interferograms AB and BC, but will be absent from AC. We can reasonably associate the fringe pattern with image B and dismiss the 

possibility that it represents surface deformation. By extending 

this approach to a larger suite of interferograms we can partly 

eliminate spurious effects.

Lohman and Simons (2005) studied the covariance 

structure of noise in InSAR data by comparing several 

independent interferograms with GPS observations of tropospheric delay in the same target area. They concluded that 

downsampling InSAR data by spatial averaging can be an 

effective means to reduce the number of highly correlated 

data points and proposed a resolution-based resampling 

method as an alternative to the approaches used by Jónsson 

and others (2002) and Simons and others (2002). Regardless of the resampling algorithm used, Lohman and Simons 

(2005) stressed the importance of using the full noise covariance matrix to estimate the uncertainty in model parameters 

obtained from inversions of InSAR data.

It should be obvious from the preceding discussion that it 

is not yet possible to specify in a rigorous way the uncertainty 

in displacement measurements made with InSAR. How large 

are propagation delay anomalies in the SAR images used to 

produce the interferograms? Were their effects correctly identified in the interferograms and eliminated from the displacement 

measurements? Which spatial averaging algorithm was used to 

downsample the InSAR data? How was the covariance structure 

of noise in the InSAR data estimated? These questions cannot be answered in a uniform way for a set of interferograms, 

even if all of the interferograms were produced from images 

acquired by the same SAR. On the basis of our own experience 

and consultations with colleagues who specialize in InSAR, 

we assign an uncertainty to the InSAR displacement measurements reported here of ±10 mm. We report InSAR results to one 

significant figure in units of millimeters, rather than centimeters, 

for ease of comparison with leveling and CGPS results. Even if 

our uncertainty estimate is too small by a factor of two, which 

seems unlikely, it does not invalidate the InSAR results reported 

here because most of the displacements are larger than 20 mm 

and corroborated by leveling results.
58 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in 

Yellowstone National Park

This section is intended as an enduring repository for information concerning more than 500 benchmarks in Yellowstone 

National Park that have been used primarily for leveling surveys 

from 1923 to 2007 and also used for more recent gravity and 

GPS surveys. In compiling this information, we included many 

more benchmarks than are available in the National Geodetic 

Survey (NGS) online database (http://www.ngs.noaa.gov/cgibin/datasheet.prl), including marks than have not been used for 

any NGS- or USGS-led surveys (for example, a large number of 

Federal Highway Administration (FHWA) marks along paved 

roads). Also included are many benchmarks, most of them away 

from roads, that have been used by the USGS and University 

of Utah for GPS surveys that are not discussed in this report. 

Nonetheless, we are aware that this compilation is not comprehensive for all benchmarks in Yellowstone National Park. 

Rather, it includes all marks in close proximity to roads that 

have been used for leveling or gravity surveys (or could be used 

in future surveys) plus a selection of backcountry marks that we 

were able to visit on foot. We hope that the availability of these 

descriptions will facilitate the design and execution of future 

geophysical surveys, especially using gravity and GPS.

We searched for benchmarks along Yellowstone National 

Park roads during the summers of 2002, 2003, 2004, and 

2008. Information included in these descriptions is current as 

of October 2008. More than 300 benchmarks were recovered 

in good condition. A significant number (approximately 40) 

that were searched for and not found probably still exist but 

are obscured by vegetation, soil, or other debris. Another 

approximately 30 marks that were not found are presumed 

destroyed, based on evidence for roadwork that probably disturbed the mark (for example, removal of a culvert headwall 

into which a mark was set). Only a handful of benchmarks 

(about 5) could be confirmed destroyed based on evidence 

including broken concrete piers, empty drill-holes, or the 

marks themselves lying on the ground. At the time of this 

writing, a segment of the highway between Madison Junction and Norris Junction was under construction. Realignment will make several benchmarks difficult to access and 

might result in the destruction of at least one mark in the 

headwall of a bridge. Similar work has occurred previously 

elsewhere in the park, where some marks can still be found 

along abandoned road segments that have been remediated 

to their original state. The National Park Service (NPS) and 

FHWA have demonstrated commendable awareness of the 

historical importance of benchmarks throughout the park and 

try to avoid disturbing them or, in cases when disturbance is 

unavoidable, to reset replacement marks nearby. Such was the 

case near Mount Washburn, where several marks were reset 

in 2002 before extensive road construction between Canyon 

Junction and Tower-Roosevelt Junction.

The locations of all benchmarks included in this compilation are shown on the maps in figures B1–B15, which 

are color coded according to the status of the marks (found; 

searched for, not found; presumed destroyed; destroyed; not 

searched for). Interactive versions of the maps are included 

in the data files that accompany this report, together with 

descriptive information for each of the marks. The same 

information is available online at http://pubs.usgs.gov/

pp/1788/pp1788_benchmarks/.

In addition to the interactive maps of figures B1–B15, 

the following information is included in the data files for this 

appendix. Each benchmark description includes a standard set 

of information. The benchmark name is assigned by the organization that installed the mark, generally NGS, USGS, FHWA, or 

NPS. The “STAMPING” field gives the exact stamping that is 

on the face of the mark, and the “USGS QUAD MAP” indicates the USGS 7.5-minute quadrangle map name that covers 

the location of the mark. “NGS PID” refers to the Permanent 

Identifier (PID) of the mark if it is included in the NGS database. Using this unique identifier, one can easily locate the mark 

in the online NGS database (http://www.ngs.noaa.gov/cgi-bin/

datasheet.prl). If this field is empty, then the benchmark is not 

included in the NGS database. The “LATITUDE,” “LONGITUDE,” and “ELEVATION” fields show the most precise 

WGS84 coordinates available for the mark. The locations were 

established by (1) rapid-static GPS measurements by E. Endo 

of the USGS Cascades Volcano Observatory (CVO) in 1996, 

(2) differential GPS measurements during 2002, 2003, and 2004 

using equipment owned by NPS, or (3) handheld GPS measurements in 2008. Coordinates are accurate to within 10 m but are 

much better in most cases.

Below the fields described above is the benchmark 

description (“DESCRIPTION”) and the year in which the 

mark was located or searched for most recently (“DATE OF 

DESCRIPTION”). The text of the descriptions and history 

of the marks are drawn from the online NGS database, from 

USGS descriptions by Dzurisin and Yamashita (1986), Dzurisin and others (1986), Stoopes and Yamashita (1988), and 

Yamashita and others (1997), or from information provided 

by Jeff Place (written commun., 2002) of the FHWA Western 

Federal Lands Highway Division in Vancouver, Washington. 

Each description has been modified to account for differences 

between the original description and current road configuration, mileages, distances to landmarks, or other details as 

determined during reconnaissance work in 2002, 2003, 2004, 

and 2008. In addition to the text, photographs of each mark, 

usually including one closeup and two oblique views from 

opposing vantage points, are provided to aid with locating 

and identifying the mark. For additional information, please 

contact Dan Dzurisin (dzurisin@usgs.gov).
Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone National Park 59

A

A

A

A

1

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

111° 110.5° 110°

45°

45.5°

45°

45.5°

MADISON

JUNCTION

NORRIS

JUNCTION

MAMMOTH

JUNCTION

TOWER

JUNCTION

NORTHEAST

ENTRANCE

WEST

ENTRANCE

CANYON

JUNCTION

FISHING BRIDGE

JUNCTION

LAKE

EAST

ENTRANCE

WEST THUMB

JUNCTION

OLD FAITHFUL

JUNCTION

SOUTH ENTRANCE

111° 110.5° 110°

Yellowstone National Park boundary

Yellowstone National Park boundary

Yellowstone National Park boundary

N

NORTH

ENTRANCE

MARY MOUNTAIN

TRAIL

Figure B2

Figure B3

Figure B4

Figure B5

Figure B6

Figure B7

Figure B8

Figure B12

Figure B11

Figure B10

Figure B9

Figure B15

Figure B14

Figure B13

EXPLANATION

0 10 20 km

0 5 10 miles

Figure B1. Location and status of benchmarks in Yellowstone National Park and vicinity. Benchmarks that were found in good 

condition during reconnaissance work in 2002–8 are green. Several of the marks shown as “searched for, not found” (yellow) 

probably exist but might be buried or obscured; they likely could be found with a concerted effort (this includes several marks along 

the Mary Mountain Trail in areas not likely to have been disturbed since the marks were set). Marks shown as “presumed destroyed” 

(red) were set in areas known to have been disturbed by subsequent roadwork and are unlikely to have survived. Benchmarks that 

could be confirmed destroyed (either because the benchmark itself was found out of place or because the mounting itself, usually

either a drill hole or the base of a concrete pier, was found but the benchmark was absent) are shown in white. Time constraints or 

lack of an adequate description precluded searches for a few marks shown as “not searched for” (blue). An interactive version of 

this map is included on the data files that accompany this report, together with descriptive information for each of the benchmarks. 

The information is available online at http://pubs.usgs.gov/pp/1788/pp1788_benchmarks/. Figures B2–B15 show enlarged sections of 

this map that can be used to locate benchmark descriptions in the data files or online. The one-page descriptions are arranged by 

benchmark name in numerical order, followed by alphabetical order.
60 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

1 MDC 1975

10-7 FHA

A

D339

Q158

A14 RESET

B14

B144

B339

C14

C144

C339

C566

D566

E130

HARLEQUIN

M33

HULL

O33

R158

S158

T158

U158

V158

W162

WE-2 FHA

WE-5 FHA

Y13

Z13

6802.32 BPR

WE-1 FHA

WE-3 FHA

WE-4 FHA

AIRPORT 0 1 32 4 5 km

0 1 2 3 mi

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

Yellowstone National

Park Boundary

Figure B2. Location and status of benchmarks between the West Entrance of Yellowstone National Park and 

Madison Junction. Status is indicated by color code shown in the explanation. See appendix B text and figure B1 

caption for additional information.

1 MDC 1975

1 MDC 1976

10-10 FHA

10-11 FHA

10-12 FHA

10-7 FHA

A

A158

D158

D339

F365

G158

G339

H339

H9

J158

J339

K158

K339

L158

Q158

O9

R158

D365

6802.32 BPR

F158

H158

M158

Q9

P158

N158

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

PIN E

PIN D

STEPHY

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B3. Location and status of benchmarks between Madison Junction and Norris Junction. Status is indicated 

by color code shown in the explanation. See appendix B text and figure B1 caption for additional information.
Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone National Park 61

1 MDC 1976

10-12 FHA

10-14 FHA

10-15 FHA

10-16 FHA

10-17 FHA

19 DI

3 MDC 1978

4 MDC 1975

A158

A339

C157

C9

D157

D9

E365

F157

F9

G157

G365

H157

H365

H9

J157

K157

L157

M157

N157

P157

Q157

R157

SLF6 

T157

U157

V157

X157

Y157

Y8

Z157 (aka 10-13 FHA) D365

S157

W157

11-1 FHA N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B4. Location and status of benchmarks between Norris Junction 

and Mammoth Hot Springs Junction. Status is indicated by color code 

shown in the explanation. See appendix B text and figure B1 caption for 

additional information.
10-19 FHA

12

17 MDC 

19 DI

4 RDS

5 RDS

6344.60

6558.70

6964.69 C157

D157

G365

H365

J365

K365

L365

M365

N365

Q365

R365

RM2

T365

U365

W365

V365

X11

X365

Y11

Y8

Z11

6603.27 BPRA366

S365

1RDS

2RDS

P365

Z365

Y365

11-1 FHA

10-18 FHA

10-20 FHA

10-21 FHA

10-22 FHA

N

Fo

PrSeDeNo

0 1 32 4 5 km

0 1 2 3 mi

Figure B5. Location and status of benchmarks between the North Entrance of Yellowstone National Park and Tower-Roosevelt Junction by way of Junction. Status is indicated by color code shown in the explanation. See appendix B text and figure B1 caption for additional information.
Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone National Park 63

1

10-24 FHA

A

11 MDC 1976

12 MDC 1976

12-1 FHA

12-2 FHA

13 MDC 1976

14 MDC 1976

15 MDC 1976

16 MDC 1976

16 MDC 1977

17 MDC 1976

17 MDC 1977

18 MDC 1976 RESET

19 MDC 1976

21 MDC1976

B366

CVO85-230 D367

DWU

E366

F366

G366

J366

J368

K11

L366

M366

N11

N366

P366

Q11

Q366 RESET

R11

R366

RM2

S366

T366

VS41

X365

20MDC 1976

P11

C366

D366

L11

H366

K366

10-23 FHA

10-25 FHA

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B6. Location and status of benchmarks between Tower-Roosevelt 

Junction and Canyon Junction. Status is indicated by color code shown in the 

explanation. See appendix B text and figure B1 caption for additional information.
64 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

1 MDC 1976

10 MDC 1976

10-12 FHA

10-26 FHA

A

11 MDC

1976

2 MDC 1976

3 MDC 1976

4 MDC 1976

5 MDC 1976

6 MDC 1976

7 MDC 1976

8 MDC 1976

9 MDC 1976

A158

A367

B367

C367

CVO85- D367

230

F365

H9

K11

K339

LC38

NORCAN-1

T9

V9

D365

E367

N Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for 

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B7. Location and status of benchmarks between Canyon Junction and Norris Junction. Status is indicated by color code shown 

in the explanation. See appendix B text and figure B1 caption for additional information.

10-26 FHA

A

10-27 FHA

A

10-28 FHA

10-29 FHA

10-30 FHA

11 MDC 1976

22 MDC 1976

23 MDC 1976

24 MDC 1976

25 MDC 1976

26 MDC 1976

27 MDC 1976

28 MDC 1976 29 MDC 1976

30 MDC 1976 31 MDC 1976

7743.43

A11

ARBEE / CVO 84-22

B11

CVO 85-229 / 84-2

CVO 85-228 / 84-3 CVO 85-227 / 84-6

CVO 85-220 / 84-11

CVO 85-221 / 84-10

CVO 85-222 / CVO 84-9

CVO 85-224 / 84-7

CVO 85-225 / 84-5

CVO85-226 / 84-4

CVO85-230 / CVO 84-1 D367

DA3

E11A2

F11A

HOLLIS

K11

 KAYGEE

L19

LC58

LE HARDY

U367

V367

X367

Y367

Z367

6RDS

G11

W367

CVO 84-8 / CVO 85-223

J11

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

0 1 432 5 km

0 1 2 3 mi

Figure B8. Location and status of benchmarks 

between Canyon Junction and Fishing Bridge 

Junction. Status is indicated by color code shown 

in the explanation. See appendix B text and figure 

B1 caption for additional information.
Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone National Park 65

10-30 FHA

10-32 FHA

10-33 FHA

2 MDC 1977

3 MDC 1977

31 MDC

1976

4 MDC 1977

5 MDC 1977

POT /

54 MDC 1976

6 MDC 1977

7 MDC 1977

8 MDC 1977

9 MDC 1977

A11

A368

B368

BIG THUMB

C368

D368

E368

F368

G368

H368

L19

PIT

R10A

S10A

T10

T367

U367

V10

W10

WT2

U10 10-31 FHA

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

LAKE ASTRO

YELLOWSTONE

L13

Z10 RESET

0 1 32 4 5 km

0 1 2 3 mi

Figure B9. Location and status of 

benchmarks between Fishing Bridge 

Junction and West Thumb Junction. Status 

is indicated by color code shown in the 

explanation. See appendix B text and 

figure B1 caption for additional information.

10-1 FHA

10-2 FHA / OF4

10-33 FHA 10-34 FHA

48 MDC 1976

49 MDC 1976

S367 / 53 MDC 1976

POT / 54 MDC 1976

7273.15 BPR

8 MDC 1977

A

CVO 86-17

D10

F10

F367

G367

H10

J367

L10 RESET

L367 M367

MEERTENS

N367

SHOSHONE

T367

WT2

Y366

Z366

Q367 M10 R367

N10

O10

P10

50MDC 1976

51MDC 1976

H367

K367

10-35 FHA / 52 MDC 1976

K 10 RESET

P367

CVO 86-18

CVO 86-19

CVO 86-20

CVO 86-21

CVO 86-22

CVO 86-23

10-34 AZI FHA

0 1 32 4 5 km

0 1 2 3 mi

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

MEERTENS2

Figure B10. Location and status of benchmarks between West Thumb Junction and Old Faithful Junction. Status is indicated by color 

code shown in the explanation. See appendix B text and figure B1 caption for additional information.
66 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

1 MDC

1975

10-1 FHA

10-2 FHA / OF4

10-4 FHA

10-6 FHA

10-7 FHA

A

41 MDC 1976

42 MDC 1976

43 MDC 1976

45 MDC 1976

48 MDC 1976

7273.15 BPR

A

C10

CVO 94-11

D10

D339

E339 A

MEERTENS

NEZ PERCE

Q158

THERMAL

Y366

Y9

15MDC 1977

R158

40MDC 1976

44MDC 1976

F339

U366

V366

W366

X366

Z9

6802.32 BPR

46MDC 1976

47MDC 1976

DA1

DA3

10-3 FHA

10-5 FHA

0 1 32 4 5 km

0 1 2 3 mi

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

MEERTENS2

Figure B11. Location and status of benchmarks between Old 

Faithful Junction and Madison Junction. Status is indicated by 

color code shown in the explanation. See appendix B text and 

figure B1 caption for additional information.

10-6 FHA

E11A2

F11A

NEZ PERCE

10RDS

11RDS

13RDS

12RDS

14MDC 1977

14RDS

15MDC 1977

6RDS

7RDS

9RDS

8RDS

A365

B365

C365

CHISELED CROSS

CHISELED SQUARE

P368

Q368

R368

S368

SPIKE1

SPIKE2

SPIKE3

SPIKE4

SPIKE5

T368

TROUT

U368

V368

V366

W367

DA1

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B12. Location and status of benchmarks along the Mary Mountain Trail. Status is indicated by color code shown in the 

explanation. See appendix B text and figure B1 caption for additional information.
Appendix B. Descriptions, Coordinates, and Photographs of Benchmarks in Yellowstone National Park 67

1

1 GWM

12-1 FHA

12-2 FHA

12-3 FHA

12-4 FHA

12-5 FHA

12-6 FHA

12-7 FHA

12-8 FHA

16 MDC 1976

17 MDC 1977

18 MDC 1977

2 GWM

3 GWM

4 GWM

5 GWM

6 GWM

6449.66

6552.22

6570.59

6580.56

6619.26

6626.42

6652.02

6684.92

6738.91

6830.02

6835.25

6902.10

6980.85

7008.17

7123.74

7194.05

7171.69

B366

DWU

J368

K368

L368

M368

N368 RM2

VS39

VS40

VS41

X365

X368

Y368

Z368

W368

Y365

6183.85

6188.52

0 1 32 4 5 km

0 1 2 3 mi

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

Yellowstone National Park Boundary

Figure B13. Location and status of benchmarks between the Northeast Entrance of Yellowstone National Park and Tower-Roosevelt 

Junction. Status is indicated by color code shown in the explanation. See appendix B text and figure B1 caption for additional information.

10-30

FHA

13-1 FHA

13-4 FHA/13-4 AZI FHA

31 MDC 1976

32 MDC 1976

38 MDC 1976

7743.43

A11

CVO 84-15

CVO 84-12 / 84-27

CVO 84-16R

CVO 84-19R

CVO 84-24

CVO 93-004

H12

K12

L19

LAKE ASTRO

M369

U367

YELLOWSTONE

Z10 RESET

33 MDC

34 MDC

CVO 84-20

CVO 84-26

D12

E12

F12

13-3 FHA

36MDC 1976

37MDC 1976

39MDC 1976

L369

CVO84-13

CVO84-14

CVO 84-17

CVO84-18

CVO 84-25

CVO 84-28

CVO 84-21

13-5 FHA 13-2 FHA

CVO 93-008

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

Yellowstone National

Park Boundary

0 1 32 4 5 km

0 1 2 3 mi

Figure B14. Location and status of benchmarks between the East Entrance of Yellowstone National Park and Fishing Bridge Junction. 

Status is indicated by color code shown in the explanation. See appendix B text and figure B1 caption for additional information.
68 History of Surface Displacements at the Yellowstone Caldera, Wyoming, from Leveling Surveys and InSAR Observations, 1923–2008

10 MDC 1977

10-33 FHA

11 MDC 1977

12 MDC 1977

13 MDC 1977

14-1 FHA

14-2 FHA

A

14-4 FHA

19 MDC 1977

2 BPR

50 RDS

S367 / 53 MDC 1976

POT / 54 MDC 1976

8 MDC 1977

9 MDC 1977

A369

B369

BIG THUMB

C369 CVO 86-10

CVO 86-11

CVO 86-12

CVO 86-13

CVO 86-15

CVO 86-16

CVO 86-17

D369

E369

F369

FLAGG

G368

GS1

GV2

H368

J369

K369

L13

LEWIS FALLS

N13

N369

O13

P369

Q369

S13

T13

T367

WT2

G13

H13

J13

K13

14-3 FHA

CVO 86-14

Yellowstone National Park Boundary

N

Found

Presumed destroyed

Searched for, not found

Destroyed

Not searched for

EXPLANATION

0 1 32 4 5 km

0 1 2 3 mi

Figure B15. Location and status of benchmarks 

between the South Entrance of Yellowstone National 

Park and West Thumb Junction. Status is indicated by 

color code shown in the explanation. See appendix B 

text and figure B1 caption for additional information.
Produced in the Menlo Park Publishing Service Center, California

Manuscript approved for publication December 28, 2011

 Text edited by James W. Hendley II

Layout and design by Judy Weathers
Dzurisin and others—History of Surface Displacements at the Yellowstone Caldera from Leveling Surveys and InSAR Observations—Professional Paper 1788