Yellowstone Plume Head: Postulated Tectonic Relations to the 

Vancouver Slab, Continental Boundaries, and Climate 

by Kenneth L. Pierce,1 Lisa A. Morgan,2 and Richard W. Saltus2 

Open-File Report 00-498 

Online Edition 

2000 

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

(or with the North American Stratigraphic Code). Any use of trade, product, or firm names is for descriptive 

purposes only and does not imply endorsement by the U.S. Government. 

U.S. DEPARTMENT OF THE INTERIOR 

U.S. GEOLOGICAL SURVEY 

1

 Bozeman, Montana 

2

 Denver, Colorado 
ABSTRACT 

We trace the Yellowstone hotspot track back to an apparent inception centered near the OregonNevada border. We and others have concluded this is the locus of a starting plume or plume head. Consideration of this plume-head model leads us to discuss the following three implications. 

(1) The apparent center of the relic plume head is about 250 km west of the location where both the 

trend of the younger hotspot track and the inferred plate motions would place the hotspot at 16 Ma. A 

possible explanation for this discrepancy is the westward deflection of the plume up the inclined 

Vancouver slab. Plate tectonic reconstructions and an intermediate dip for the Vancouver slab indicate a 

plume head would have encountered the Vancouver slab. 

(2) The postulated arrival of the plume head at the base of the lithosphere is temporally associated 

with eruption of the Columbia River and Oregon Plateau flood basalts at 14-17 Ma; however, these 

basalts were erupted several hundred kilometers north of the apparent plume center. The postulated plume 

center is symmetrically located near the midpoint of the 1,100-km-long Nevada-Oregon rift zone (see fig. 

1). Strontium isotopic variations reflect crustal and mantle lithosphere variations along the trend of this 

rift zone, with the basalt area of Oregon and Washington lying west of the 0.704 line in oceanic crust, the 

apparent center in northern Nevada between the 0.704 and 0.706 line in intermediate crust, and the area of 

central and southern Nevada east of the 0.706 line in Precambrian continental crust. Geophysical modeling is consistent with a dense crust north of the Nevada-Oregon border and an asthenospheric low-density 

body that extends several hundred kilometers south and north of the Nevada-Oregon boundary. A reconstruction of the initial contact of the plume head with the lithosphere suggests relatively thin lithosphere at 

17 Ma beneath Oregon and Washington, which would favor the spreading of the plume northward in this 

direction, more decompression melting in this "thinspot" area, and the eruption of basalt through dense, 

oceanic lithosphere. Thus, preferential extrusion of flood basalts north of the plume center may be the 

result of differences in the pre-plume lithosphere, and not the location of the center of the plume head. 

(3) A plume head rising into the base of the lithosphere is expected to produce uplift, which we 

estimate to be on the order of 1 km with a north-south dimension of 1,000 km. This plume-head uplift, 

followed by subsidence, is consistent with Cenozoic paleobotanical altitude estimates. Other climatic 

indicators show major aridity about 15 Ma in areas in the inferred precipitation shadow east of the inferred uplift. Indicators of climate about 7 Ma are compatible with an eastward migration of uplift to a 

site between the plume-head area and the present Yellowstone crescent of high terrain. The warm Neogene "climate optimum" correlates with 14-17-Ma flood basalt and rhyolite volcanism. The continued 

effects of Yellowstone plume-head uplift and ensuing plume tail uplift, if real, could provide regional 

uplift that is geophysically plausible. Climatic modeling has shown that uplift of the age and latitude of 

the postulated Yellowstone plume-head uplift, if allied with Himalayan and perhaps other uplifts could 

result in the late Cenozoic cooling leading to the Pliocene-Pleistocene ice ages (Kutzbach and others, 

1989; Ruddiman and others, 1989, 1997). 

Thus, the postulated Yellowstone plume head could have played an important role in the late Cenozoic geologic history of the northern, interior part of the U.S. Cordillera. Future studies of the kind briefly 

discussed here should provide a better evaluation of the Yellowstone plume head concept. 

INTRODUCTION 

A common explanation of hotspot tracks, such as the one responsible for the Hawaiian Islands, is 

that they are generated by the interaction of a relatively fixed deep thermal plume with a moving lithospheric plate. Richards and others (1989) suggested that mantle plumes initiate as larger plume heads that 

produce associated flood basalts and then evolve into narrower plume tails associated with later, smaller 

volume volcanism. Thermal mantle plumes are controversial in part because they cannot be directly 

2 
investigated; the narrow plume tails are too small in cross section to be imaged seismically. Mantle 

plumes are thought by many, however, to explain an important and still poorly resolved piece of the 

generally accepted plate-tectonic model. A starting plume head rises slowly from deep in the mantle and 

is fed from below by a much thinner tail or chimney in which heated material is rising through the "tail 

pipe" an order of magnitude faster than the plume head rises. Thus, as the plume head rises, it inflates to 

diameters of hundreds or perhaps more than a thousand kilometers before it slowly impacts the base of 

the lithosphere. Upon rising the last 150 km or so, decompression melting in the plume head results in the 

eruption of flood basalts. In his presidential address to the Geological Society of America, George 

Thompson (1998) argued for the importance of deep mantle plumes in general and the Yellowstone plume 

head and tail in particular, concluding that The paradigm of deep mantle plumes, like plate tectonics or 

asteroid impacts, supplies a wonderful unifying concept for geoscientists and for communicating our 

science to the world at large. 

Plume heads produce a large area of uplift that starts before their actual arrival at the base of the 

lithosphere (Hill and others, 1992). Hotspot swells with heights of 1-2 km and diameters of 800-1,200 km 

(Crough, 1978) are associated with current hotspot positions commonly inferred to be plume tails. The 

partial melting of upwelling hot mantle produces two materials that are both lighter than the original 

mantle (density about 3.3 g/cc): basalt melt (density about 3.0 g/cc) that rises upward and restite residuum 

(density perhaps about 3.0 g/cc) that stays in the mantle (Morgan and others, 1995). 

Despite their apparent unifying appeal, many dispute the mantle plume idea. For example, in the 

Pacific plate near the Hawaiian, Society and Marquesas hotspots, Katzman and others (1998) determined 

that upper mantle velocities were high and not low as predicted by traditional hotspot thermal models. 

They suggest Richter-type convective rolls above the 660-km discontinuity and oriented parallel to plate 

velocity rather than a deep thermal plume. Don Anderson (1998 and references therein) argues against the 

mantle-plume explanation of hot spots and particularly the deep-seated mantle plume idea and presents 

alternative explanations to plume arguments. 

In contrast, the recent book "The Earth's Mantle" contains many chapters by different authors 

support the concept of deep mantle plumes, and the concept of starting plumes or plume heads. Griffiths 

and Turner (1998) present argument for the concept of a starting plume head, fed by a plume tail. Davies 

(1998) describes mantle convection with descending slabs and rising plumes. Based on chemical considerations, Campbell (1998) argues for a deep plume source for ocean island basalts. Based on seismological and experimental studies, Jackson and Rigden (1998) argue that the mantle is grossly uniform in 

chemical composition throughout and that phase transformations provide an adequate explanation for the 

seismically observed radial structure. In another book, Garnero and others (1998), recognize an ultralow 

velocity zone (ULVZ) that may represent partial melting at the core-mantle boundary. Correlation of this 

ULVZ with surface hotspot distribution and anti-correlation with descended slabs suggests whole mantle 

convection may intersect just above the core-mantle boundary. 

As a test of the viability of the Yellowstone plume hypothesis, we were led into the problem of 

addressing the initial stages of the Yellowstone plume head by tracking the Yellowstone hotspot back to 

the west from its present location at the Yellowstone Plateau (Pierce and Morgan, 1990, 1992; fig. 1, 

hotspot track; Anders and others, 1989; Rodgers and others, 1990; Malde, 1991; Smith and Braile, 1993). 

The hotspot track is associated with a northeastward progression of rhyolitic volcanic fields, faulting, and 

uplift. The segment containing the 10-Ma-and-younger rhyolitic volcanic progression is linear in both rate 

and trend and is reflected in the topographic depression of the eastern Snake River Plain (SRP). In 

contrast, the segment containing the 10-Ma-and-older rhyolitic progression includes two calderas active 

about 10 Ma but spaced 100-200 km apart (Morgan and others, 1997). The progression associated with 

rhyolitic volcanism becomes increasingly diffuse to the southwest. The earliest stages of this Yellowstone 

style volcanism can be projected back in space and time along this trend to the McDermitt caldera com3

plex along the Nevada-Oregon boundary about 16 Ma (Malde, 1991; Thompson and Gibson, 1991; Pierce 

and Morgan, 1992) and to the northern part of the northern Nevada rift (Zoback and Thompson, 1978). As 

pointed out by Pierce and Morgan (1992), the 10-Ma and younger trend has an orientation of N. 54o E. 

and an apparent rate of 2.9 centimeters/yr while the 10-Ma and older trend has an orientation of N. 75o E. 

and an apparent rate of 7.0 centimeters/yr. These differences in apparent orientation and rate compound 

the problem of clearly identifying the initial hotspot starting location. 

Based largely, however, on the track of the Yellowstone hotspot and its Columbia River-Oregon 

Plateau flood basalt association, the initial stages of the Yellowstone starting plume (head) have been 

inferred to begin at about 17 Ma beneath the northern Basin and Range (Zoback and Thompson, 1978; 

Hooper, 1990; Duncan and Richards, 1991; Thompson and Gibson, 1991; Draper, 1991; Zoback and 

others, 1994; Camp, 1995; Parsons, 1995). The large sublithospheric density deficit required to support 

high topography at this location is interpreted by Parsons and others (1994) and Saltus and Thompson 

(1995) to be the remnants of the Yellowstone plume head. 

If we assume the nascent Yellowstone starting plume intercepted the lithosphere centered near the 

Nevada-Oregon border about 17 Ma, then the following possible relationships need to be examined and 

have implications on understanding the region: (1) The rising plume head may have interacted with and 

been diverted westward by the inclined Vancouver slab of oceanic lithosphere. This interaction may 

explain the anomalously high apparent volcanic migration "rate" from 16 to 10 Ma which does not 

correspond with a known change in rate of the North American Plate. (2) Differences in the character of 

the pre-plume lithosphere along a NNW-SSE 1,000-km-long trend centered near McDermitt may account 

for the different geologic features observed, particularly the 14.5- to 17.5-Ma major pulse of Columbia 

River-Oregon Plateau flood basalt eruptions north of the center of the plume head (assumed to be near the 

McDermitt caldera complex) and the dike injection along the 500-kilometer-long northern Nevada rift 

and associated 14- to 17-Ma basaltic and silicic volcanism near and south of this center. (3) Significant 

topographic domal uplift above the plume head may have had great impact on past global and northwest 

U.S. climates and may explain regional changes in topography since the middle Miocene. If geologic 

observations support these three relations, this may provide support to the plume-head hypothesis. In 

addition to the three relations described above, the current high regional heat flow (Blackwell, 1989; 

Lachenbruch and Sass, 1978) and ongoing extension of the Basin and Range region of the North American plate may also relate to the stagnant Yellowstone plume head. 

WESTWARD DISPLACEMENT OF PLUME HEAD BY INCLINED VANCOUVER SLAB 

Zoback and Thompson (1978) first suggested that the Yellowstone hotspot surfaced in northern 

Nevada about 16 Ma. Malde (1991) noted that the oldest set of calderas along the Yellowstone hot-spot 

track erupted the 16.1-Ma rhyolites of the McDermitt field (Rytuba and McKee, 1984, fig. 1). Several 

others have also postulated the McDermitt area near the Nevada-Oregon border (fig. 1) to be the approximate location of a large starting-plume (Yellowstone plume head, Thompson and Gibson, 1991; Draper, 

1991; Pierce and Morgan, 1992; Parsons and others, 1994; Zoback and others, 1994; Camp, 1995). 

Placing the starting position in the McDermitt area raises some problems, as noted but not explained 

by Pierce and Morgan (1992). The apparent rate of hotspot migration from 16 to 10 Ma is 7 centimeters/ 

yr, more than twice the 10-Ma to present 2.9 centimeters/yr apparent rate (Pierce and Morgan, 1992). On 

the basis of intervals of very high fault offset, Anders and others (1994) determined a plate migration rate, 

accounting for extension, of 22 km/m.y. (km/m.y. = mm/yr), quite compatible with a global hotspot (but 

excluding Yellowstone) plate tectonic rate of 2.2±0.8 km/m.y. (Alice Gripps, written commun., 1991, in 

Pierce and Morgan, 1992). 

Projection of the trend of the 2- to 10-Ma hotspot track back to 16 Ma yields a location near the 

4

125° 120° 115° 110°


45° 

40° 

35° 

X 

X X 

X 

6.6 

Ma 

10 

Ma 

2 

Ma 

➤

➤ 

A' 

A 

~260 km 

offset 

ZONE 

704 

706 

704 

706 

X X X 

X X 

X 

X 

➤ ➤ ➤ ➤ ➤ 

RIFT 

➤ CALIFORNIA

OREGON 

NEV AD A 

UTAH

COLORADO 

IDAHO 

WYOMING 

MONTANA 

WASHINGTON 

OREGON 

Active

Basin 

a

nd 

Projected position 

of plume @ 16 Ma, 

S54°W@2.9 cm/yr 

from 10 Ma field 

McDermitt Calderas 

~Plume head center @ 16 Ma 

NEV AD A 

Flood basalt, mostly 14-17 Ma, Columbia 

River and Oregon Plateau 

87Sr/86Sr 

Northern Nevada rift and similar features 

to west. Southern extension dashed 

Feeder dikes, flood basalts 

Area of rhyolite eruptions, 14-17 Ma 

Volcanic fields and track of Yellowstone 

hotspot, 10 Ma and younger 

X - ignimbrite ages, - rhyolite ages 

EXPLANATION 

0 500 

Kilometers 

Scale 

6.6 

Ma 

Act i ve Basi n an d R ang e 

Range 

0.704 and 0.706 lines 

Figure 1. Map of western United States showing the track of the Yellowstone hotspot (after Pierce and Morgan, 1992, fig. 23). The 10- to-2-Ma track 

is compatible with plate motion, whereas the 10- to-16-Ma track has apparent rates three times faster than predicted. Projecting back from 10 Ma, we 

infer the plume was located at the site of the hexagon. As shown by the arrowed line, the geologic location of the inferred plume head was about 260 

km further west at about the McDermitt caldera. Assuming the Vancouver slab was in the subsurface (figs. 2 and 3), buoyant rise up the slab may 

have deflected the plume head westward the 260 km shown by the arrowed line. Strontium lines after Reed (1993) with modifications in Washington 

from Robert Fleck (written commun., 1998) and southeastern Idaho from Leeman and others (1992). Transect A-A’ is portrayed in Figure 5. 

5

Idaho-Utah border about 250 km east of the apparent starting plume centered near McDermitt. West of the 

10-Ma Picabo volcanic field, successively older fields have a very crude trend that ends up at McDermitt, 

but this direction of hotspot migration from 16 to 10 Ma is N. 75° E., about 20° different from the direction of N. 54° E. from 10 to 2 Ma. 

We have not accounted for the amount of extension (Rodgers and others, 1990) after 16 Ma over the 

present distance of 430 km between the McDermitt area and the 10-Ma Picabo volcanic field, other than 

to assume the observed post-10-Ma rate of 29 km/m.y. is the vector sum of the plate rate and the extension rate. If extension has been constant after 16 Ma, this is only about 25% of the total calculated rate of 

"migration" based on Anders (1994) estimation of 22 km/m.y. for plate motion combined with 7 km/m.y. 

of tectonic extension to result in a combined rate of 29 km/m.y. One observation that suggests extension 

has been minor is the smooth, relatively unfaulted topography between McDermitt and the central SRP as 

shown on the shaded relief map of the U.S. (Thelin and Pike, 1991). This broad plateau is formed largely 

of 10- to 16-Ma volcanic strata (Luedke and Smith, 1981, 1982, 1983) suggesting little faulting since that 

time. 

Other evidence based on fission track studies in the eastern Basin and Range province, however, 

suggests major extension occurred between 20 and 15 Ma (Elizabeth Miller, written commun., 1999). 

Local extension east of McDermitt is indicated by opening of the Western Snake River Plain (a few 10's 

of km), and by opening of the Raft Valley by the eastward sliding of the Black Pine and Sublette Ranges 

off the Albion Range (Covington, 1983, also shown in Pierce and Morgan, 1992, fig. 11). If extension east 

of McDermitt is as large as 260 km, this could then explain the 260 km westward offset of the 16-Maplume location (fig. 1) in the last 16 Ma, and following hypothesis of westward deflection by the 

Vancouver slab would not be needed. 

RECONCILIATION BASED ON WESTWARD DEFLECTION BY THE VANCOUVER SLAB 

We suggest here that the apparent discrepancy between both rate and azimuth of hotspot migration 

can be explained by the westward displacement of the plume head when it rose up into the east-dipping 

Vancouver (or Juan de Fuca) slab (figs. 2, 3). Geist and Richards (1993) suggested that the Yellowstone 

plume intercepted the Vancouver slab before 17 Ma, but they argued that the downward and strong 

northeast motion of the Vancouver plate trapped the plume for some time and carried it northward to near 

the common borders of Washington, Oregon, and Idaho where it eventually broke through the slab and 

produced the Columbia River flood basalts (fig. 1). We modify their idea of interaction with the 

Vancouver slab by suggesting that buoyant gravitational forces would be more effective than tractive 

forces and that the plume would buoyantly rise up along the lower surface of the inclined slab, rather than 

being dragged northward and held down by the northeastward-moving and descending Vancouver slab. 

For a plume beneath a horizontal plate, tractive forces, where not opposed to gravity, do appear significant. We also propose a starting plume (or plume head) interacting with the Vancouver slab rather than an 

already existing plume tail (or chimney) that previously had been located beneath the Pacific Ocean as 

proposed by others (Geist and Richards, 1993; Duncan, 1982). The possible interaction of the Vancouver 

slab with the Yellowstone plume has been proposed by several others (Leeman, 1982; Duncan, 1982; 

Draper, 1991; Hill and others, 1992; and Hooper and Hawksworth, 1993). 

To reconstruct the interaction of the Vancouver slab with the Yellowstone hotspot, we have used the 

plate tectonic reconstructions of Severinghaus and Atwater (1990; fig. 2). We have projected the track of 

the Yellowstone hotspot back in time on the basis of its 2- to 10-Ma rate of 2.9 cm/yr at S. 54° W. (includes both plate rate and extension rate, Pierce and Morgan, 1992), using reconstructions for 30, 20, 10, 

0 Ma (Severinghaus and Atwater, 1990). Projecting the direction and rate from 2-10 Ma back to 16 Ma is 

warranted for the "paleomap" models of global plate motion (Malcolm Ross, written commun., 1998, 

6

based on paleomap programs). Such models show the North American Plate having no more than a 4-

degree change in direction and 3 percent variation in rate over the last 20 Ma. However, a 12-degree 

change in direction and a 17 percent increase in rate are indicated about 20-21 Ma, which is before the 

volcanism we associate with the Yellowstone hotspot. 

Figure 3 is a time sequence of cross sections incorporating the above rate with plume rise models 

and plate tectonic reconstructions. For the plume head, we have used a rise rate of 0.1 m/yr (100 km/m.y.) 

suggested by Richards and others (1989) going back in time (and downward) from its interception with 

the North American plate (lithosphere) about 17 Ma ago. We estimate the volume of the plume head in the 

upper mantle to be roughly 400 km3

 based on: (1) the volumes erupted from the Columbia River-Oregon 

Plateau flood basalts (Hooper, 1997; Carlson and Hart, 1988), assuming 5-30 percent of partial melting 

(202-368 km3

; Coffin and Eldholm, 1994), plus (2) the linear extent of erupted and injected material along 

the 500-km Nevada rift zone (Zoback and others, 1994); and (3) the residual mass deficit shown by 

Parsons and others (1994) for the northern Nevada area that is comparable to the expected deficit produced by a small plume head with a diameter of about 400 km and 100-300 degrees C hotter than the 

surrounding asthenosphere (Hill and others, 1992). We assume the chimney (tail) of the Yellowstone 

plume has been held fixed over time (fig. 3) and that the North American plate has migrated at 22 km/m.y. 

and the hotspot has migrated at 29 km/m.y. (plate motion plus extension). As already noted, we assume 

the heated material in the tail is rising at about 1.0 m/yr, which is an order of magnitude faster than the 

rate of the ascending plume head (0.1 m/yr; or 100 km/m.y.). Thus, ascending material in the plume tail 

would contribute to inflation of the more slowly rising plume head. 

The inclination of the Vancouver slab is difficult to constrain. About 50 Ma, it is thought to have 

been flat (see discussion in Atwater, 1989, p. 46-49). At present the dip as far inland as the CascadesColumbia Plateau boundary is 12 degrees (Parsons and others, 1998). Further inland and with more 

difficult techniques, Rasmussen and Humphreys (1988) estimate a dip of about 65 degrees near the 

Washington-Oregon boundary, and VanDecar (1991) estimate about 60 degrees beneath Washington. 

Seismic imaging by van der Lee and Nolet (1997) shows remnants of the Vancouver slab at 500 km 

beneath the area near Salt Lake City, suggesting a time-averaged dip of about 30 degrees; they suggest a 

flat slab extending 1,000 km inland at 50 Ma, and at about 30 degrees dip extending more than 50 km 

inland at 30 Ma. Geist and Richards (1993) draw about a 35 degrees slab dip at 20 Ma. We have drawn a 

dip of about 40 degrees at 30 Ma that steepens to 45 degrees at 16 Ma and 50 degrees at present. 

By about 20 Ma, the starting plume intercepted the Vancouver slab (fig. 3). If gravitational buoyancy 

dominates over tractive forces, the buoyant plume head would be deflected to the west by the west-rising 

lower surface of the Vancouver slab. We suggest the tractive force that the Vancouver slab could exert on 

the relatively low viscosity plume head, particularly by dragging the plume head down, would be minor 

compared to the buoyant rise of an expanding starting plume. Put more simply, how would the hot, low 

viscosity plume head stick to the steeply inclined slab? Also, consideration must be given to the slope 

angle of the descending slab. If the slope is as steep as 60 degrees, as inferred by Rasmussen and 

Humphreys (1988), traction would have even less of an effect on the ascending plume and thus would not 

block it. 

Buoyant rise directly up the slab would displace the plume head to S. 70° W. That the deflection 

appears to be essentially to the west may reflect a small component of northward drag due to the northeasterly movement of the Vancouver slab relative to the North American plate. 

With this westward offset by the Vancouver slab, the apparent rate of migration of volcanism between 16 and 10 Ma is an additive combination of the projected plate motion plus extension of 29 km/ 

m.y. for 6 m.y. equaling 175 km plus the 250 km net offset by the Vancouver plate. This results in an 

apparent migration of approximately 420 km in 6 m.y., solved so as to yield the apparent migration rate of 

70 km/m.y. 

7

120°W 100°W


40°N 

40°N 

40°N 

40°N 

✪ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

▲ 

30 Ma 

20 Ma 

10 Ma 

0 Ma 

Sila F.Z. 

Sedna F.Z. 

Sila F.Z. 

Surveyor F.Z. 

Sedna F.Z. 

Sedna F.Z. 

Mendocino F.Z. 

Pioneer F.Z. 

JUAN 

DE 

FUCA 

PLATE 

10 

10 

3 

2 

1 

3 5 

1 2 

5 

3 

2 

1 

JUAN 

FUCA 

VANCOUVER 

SLAB 

VANCOUVER 

SLAB 

3 5 

2 

1 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

✪ 

✪ 

✪ 

➤ 

➤ 

➤ 

➤ 

LEGEND 

1,000 Km 

Scale @ 43°N 

0 

➤ ✪ 

✪ Inferred past locations based on 

past 10 Ma motion of N54°E 

@ 29km/m.y. 

Track of the Yellowstone hotspot 

✩ 

Present 

DE 

SLAB 

Figure 2. Location of the Yellowstone plume relative to the North American Plate and the Vancouver slab (or Juan de Fuca slab; from 

Severinghaus and Atwater, 1990). The plume is assumed to be fixed in the mantle and now 50-km northeast of the 2.0-Ma caldera that started 

the Yellowstone Plateau volcanic field. Hotspot migration of 29 km/m.y. from 2 to 10 Ma is plotted as a circled star on the line of section. This 

rate of 29 km/m.y. is close to the plate motion rate of 22 ± 8 km/m.y. and includes tectonic extension that probably accounts for the apparent 

difference of 7 km/m.y. Cross sections are drawn to go through the McDermitt caldera and are nearly parallel to plate motion and to the 

inclination of the Vancouver slab. The reconstruction of Severinghaus and Atwater (1990) accounts for deformation, and the numbers on the 

Vancouver slab indicate the thermal state of the slab with 1 meaning solid enough to have earthquakes and 10 meaning nearly the same as the 

surrounding mantle. 

8

➤ ➤ ➤ 

➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ 

➤ 

➤ 

10 Ma 

0 Ma 

16 Ma 

20 Ma 

30 Ma 

WEST EAST 

Plume head 

400 km across 

Plate also moving into section 

Plate also moving 

into section 

stagnant head 

stagnant head 

= restite enriched, 

buoyant 

490 km 

Plate motion 

➤ Plate motion 

➤ Plate motion 

➤ Plate motion 

➤ Plate motion 

chimney YNP 

YNP 

YNP 

YNP 

YNP 

OR/NV rift 

Chimney 

0 

Scale @ 43°N 

1,000 Kilometers 

➤ 

➤ ➤ 

➤ 

400 km ➤ ➤ 

290 

km 

➤ ➤ 

= restite enriched, 

buoyant 

➤ 

➤ 

➤ 

➤ 

➤ 

Figure 3. Postulated westward deflection of the Yellowstone plume head by buoyant rise up the inclined Vancouver slab. Time sequence from 30 Ma 

(top) to present (bottom) with position of plume held fixed in the mantle. At 20 Ma, the plume head has intersected the inclined Vancouver slab and is 

being displaced westward, and by 16 Ma the plume center is beneath McDermitt (along the OR/NV rift), about 260 km west of its feeding plume tail. 

As the plume flattens, the area of the plume in this cross section diminishes to about half because of spreading in the third dimension. The inclination 

of the Vancouver slab is shown to increase from 30 Ma to present (see text under heading "Reconciliation Based..."). Numbers on slab (1, 3, 5, and 10) 

are the thermal state of slab from Severinghaus and Atwater (1990); see also Figure 2. This cartoon does not account for interaction with the 660- and 

410-km discontinuities. 

9

OTHER PROPOSED LOCATIONS OF THE YELLOWSTONE HOTSPOT 

A number of proposals about the location of the Yellowstone plume at different times have been 

advanced (fig. 4). Geist and Richards (1993) suggest the Yellowstone plume was trapped for some time 

and carried northward by the Vancouver slab to a position beneath the main extrusion area of the Columbia River basalts. Our idea builds on the Vancouver slab interaction suggested by Geist and Richards 

(1993). We propose that buoyant forces dominated tractive ones and that the plume head was centered 

near the Nevada-Oregon border. 

In the next section and earlier (Pierce and Morgan, 1992), we argue that the nascent Yellowstone 

plume intercepted the lithosphere near the McDermitt caldera area and that the apparent asymmetry 

related to the rhyolitic track of the Yellowstone hotspot with respect to the location of the Columbia 

River-Oregon Plateau flood basalts is due to changes in pre-plume lithosphere thickness and crustal 

composition. We argue herein and in our prior paper (Pierce and Morgan, 1992) that the asymmetry is 

merely apparent. In fact, the 1,100-km-long Nevada-Oregon rift can be considered an equal, but overlooked, component to the early stages of the plume head intercepting the base of the lithosphere, as are 

the voluminous flood basalt eruptions to the north. Both components are symmetrical to the McDermitt 

area. Specifically, the Columbia River-Oregon Plateau basalts, located primarily north of the McDermitt 

center and north and off the trend of the hotspot track, are the product of a thinner, more mafic, younger 

accreted crust (Hooper, 1997) where the flood basalts surfaced while the northern Nevada rift developed 

in thicker, Proterozoic to Paleozoic, mafic to transitional accreted crust. From 14 to 16 Ma in northern 

Nevada and adjacent parts of Oregon and Idaho, there was widespread silicic volcanism, although the 

locations of the calderas have not been well established (fig. 1; Luedke and Smith, 1981, 1982. 1983). 

Oppliger and others (1997) suggest that the 34-to 43-Ma "Carlin" gold mineralization at Carlin, 

Nevada, is associated with an "incubation period" of an early Yellowstone plume. They cite the rich 

abundance of gold and related siderophile elements in the Carlin deposits as being plume related and note 

that these enriched elements are thought to be derived from the core-mantle boundary where many think 

thermal plumes originate. We suggest, however, that having the Yellowstone plume beneath northern 

Nevada at 34-43 Ma is difficult to reconcile considering the rate and direction of motion of the North 

American plate (fig. 4) and the current location of the Yellowstone plume. Assuming a rate of 29 km/m.y. 

in a direction of S. 55° W. from Yellowstone places a 40-Ma plume somewhere near Sacramento, California, quite distant from Carlin, Nevada. The lack of volcanic deposits or related rifts that may have recorded the passage of a thermal plume in a time-transgressive pattern to the southwest from McDermitt is 

striking. Furthermore, given our conclusion that the Yellowstone plume head intercepted the lithosphere at 

about 17 Ma near McDermitt, we cannot reconcile the proposal by Oppliger and others (1997) that the 

plume head was southeast of McDermitt and only 380 km from the 10 Ma position of the hotspot. Such a 

location would yield an anomalously slow plate tectonic migration rate of 11 km/m.y.; this rate would 

also include any basin and range extension (fig. 4). 

In a related paper by the same authors, Murphy and others (1998) suggest that before 40 Ma, the 

Yellowstone plume was further west (fig. 4) and the plume track had been subducted in advance of the 

time when the actual plume went under the North American plate about 55 Ma. Thus, the plume track and 

associated swell more than 1,500 km long was subducted prior to 55 Ma. Their paper stresses the idea that 

subduction of a plume track swell could add buoyancy to the subducted slab and affect the Cretaceousearly Tertiary orogeny. Their hotspot track differs in rate and orientation from that determined by the 

hotspot track volcanism (Pierce and Morgan, 1992), by hotspot faulting (Anders 1994), and plate motion 

(Alice Gripps in Pierce and Morgan, 1992, p. 6). 

Bob Duncan (1982) proposed "a captured island chain in the Coast Range of Oregon and Washington" and attributes this to the Yellowstone hotspot (fig. 4). He suggested that the plume was shielded from 

surfacing between about 17 Ma and perhaps 30 Ma and inferred that the plume was trapped during this 

10

interval beneath the Vancouver slab. More recently, part of this "captured island" terrain, termed Siletzia 

and dated at 51-55 Ma, has also been ascribed to a Yellowstone hotspot origin by Pyle and others (1997). 

If the hotspot existed before 17 Ma and was active offshore between 50 and 60 Ma (Duncan, 1982; Pyle 

and others, 1997), yet was shielded from surfacing between 17 and perhaps 25 Ma by the sinking 

Vancouver slab, we would expect to see some surface manifestation of this plume west of McDermitt 

between 25 and 50 Ma. Furthermore, if the plume existed before to 17 Ma, why does the Nevada-Oregon 

rift (Pierce and Morgan, 1992; Zoback and others, 1994; Parsons and others, 1994) appear to represent a 

17-Ma event extending about 500 km both north and south of the Oregon-Nevada border centered in the 

McDermitt area and overlain by rhyolitic hotspot track volcanic rocks at McDermitt? How do these 

models (Duncan, 1982; Pyle and others, 1997) account for such a change in magnitude of processes going 

from relatively small volcanic events in the late Eocene in areas well away from the current trend of the 

hotspot track to an approximately 40-Ma period of quiescence and a sudden large event at 17 Ma? 

125° 

45° 

40° 

120° 115° 110° 

➤ 

➤ 

Figure 4. Some positions for the Yellowstone hotspot at different times as postulated by others and by us. The location between 50 and 

17 Ma presents a problem that is avoided if the Yellowstone plume starts at 17 Ma. 

6.6 

Ma 

10 

Ma 

0 500 

Kilometers 

55 Ma 50 Ma

Murphy and others, 1998 

(Pre-55 Ma swell 

east of hotspot) 

40 Ma, Carlin, NV

Oppliger and others, 1997 

Pierce and Morgan, 1992 

17 Ma, Beneath Columbia

River Basalts

Geist and Richards, 1993 

~50 Ma, 

Accreted Islands

Duncan, 1982 

X X 

70 Ma, Yukon

62°N, 138°W

Johnston and others, 1996 

2 

Ma 

16 

Ma 

11

We also suggest that given our understanding of plate motions and rates during this period, it is difficult to 

reconcile a Yellowstone plume off the coast of the Oregon-Washington border considering its present 

location under the Yellowstone Plateau and its track over the past 16 m.y. which points towards Sacramento and the Great Valley of California. 

A late Cretaceous location of the Yellowstone hotspot in the Yukon (fig. 4) was proposed by Johnston 

and others (1996). This idea depends on the 50-Ma location now on the Washington-Oregon coast, and 

has the same problems as discussed above. 

In conclusion, we concur with Draper (1991) that no connection exists between the 0- to 17-Ma 

Yellowstone plume and the various models for an inferred "Yellowstone" plume 50-60 Ma off the coast of 

Oregon and Washington (Duncan, 1982; Pyle and others, 1997). Furthermore, westward displacement of 

the Yellowstone plume head by the inclined Vancouver slab as predicted from plate tectonic and plume 

histories seems physically plausible and explains the problem of an apparent increase in plate motion 

(plus extension) by more than 250 percent (from 7 cm/yr between 10 and 16 Ma to 2.9 cm/yr after 10 

Ma). 

One possible test of the interrelation of the Yellowstone plume head with the Vancouver slab might 

be to look for the effects of the slab and associated subducted sediments in the chemistry of the Washington and Oregon basalts. Takahahshi and others (1998) suggest that the Grande Ronde units of the Columbia River flood basalt result from the melting of a plume head containing very old oceanic crust (Precambrian?). The arc-like geochemical signature of these basalts (high Ba/La ratios; Hooper and 

Hawkesworth, 1993) leads Hooper (1997) and Shervais and others (1997) to suggest the entrainment of a 

subduction component into the plume head that is consistent with our proposed interaction with the 

Vancouver slab. 

THE YELLOWSTONE STARTING PLUME:


TRANSECT ALONG THE NEVADA-OREGON RIFT ZONE


The apparent asymmetry associated with the Yellowstone starting plume at 14-17 Ma can be explained by changes in the thickness and composition of the lithosphere along a north-south axis parallel to 

the Nevada-Oregon rift zone (figs. 1 and 5). We place the initial plume head center at the McDermitt 

caldera complex along the Nevada-Oregon border (fig. 1). The Columbia River-Oregon Plateau flood 

basalts extend for hundreds of kilometers to the north of, and give a lop-sided appearance to, the track of 

the Yellowstone hotspot, although we emphasize that this plume-head center is at about the midpoint of 

the Nevada-Oregon rift zone. Nonetheless, the location of these flood basalts with respect to the inferred 

initial location of the starting plume requires explanation. Geist and Richards (1993) explained this 

northward location of the Columbia River flood basalts by a northward diversion of the plume head by 

the northward-moving Vancouver slab. 

We suggest that a south-to-north change from an older cratonic crust with a thicker lithospheric 

mantle to a younger, more oceanic crust with a thinner lithospheric mantle (fig. 5) controlled the surface 

eruption of flood basalts from the plume (Zoback and Thompson, 1978; Pierce and Morgan, 1992; 

Zoback and others, 1994; Parsons and others, 1994; Takahahshi and others, 1998). As noted by Hooper 

(1997), nearly all of the Columbia River basalts erupted through fissures in thinner lithosphere made up 

of accreted oceanic crustal material. The earliest eruptions occurred along a north-south suture on the 

west side of the Precambrian crust and perpendicular to the minimum principal stress of the regional midMiocene state of stress (Zoback and others, 1994) and parallel to a Miocene back-arc spreading system 

(Zoback and others, 1981; Parsons, 1995). This denser, more mafic crust would facilitate basaltic melts 

rising to the surface and being erupted as (1) the denser crust increases the lithostatic pressure per crustal 

unit depth on the magma chamber, and (2) the mafic material is more refractory and thus less likely to 

12

melt. That the plume head uplift resulted in north-south rifting rather than the classic tensional rifts at 120 

degree angles appears to relate to this east-west orientation of the minimum principal stress. 

The extent of flood basalt now appears to include northern California and the Lovejoy Formation 

dated about 16 Ma (Wagner and Saucedo, 1990; Page and others, 1995). Wagner (oral commun., 1999) 

estimates the volume of this basalt to be near 75,000 km3

, including much basalt beneath the northern 

Sacramento Valley. 

The northern Nevada rift zone was intruded by 14- to 17-Ma mafic magmas (Zoback and Thompson, 

1978; Zoback and others, 1994). This extension zone continues northward to the feeder dikes of the 

Oregon Plateau and Columbia River flood basalts and southward along the extension of the northern 

Nevada rift into southern Nevada (Blakely and Jachens, 1991) and may include the 14- to 17-Ma rhyolites 

of southern Nevada. All together, these extensional features form the 1,100-km-long, 17-Ma NevadaOregon rift zone and centers on the inferred initial sublithospheric position of the Yellowstone plume head 

(Pierce and Morgan, 1992). Although feeder dikes for the Columbia River flood basalts are as much as 

500 km north of this inferred center, White and McKenzie (1989) note that if mantle plumes coincide with 

active rifts, large-volume basalt eruptions can extend along rifts for 2,000 km (half distance = 1,000 km). 

Although some plume heads are associated with radial dikes extending thousands of kilometers (Thompson, 1998), the linear N. 20°-25° W. Nevada-Oregon rift zone indicates that at 17 Ma from eastern 

Washington to southern Nevada, the least principal stress was horizontal and oriented N. 65°-70° E. 

(Zoback and Thompson, 1978; Christiansen and McKee, 1978; Zoback and others, 1994). 

Figure 5 shows modern geologic and geophysical characteristics along a north-south cross section 

parallel to the 14-to 17-Ma Nevada-Oregon rift zone (A-A', fig. 1). West of the 87Sr/86Sr 0.704 line is 

young, mafic oceanic terrane in Oregon and Washington in the area of flood basalt extrusion. East of the 87Sr/86Sr 0.706 line is Precambrian, continental crust in southern Nevada. General differences along the 

north-south section are as follows using McDermitt as the midpoint: Transect A shows basalts occurring 

in the farthest north segment; rhyolites, basalts, and tuffaceous sediments are exposed in the middle with 

trachybasalts exposed farther south; mafic intrusives with rhyolites occur in the southernmost segment. 

Transect B shows the terrane to the north as mostly Mesozoic and younger mafic oceanic and oceanic 

accreted terranes and the terrane to the south as increasing in age from a Paleozoic magmatic arc (Elison 

and others, 1990) to an interval along the Sr 0.706 line between a Paleozoic arc (to the west) and late 

Proterozoic sedimentary rocks (to the east; Link and others, 1993). Transect C shows that heat flow has a 

broad culmination over the inferred position of the plume head near the Nevada-Oregon border. Transect 

D shows terrain altitude changes from north to south: higher terrain in Canada descending to about 300 m 

on the Oregon-Washington border, then a bench above 1,500 m in southern Oregon and northern Nevada, 

to a culmination in central Nevada at about 2,400 m, and finally a decrease to about 300 m in southern 

Nevada. Transect E is the average of complete Bouguer gravity anomaly values based on stacking (averaging) of the five parallel profiles centered on A-A' (fig. 1). The averaged profile shows prominent steps 

near the Washington-Oregon border and in southern Nevada separated by a major low (Saltus and Thompson, 1995). 

Transect F is a two-dimensional model of the gravity profile. The southern boundary of anomalous 

asthenosphere on it is based on seismic, heat-flow, and isotopic constraints (Saltus and Thompson, 1995); 

the northern boundary of anomalous asthenosphere is based, by analogy, on the position of the complementary step in gravity and topography near the Washington-Oregon border. The crust/mantle boundary 

in the model is an average of three different western U.S. Moho maps presented in Geological Society of 

America Memoir 172 (Mooney and Weaver, 1989; Pakiser, 1989; Braile and others, 1989). We have 

allowed the midcrustal interface between felsic upper crust and mafic lower crust to vary in order to fit 

the remaining anomalies. To aid understanding, we have used absolute densities, albeit these are somewhat arbitrary; the gravity model is only sensitive to lateral variations in relative density, not to the 

13

absolute values. The model indicates that the gravity data are consistent with the seismically determined 

average Moho, with our hypothesized anomalous hot or light asthenosphere (hot plume head and restite, 

both assumed to be 0.05 g/cm3

 lighter than regular asthenosphere, fig. 5), and with a crust which ranges 

from generally mafic (shown as lower crust) in the north to generally felsic (shown as upper crust) in the 

south. 

Figure 6 illustrates schematically along a north-south axis how a spreading plume head 500-1000 km 

across (Hill, 1972) might interact along the section A-A' with crustal changes reflected by the 87/86Sr 0.704 

and 0.706 lines (fig. 1). To the north of center is relatively thin, dense, accreted oceanic crustal lithosphere, whereas to the south of center is progressively older, more silicic, more continental crustal material 

(Kistler, 1983; Elison and others, 1990; Mooney and Braile, 1989; Camp, 1995, Link and others, 1993). 

The rising and spreading plume head (density of 3.25 g/cc) would undergo decompression melting above 

a depth of about 150 km, producing a basaltic melt (density 3.0 g/cc) and also leaving a restite (also 

density 3.0 g/cc). North of the McDermitt area, voluminous flood basalts erupted through the Chief 

Joseph dike swarm that is associated with the foliated and sheared suture zone between the Precambrian 

continental crust on the east and the oceanic lithosphere and intraoceanic island arc terranes on the west 

(Hooper, 1997; Camp, 1995; Snee and others; 1995; Vallier, 1995) also delineated by the 87/86Sr 0.704 and 

0.706 lines (fig. 1). As observed by Hooper (1997), a significant lithospheric signature is present in all but 

the earliest of the Columbia River basalts and, in fact, varies spatially through time. The Columbia River 

basalts can be divided into three basic subgroups of enriched subcontinental lithospheric mantle that were 

entrained into the plume head. 

For the Mesozoic and younger oceanic accreted terrane north of the midpoint, plume-head material 

would flow towards higher areas or “thinspots” at the base of the lithosphere, thereby allowing for an 

increase in decompression melting and forming basaltic magmas (Thompson and Gibson, 1991; Sleep, 

1990). According to Hooper (1997), the Columbia River basalts were mantle-generated magmas that had 

relatively long residence times in reservoirs at the base of the crust and that erupted periodically and 

rapidly through NNW-SSE-oriented fissures to form the voluminous flood basalts. Because crust in this 

area was relatively dense, thin, and weak, and was adjacent to a major tectonic boundary of thicker, more 

competent, continental lithosphere (Hooper, 1997), the basaltic magma would be readily able to rise to the 

surface where it formed flood basalts of the Columbia River and Oregon Plateau. As described by Camp 

(1995), basalt flows become younger to the north and reflect the northward spread of the plume head as it 

intersected the lithosphere and pancaked outward. In addition to the northward migration of basalt flows, 

Camp (1995) also documents a northward progression of uplift during Columbia River basalt time. 

Between 14 and 17 Ma, northward-directed compression formed the east-west Yakima folds and associated thrusts (Hooper, 1990; Reidel and others, 1989), also consistent with plume head uplift from the 

south. 

Magmatism from near and to the south of the Nevada-Oregon border was different from the flood 

basalts farther north (fig. 1): (1) rhyolites and basalts erupted in the region of the northern Nevada rift, (2) 

trachybasalts erupted in the central area, (3) unexposed basaltic dikes, inferred by observed magnetic 

anomalies, and isolated rhyolitic volcanic fields are in the south (Zoback and others, 1994). We think the 

following processes occurred beneath this area: decompression melting generated hot basalt that rose into 

the silicic crust where, due to the higher melting temperature of basalt, heat from the basaltic melt produced silicic magmas that then rose higher in the crust to form high-level magma chambers, from which 

erupted rhyolitic ignimbrites and lava flows. 

Undoubtedly, the cross section along A-A' would have been somewhat different at 17 Ma. In 

particular, in northern and central Nevada, the lithosphere was probably thicker and has been subsequently thinned by basin and range extension. A significant amount of basaltic underplating, which we 

associate with the starting plume, has thickened the crust near the Moho in Nevada (Thompson and 

14

706 706 

704 704 

706 706 

➤ ➤ ➤ ➤ ➤ ➤ ➤ ➤ 

➤ ➤ ➤ ➤ 

-600 -400 -200 0 200 400 600 800 -600 -400 -200 0 200 400 600 800 

OREGON 

NEVADA 

upper crust 

mantle lithosphere 

hot asthenosphere and restite Asthenosp

Feeder dikes Feeder 

dikes 

Rhyolites 

34-37 28-34 21-28 

N. NEV. RIFT 

Flood Basalts 

Early Proterozoic Parallel to line 

of section 

Parallel to line 

of section 

Mesozoic and younger oceanic, mafic lithosphere with accreted oceanic terranes 

Paleozoic Arc 

Older volcanism 

Rhyolites 

Late Proterozoic sed. rx apparently on Early Proterozoic or Archean 

Observed Calculated 

➤ 

A 

North 

lower crust 

D=2.67 

D=3.35 

D=3.25 

➤ 

D=2.85 

A 14-17 Ma 

Magmatism 

B Crust, Age 

& Origin 

C Heat flow

105 

63 

D Topography 

3 2 1 

E 

0 

Bouguer -800 Gravity 0 

-50 

-100 -150 -200 -250 

F Model 0 

50 

100 150 

-800 

V.E.=2 Scale=7500000 704 704 

Figure 5. See list of Figure Captions for full explanation. Distance (km) 

Kilometers 

15

North 704 704 So

706 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ ➤ 

X X X X X X 

Flood Basalt, 14-17 Ma 

Yakima 

foldbelt 

Plume 

head 

➤ ➤ 706 ➤ ➤ 

➤ 

0 

100 200 

Km 

100 Km 

0 

0 

Scale 

Vertical Exaggeration = 2x 

OREGON 

NEVADA 

Mafic "rift" intrusion 

Young oceanic (mafic) crust with accreted island-arc terranes 

Parallel 

to line of 

section 

Parallel to line 

of section 

Early Proterozoic 

Upper 

Crust 

Late Proterzoic seds 

on Pre C crust ? 

Decompression melting producing 

basalt and restite 

Basalt underplating, dikes 

parallel to section, not shown 

21-34 Ma magmatic arcs, depleted crust (?) Rhyolite magma chamber 

EXPLANATION 

X X

Mantle 

Lithosphere 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ ➤ ➤ ➤ ➤ 

Rhyolite ignimbrites and flows, 14-16 Ma 

Paleozoic 

Arc 

Basalt feeder 

dikes 

Crust Lower 

➤ ➤ 

A A

Figure 6. Geologic cartoon along line of section A-A’ showing emplacement of postulated Yellowstone plume head. The plume head spreads out (pancakes, mushrooms) upon rising into the lithosphere, bu

X 

➤ 

thinner lithosphere (young thin mafic crust and lithospheric mantle) beneath Oregon. Decompression melting (squiggly vertical lines) produces basalt magma that rises upward and interacts differently with magma rises through the dense, more oceanic crust, and surfaces as the Columbia River and Oregon Plateau flood basalts, and (B) in Nevada and southern Oregon, basalt magma melts silicic lower crustal upper crust forming magma chambers (oval pods) and rhyolitic ignimbrite eruptions, shown as volcanoes, and (C) in northernmost Nevada and southern Oregon, both processes have operated. The restite fremains in the mantle but is buoyant and contributes to long-lasting uplift (Morgan and others, 1995). At the time of emplacement, doming above the plume head may be 1-2 km (exaggerated in drawing). push from this doming may have produced the Yakima fold belt near the northern end of the section. 

16

others, 1989). Crustal thickness may also have been thinner to the north before basaltic underplating 

associated with the 15-Ma and younger volcanism and the mantle lithosphere was likely thin beneath this 

young oceanic terrane. The hot, spreading plume head may have thinned the mantle lithosphere by 

thermal and mechanical erosion. Because topographic changes also have transpired over the last 17 m.y., 

the area above the plume head would have been topographically higher when the plume head first intersected the North American plate at 17 Ma than today. As suggested by Parsons and others (1994), the onekm high present topography centered on McDermitt results from the residual effect of low-density 

Yellowstone plume material. 

The plume head and associated Nevada-Oregon rift are spatially near the center of the active Basin 

and Range (fig. 4). We suggest a causal relationship exists between the inferred plume head and changes 

in the Basin and Range. Upon impact with the lithosphere, the plume head would decrease its ascension 

rate; this impact is associated in time and space with a change in Basin and Range extension and volcanism. While not advocates of a plume origin for this region, Christiansen and Yeats (1992) note, "The 

bimodal rhyolite-basalt magmatism of the Great Basin region is mostly younger than 17 Ma, following a 

widespread magmatic lull. By about 17 Ma, significant uplift had begun to be the dominant factor in the 

1,600-km region that encompassed the Great Basin Region, the Columbia Intermontane region, and 

surrounding areas." For the 14-10-Ma interval, they write (p. 388 and Plate 7): "Regional extension in the 

continental interior changed to widely distributed normal faulting between about 14 and 10 Ma with 

accelerated uplift of the region from the Sierra Nevada to the High Plains." 

Furthermore, northward migration of the source of flood basalts is represented in the spreading of the 

plume head (Camp, 1995). The remarkable N. 20°-25° W. orientation of the 1,100-km Oregon-Nevada 

rift zone indicates the minimum principal stress over this great length at 17 Ma was about N. 65°-70° E. 

and subparallel to the west coast (Zoback and Thompson, 1978; Zoback and others, 1994). 

SOUTHERN OREGON RHYOLITE BELT AND THE PLUME HEAD 

Many have suggested that the southern Oregon rhyolite belt appears to contradict the plate-tectonic 

Yellowstone hotspot hypothesis in that it is a similar-aged volcanic progression that advances WNW 

across southern Oregon (MacLeod and others, 1976; Christiansen and Yeats, 1992, p. 381-382). The 

vector of this progression makes a 120-degree angle with the vector of the SRP-YP volcanic progression 

(fig. 1). Both emanate from the tri-state boundary area and have similar ages. Pierce and Morgan (1992, p. 

32-33) note significant differences in the rhyolitic volcanism between the eastern SRP-YP trend and the 

southern Oregon trend that suggest different processes for the two progressions. Differences in the two 

volcanic provinces include the time of inception (the southern Oregon belt began about 10 Ma and 

McDermitt erupted at 16.1 Ma); the style of volcanism (small rhyolitic domes and small-volume ignimbrites are typical in southern Oregon whereas large-volume ignimbrite eruptions are typical in the SRPYP province), and the volcanic migration rates. The Brothers strike-slip fault zone forms the northern part 

of the southern Oregon rhyolite trend. It also separates basin and range and associated extension on the 

south from the High Lava Plains with much less extension on the north. The map pattern of the Brothers 

fault zone (Walker and others, 1981; Pezzopane and Weldon 1993) is represented by short, small-offset, 

normal faults that are arranged in en echelon patterns oriented 10-20 degrees clockwise from the overall 

Brothers fault-zone trend. These extensional openings are associated with right lateral shear that may 

have provided conduits for volcanic eruptions. Draper (1991) suggests that both the spreading 

Yellowstone plume head and the activity on the Brothers fault zone started in northern Nevada-southern 

Oregon and migrated to the WNW. Volcanism followed this WNW migration of fault activity and plume 

spreading. 

Some possible mechanisms that might explain the northwesterly volcanic progression in southern 

17

Oregon are: (1) counterflow associated with the WNW flow at the base of the lithosphere; (2) plume 

spreading (Sleep, 1997); and (3) the "up-and-out-welling" at the edge of a hot, thick body that might 

produce a WNW drag (fig. 7). A mechanism that might drag the western margin of the Yellowstone plume 

head further westward and produce the volcanic trend observed in southern Oregon is the counterflow or 

backflow occurring in the acute angle between the descending Vancouver slab and southwest-advancing 

North American plate (Draper, 1991). In addition to this counterflow, Eugene Humphreys (oral commun., 

1997) and others suggest a thermal convective upwelling along the gradient of the edge of any hot mass in 

the mantle. 

McDermitt Calderas 

Quaternary 

Basaltic lavas 

Rhyolitic 

ash flows 

Rhyolitic domes 

showing K-Ar age 

OREGON 

Area of figure 0 20 40 60 80 100 

Kilometers 

Figure 7. Age progression of silicic lava domes of the southern Oregon rhyolite belt. This figure shows upper Miocene and younger 

volcanic rocks and is copied from Christiansen and Yeats (1992, fig. 55). The rhyolite domes and their ages were shown by MacLeod and 

others (1976) to have such a systematic northwest decrease in age that ages could be contoured in one-million-year increments. This 

northwest age progression is a key argument against a plume origin for the Yellowstone hotspot (Lipman, 1992; Christiansen and Yeats, 

1992). See Figure 9 for our suggestions about how the southern Oregon rhyolite belt might be reconciled with a thermal plume origin of 

the Yellowstone hotspot. Both the southern Oregon rhyolite belt and the Yellowstone hotspot track start in the region near McDermitt, but 

they become younger in nearly opposite directions at the same time. The silicic domes (dark) form two belts that trend N. 75o W. The 

northern belt ends at Newberry caldera and parallels the Brothers fault zone (wide gray line) which is actually a complex en echelon 

pattern of normal faults whose trend of about 10o clockwise to the overall pattern (Walker and Nolf, 1981) suggests a right lateral 

component (Pezzopane and Weldon, 1993). Between 10 and 17 Ma, ignimbrites and other silicic volcanism are widespread in the 

southern Oregon, northern Nevada and southwestern Oregon area (fig. 1; Pierce and Morgan, 1992;) but have no systematic progression 

with age in southeast Oregon (MacLeod and others, 1976, p. 470). 

18

Draper (1991) notes that no younger basalts overlie the Columbia River basalts. He attributes this to 

a combination of (1) magma reservoir depletion exhausted by the eruption of the Columbia River and 

Steens Mountain basalts and (2) extension beginning at about 10 Ma that forced the plume head to spread 

out laterally at the base of the crust. This process resulted in abbreviated magma residence times and 

generated primitive HAOT (high-alumina olivine tholeiitic) magmas to the west of the Columbia River 

basalts. Draper (1991) suggests that as extension increased with time, smaller volumes of primitive 

magmas erupted. According to Draper (1991), the processes of extension combined with a laterally 

expanding plume head under a crust depleted in low-melting-point components. The result produced 

decreasing volumes of rhyolitic material over time. The WNW progression of silicic activity in the 

southern Oregon belt (figs. 7 and 8) was concentrated along the edge of the westward-expanding plume 

head influenced by the sinking Vancouver slab. 

In contrast to this minor rhyolitic activity in southern Oregon, large volumes of rhyolitic ignimbrites, 

tephra, and lavas erupted along the SRP during this comparable interval of time (Bonnichsen, 1982; 

Perkins and others, 1998; Morgan and others, 1997; Morgan and others, 1984; Morgan and McIntosh, 

2001; Christiansen, 1984). The northeast progression of rhyolitic volcanism in the SRP-YP has been 

attributed to the North American plate directly overriding the chimney or tail phase of the thermal plume 

(Pierce and Morgan, 1992) that melted a continuous supply of undepleted crustal material (Draper, 1991). 

➤ 

➤ 

➤ ➤ 

➤ 

➤ 

➤ ➤ ➤ ➤ ➤ ➤ 

0 100 200 Km 

1 2 3 4 5 6 7 8 9 10 16 

plume head 

mantle lithosphere 

McDermitt 

Age of rhyolitic volcanism, in Ma, 

North American Plate motion, ~25 km/m.y. to WSW 

counter flow 

Coast line 

Coast Range 

Cascades 

Oceanic lithosphere 

Formation of rhyolite magmas 

and their extrusion 

Progressive westward penetration 

of mantle lithosphere (see caption) 

Slab sink ~5 cm/yr

➤ 

➤ 

➤ 

➤ 

Pacific Ocean 

➤ ➤ ➤ ➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

➤ 

WNW ESE ➤ 

showing migration of activity to the northwest 

calderas 

0 

40 Km 

50 

100 

Figure 8. Geologic sketch (after Draper, 1991) along the southern Oregon rhyolite belt parallel to the trend of the Brothers fault zone. The mounds numbered 

10 to 1 represent contours on the age or rhyolitic volcanism in Ma (MacLeod and others, 1976). Following Draper (1991), we suggest several factors may 

have acted separately or in combination to produce penetration of mantle lithosphere by basaltic melts to form rhyolite magma from crust: (1) cracking of 

mantle lithosphere by westward migration of activity on Brothers fault zone, (2) westward spreading of plume head, and (3) thermal erosion of mantle 

lithosphere associated with counterflow. Fracturing of the lithosphere along the Brothers fault zone (MacLeod and others, 1976; Carlson and Hart, 1987; 

Pezzopane and Weldon, 1993) appears to have provided the conduits for volcanism as well as its overall northwest trend. The most reasonable mechanism for 

the northwest progression of rhyolitic volcanism may be a northwest progression of faulting to form the Brothers fault zone. Such a northwest progression of 

faulting might have been accompanied by a northwestward migration of plume head material aided by both general outward migration of the plume head, and 

by counter flow between the Vancouver slab and North American plate (Draper, 1991). If the two proceeded together, the faulting may have caused fracturing 

of the mantle lithosphere permitting the erosion and stooping of the mantle lithosphere (a mechanism suggested by George Thompson, oral commun., 1998) 

and thus forming an inverted channel enhancing both outward (northwestward) plume head flow and decompression melting by thinning of the lithosphere. 

Although Draper (1991) drew a similar diagram, we include this version because many have argued that this northwestward volcanic progression negates the 

Yellowstone hotspot as a mantle plume, and we think these factors offer a mantle-plume compatible explanation. The geometry of the subducting slab is 

based on a fairly well controlled section by Parsons and others (1998) near the Oregon-Washington border. 

➤ 

19

In conclusion, we concur with Draper (1991) that the southern Oregon rhyolite belt may reflect NW 

migration of either faulting or spreading of the Yellowstone plume head and that this progression does not 

necessarily negate the SRP-YP as a hotspot track of a thermal mantle plume. 

PLUME-HEAD UPLIFT AND ASSOCIATED CLIMATE PATTERNS 

Plume heads are expected to cause uplift of one kilometer or more and have diameters of 1,500 to 

2,500 km (Hill and others, 1992). We estimate the Yellowstone plume head was originally a 400-kmdiameter sphere. A considerable size is suggested by the 15-17-Ma volcanic activity along the 1,100-km 

length of the Nevada-Oregon rift zone. Parsons and others (1994, and Tom Parsons, written commun., 

1998) consider a minimum 800-km diameter for the flattened plume head based on the current mass 

deficit in the upper mantle that extended over much of Nevada and parts of Utah and Oregon. Given the 

elongate pattern of tectonic and volcanic activity aligned subparallel to the Precambrian margin of the 

North American plate (Camp, 1995) as well as parallel to the back-arc margin inland from the Pacific 

Ocean (Zoback and others, 1981; Parsons, 1995), we favor (fig. 9) an elongate north-south spreading 

plume head similar to that proposed by Camp (1995). 

Within the Columbia River basalts (fig. 9 and table 1), northward offlap of basalt units with time is 

shown by Camp (1995, fig. 4) to reflect south-to-north migration of uplift, with local uplift rates of about 

2/3 mm/yr. This parallels the northward migration of Columbia River basalt source areas (Camp, 1995). 

Estimates of past altitudes are difficult to reconstruct, but a modern leaf-morphology technique 

based on paleo-enthalpy differences between sea level and inland localities has promising results (fig. 9 

and table 1). From detailed analysis of leaf physiognomy, Wolfe and others (1997) suggest that at about 

15-16 Ma the surface of west central-Nevada was more than 1 km higher than now. We have reservations 

about this technique, in part because its results conflict with numerous studies that we find of merit for 

late Cenozoic uplift in the Colorado Rockies. But we do find this result for Nevada surprisingly compatible in our altitude estimate, general location, and age for plume-head uplift. Further east, an Eocene 

fossil-leaf flora is compatible with eastward migration of Yellowstone hotspot uplift: just south of the 

present location of the Yellowstone hotspot is a 50-Ma locality that is estimated to have been elevated 0.9 

km to its present altitude sometime in the last 50 Ma and therefore after the Laramide orogeny (fig. 9; 

Wolfe and others, 1998). This is compatible with uplift associated with the hotspot as it migrated 650 km 

from the plume-head area at 16 Ma to its current position at Yellowstone at 2 Ma. 

Our hypothesized uplift of one km over a north-south distance of 1,000 km (fig. 9) should have a 

significant effect on local weather patterns. The expected effect would be increased precipitation on the 

west side from the orographic moisture extraction of rising airmasses coming inland (eastward) from the 

Pacific Ocean and then orographic drying and aridity east of the crest as these airmasses descend the lee 

side of the uplift. A precipitation shadow might extend further inland (east) as far as the main moisture 

source was from the Pacific. This pattern is analogous to the present Sierra Nevada with a wet west side, a 

dry east side, and a precipitation shadow extending far eastward across the Great Basin. 

The geologic history of the Tertiary Bozeman Group in the Montana-Idaho area suggests the climate 

became more arid about 15 Ma at the time the plume-head uplift would have culminated (table 1 and fig. 

9, locations 9 and 10; Thompson and others, 1982; Fields and others, 1985). In the Yellowstone Valley 

north of Yellowstone Park, 15-Ma saline lake deposits with gypsum and anhydrite formed because of 

aridity. Furthermore, associated rodent fossils at nearby localities suggest aridity was more severe than at 

20 Ma (table 1, location 11; Barnosky and Labar, 1989). 

Paleobotanical transects from eastern Washington and Oregon to the Rocky Mountains (fig. 9 and 

table 1, locality 13; Leopold and Denton, 1987,) show clear drying inland, but do not show any compelling evidence for plume-head uplift. Between 12-17 Ma, vegetation reconstructions show in eastern 

20

125°� 120°� 115°� 110°�

45°�

40°�

35°�

0 500 

Kilometers 

➤ 

➤ ➤ 

➤ 

➤ ➤ 

➤ 

➤

➤ 

➤ 

➤ ➤ 

➤ ➤ 

➤ ➤ 

➤ ➤➤ ➤ 

2 

Ma 

6.6 

Ma 

10 

Ma 

+ 0.4 Mio 

- 0.7 Mio 

3 

2 

2 

9,10 

Precipitation shadow 

(decrease) 

Yellowstone crescent 

of high terrain 

(modern) 

1, Area of inferred 

1 km plume-head 

uplift 

precipitation 

increase 

13, 

Paleovegetation transects 

13, 

P aleo v egetation transects 

- 1.4 Mio 

- 0.2 Mio 

+ 0.1 Mio - 1.2 Mio 

- 1.0 

- 1.3 

- 1.2 

- 1.1 

- 1.0 Mio 

- 0.6 Mio 

16 Ma 

6.6 

Ma + 0.2 Eoc 

7,8 

6 

6 

6 

6 

6 

6 

12 

+ 0.9 Eoc 

11 

- 0.6 Eoc 

- 0.8 Eoc 

Figure 9. Uplift and climatic patterns that may relate to Yellowstone plume (both head and tail). The dashed line shows the pancaked plume with a diameter 

suggested by Parsons and others (1994) of about 400 km but elongated north south as suggested by Camp (1995) parallel to the Nevada-Oregon rift zone. Table 1 

describes the various localities, areas, and transects shown here. One of the main climatic parameters listed is orographic dryness (light shading) in the lee of the 

inferred Yellowstone plume-head uplift. In general, observations of uplift and climate are compatible with uplift above the plume head followed by uplift migrating 

to the east associated with the plume tail (chimney) to the "Yellowstone crescent of high terrain" associated with the present plume tail (Pierce and Morgan, 1992) 

or a similar 650-m-high altitude anomaly that Smith and Braille (1993) associated with the current Yellowstone hotspot. 

21

Table 1. A sampling of possible indicators of uplift, altitude, precipitation shadow, and climate that may be related to Yellowstone hotspot plume uplift. 

Transect from Pacific Coast inland to Rocky Mountains 

Comments 

Large, low-density anomaly in the mantle interpreted to be Yellowstone plume head. Diameter about 800 km 

and associated with active basin and range. 

Offlap and thickness of subsequent CRB units indicate south to north uplift, locally at a rate of 0.67 mm/yr. 

The Yakima fold belt on the northern margin of the CRB was formed by south to north compression. CRB flows 

were being deformed shortly after their emplacement. We here suggest that the age and pattern of deformation 

are consistent with Yellowstone plume head uplift and north-directed gravitational forces on the north flank of the 

uplift (perhaps others have already suggested this). 

Paleobotanical analysis using leaf physiognomy of middle Miocene assemblages indicates subsequent 

subsidence (-) or uplift of stated amount in km. "When the standard error is applied to two coeval sites, the 

combination of the two errors produces a standard error in the estimated difference in altitude of ~760 m." 

Paleobotanical analysis using leaf physiognomy of Eocene assemblages indicates subsequent subsidence (-) or 

uplift of stated amount in km. 

Drainage divide on present Snake River Plain 75 km east of Twin Falls with drainage west of this divide down 

present Snake River Plain and out towards California, and east of this divide drainage from Jackson Hole and north of SRP southward into present Bonneville Basin. This suggests an ancestral high (hotspot?) east of Twin 

Falls in the late Miocene. 

Fish ecology and oxygen-isotope analysis on fish otolith suggest warm, moist climate with milder winters (16°C 

warmer than present) and cool summers (1° cooler than present). Oxygen isotope (SMOW) calculated for water -15.5 per mil and annual cycle in fish from -16.6 to -13.5 per mil. 

Water in glass spheres has δD value of -147 per mil, whereas present-day meteoric water at site has value of -125 

per mil. This value is as high as that for drainages in the highest country at the upper end of the hotspot track in the mountains surrounding Yellowstone (Irving Friedman, written commun., 1998), suggesting the possibility that this area was much higher at the time of deposition. 

Change from a wetter to a more arid climate at about 16 Ma may reflect inception of precipitation shadow from 

plume head. 

Climate change from wet to dry ~15 Ma (and dry to wet ~ 20 Ma) 

Saline lake deposit with gypsum and halite indicates aridity and drier than present. Rodent type suggests even 

more arid than semiarid Arikareean. May be in precipitation shadow of plume uplift to west. 

Lake sediments rich in carbonate suggest dryness. Oxygen isotopes (-5.4 to 8.7 per mil, SMOW) "are the lightest 

yet reported for a nonmarine carbonate." Lake water estimated to be -20.7 to -34.8 per mil. SMOW and modern precipitation to be -16.5 per mil, SMOW (-30 per mil oxygen SMOW converts to -160 per mil Deuterium). 

Either much higher than present or in major precipitation shadow (of plume uplift?) or a combination of the above. The similarity in age to the Chalk Hills Formation but the much lighter isotopes suggests this site may have been in a precipitation shadow related to uplift of the plume tail. 

Rocky Mountains 

Montane conifer forest with steppe 

openings. Precipitation shadow? 

Montane conifer forest with steppe 

openings. Precipitation shadow? 

Idaho 

Deciduous hardwood forest and 

Taxodium swamp 

Montane conifer and deciduous 

forest 

E. WA & OR 

Deciduous hardwood forest and 

Taxodium swamp 

Clarendonian 

Montane conifer and deciduous forest, 

Taxodium swamp 

Unit 

Low density 

anomaly in the mantle 

Columbia River 

basalt 

Yakima fold belt 

Middle Miocene 

leaves 

Eocene leaves 

Late Miocene 

Chalk Hills Fm. Chalk Hills Fm. 

Bozeman Group Six MileRenova 

unconformity 

Hepburn Mesa 

Fm. Barstovian 

Camp Davis Fm. 

Barstovian/Late 

Hemingfordian 

Clarendonian 

Age 

~15? Ma 

17.5-14 Ma 

17-14 Ma 

and 

continuing 

~14-16 Ma 

~40-50 Ma 

9?-5? Ma 

5-9.5 Ma 5-9.5 Ma 

16 Ma 

~15 Ma 

15 Ma 

~8-6? Ma 

12-17 Ma 

8-12 Ma 

Loc. 

1 

2 3 

4 = "Mio" 

5 = "Eoc" 

6 

7 8 

9 10 11 12 

13 

13a 

13b 

similar to 4 

22

Oregon/Washington, and northern Idaho- deciduous hardwood forest and Taxodium swamps; in central 

Wyoming and Colorado- montane conifer forest with steppe openings. Between 8-12 Ma, reconstructions 

show: in eastern Oregon/Washington, and southern Idaho - montane conifer forest and deciduous forests; 

in northwest Wyoming- montane conifer forests with steppe openings. The Rocky Mountain localities are 

permissible of a precipitation shadow in the lee of plume head uplift to the west 

The following history of vegetation in the SRP area is compatible with decrease in altitude (Leopold 

and Wright, 1985) that may have followed plume head and tail related uplift: (1) Miocene- deciduous/ 

conifer forest; (2) Pliocene- conifer forest with some open grassland with increasing numbers of grazing 

horses; and (3) Pleistocene, steppe vegetation and alkaline lakes. 

In Jackson Hole, <7.5- to 10.3-Ma lacustrine beds (Love and others, 1997) have carbonate that 

Drummond and others (1993) concluded had the lightest 18O/16O ratio yet recorded for nonmarine carbonate sequences (fig. 9 and table 1, locality 12). In the similar-aged 7-Ma Chalk Hills Formation in the 

western SRP (fig. 9 and table 1, locality 8), deuterium in glass spheres is 22 per mil lighter than modern 

water (Friedman and others, 1993), which may suggest that the western plain was higher at 7 Ma. Also 

from the Chalk Hills Formation, oxygen isotope studies of fish otolith vary from -16.6 to -13.5 per mil in 

water calculated to have been -15.5 per mil. These data also suggest the water source for the 7-Ma Chalk 

Hills Formation was at a higher than present elevation. Uplift associated with the inferred 7-Ma plume 

position near Pocatello (Pierce and Morgan, 1992) may have resulted both in higher terrain in the Chalk 

Hills area of the western SRP and orographically drier climates east of this highland resulting in very light 

oxygen isotopes in the Jackson Hole area. Amundson and others (1996) summarize for the western United 

States some available information on oxygen isotope distribution. 

Using the regional distribution and paleobiogeography of mollusks, Taylor and Bright (1987, fig. 5) 

locate a late Miocene (5?–9? Ma) drainage divide near American Falls, Idaho (about 75 km east of Twin 

Falls) with drainage east of this divide going south into the present Bonneville Basin (fig. 9 and table 1, 

locality 6). Wood and Clemens (2001) also invoke a northeast shifting drainage divide to explain an 

increase in drainage basin size associated with a rise of Lake Idaho from a low about 6-7 Ma to a high 

near 5.5 Ma (their fig. 7). 

GENERAL CLIMATE HISTORY

The marine record of climate change shows overall cooling throughout the Cenozoic, which in finer 

detail includes some steps, plateaus, and peaks (fig. 10). Coffin and Eldholm (1994, and references 

therein) show the correlation between the Columbia Plateau "large igneous province" and the 14-17 Ma 

Miocene warm interval, which was followed by cooling. The 14- to 17-Ma climatic optimum is recognized in deep-sea sites (fig. 10 shows compilation by Barron and Keller, 1982, fig. 1, and site 747 on 

Kerguelen Plateau by Wright and Miller, 1992). A major global warming at 18-15 Ma also is recognized 

on land by "dramatic northward movement of temperate deciduous forests north of the Arctic Circle in 

Alaska and northern Canada" (Thomas Ager, written commun., 1996). Hodell and Woodruff (1994) 

attribute this middle Miocene climatic optimum to the warming effect of carbon dioxide and other material associated with extrusion of the Columbia River flood basalts. Significant carbon dioxide emissions 

probably accompany flood basalt and other volcanic activity (Arthur and others, 1985; McClean, 1985, 

Leavitt, 1982, Rampino, 1991). 

Self and others (1997) estimate that the volatile release from the Rosa unit of the Columbia River 

basalt introduced significant amounts of S, Cl, and Fl into the upper troposphere and possibly the lower 

stratosphere, thus significantly affecting the global atmosphere. Eruptions of the type, composition, and 

magnitude as that present in the Columbia River basalts, sustained over periods of years to decades, 

would have "strong, detrimental effects on global climate" (Self and others, 1997). After 14 Ma, the great 

23

majority of the total volume of the Columbia River basalt had been erupted (Baksi, 1989) and emissions 

of carbon dioxide associated with plume head volcanism decreased, compatible with the cooling of 

climate by 14 Ma. 

Although uplifts effect climate patterns, Raymo and others (1988) and Raymo and Ruddiman 

(1992), postulate the largest effect on climate associated with uplift of the Himalayan-Tibet areas was 

through the reduction in carbon dioxide during enhanced silicate weathering of fresh material kept 

exposed by accelerated erosion on tectonically steepened terrain. Carbon-dioxide changes related to 

plume head volcanism and uplift is consistent with the climate history in the in the 17- to 10-Ma time 

interval shown in Figure 10 as follows: (1) carbon dioxide buildup in atmosphere associated with 15- to 

17-Ma Oregon Plateau and Columbia River flood basalts, and (2) carbon-dioxide reduction associated 

with silicate weathering due to erosion and dissection of plume-head uplift starting about 15 Ma. 

0 

10 

20 

30 

Large 

Igneous 

Provinces 

Columbia River 

North Atlantic 

Deccan 

4 2 0 +3 +2 

BENTHIC FORAM. 

Expanded timescale 

5-20 Ma 

Microfossil

Paleotemperature 

Evidence 

+1 

1-5 

3-8 

5.8-12 

7-12 

1.7-13 

7.9.10 

1.8-8.12.14 

1.8-8.10.12.14 

4.6.7.10.12 

4.10.15 

4.10.15.16 

10.15.16 

10.15.17 

SITE 357 

15.17 

cold warm 

5 Ma 

10 Ma 

20 Ma 

CRB 

-1 

δ18O δ18o 

SITE 289 

SITE 158 

Yellowstone 

plume-head uplift 

(postulated) 

0 mK

Atlantic 

deep waters 

3 1 1

14 Ma 

40 

50 17 Ma 

60 

70 

5 10 15 20 25 

Paleotemperature ( °C ) 

Figure 10. Temporal relations among Cenozoic climate proxies, "large igneous provinces" (flood basalts) and postulated Yellowstone plumehead uplift. Atlantic deep waters (after Miller and others, 1987, as shown in Barron and Baldauf, 1989). The about 14-17-Ma middle Miocene 

"climatic optimum" (expanded plot, right side from Barron and Keller, 1982) has been associated with the Columbia River-Oregon Plateau flood 

basalts (CRB, for example, Hodell and Woodruff, 1994). After the Miocene "climatic optimum," cooling continued eventually to the PliocenePleistocene ice ages. Uplift associated with the Yellowstone plume head is a candidate for a "plateau uplift" whose plume-head mechanism could 

produce significant late Cenozoic uplift (fig. 9). Global climatic modeling by Kutzbach and others (1989) and Ruddiman and others, (1997) 

shows that "plateau uplifts" could produce late Cenozoic climatic cooling in the Northern Hemisphere that might lead to the Plio-Pleistocene ice 

ages. Late Cenozoic Yellowstone plume uplift (head and tail) could also diminish atmospheric CO2 leading to cooling, by the mechanism of 

Raymo and Ruddiman (1992) and Ruddiman and others (1997). Plume-head uplift does not have the geophysical problem of a plausible 

mechanism of how the isostatic balance of the lithosphere could rapidly change that Molnar and England (1990a; b) used to challenge the 

validity of late Cenozoic "plateau uplifts". Millions of years before present 

➤ ➤ 

24

DISCUSSION OF PLUME HEAD-UPLIFT AND CLIMATE

Studies by Ruddiman, Kutzbach, and colleagues suggest that late Cenozoic plateau uplift could have 

forced changes that led to the "late Cenozoic climatic deterioration" that culminated in the Plio-Pleistocene ice ages (Ruddiman and Kutzbach, 1989; Kutzbach and others, 1989; Ruddiman and others, 1989). 

They suggested plateau uplift in southern Asia and the American West, including the Sierra Nevada, 

Colorado Plateau, Basin and Range, Rocky Mountains, and High Plains (Ruddiman and others, 1989). 

However, Molnar and England (1990a, b) strongly question the late Cenozoic uplift of the American 

West. Much new paleobotanical analysis using the methods of Wolfe (1993) or similar methods has also 

cast doubt on general uplift in the West and particularly that of the Rocky Mountains (Gregory and Chase, 

1992) and the Colorado Plateau. Also, House and others (1998) conclude the Sierra Nevada has been high 

since the Cretaceous based on thermal history and apatite (U-Th)/He ages. In addition, Molnar and 

England (1990a) examine the geophysical basis of uplift and question evidence suggesting true uplift as 

distinct from isostatic uplift associated with erosion. 

The mechanism of uplift associated with a plume head or tail does provide a driving process below 

the lithosphere to explain epeirogenic-type uplift. We suggest that broad plume-head uplift (see fig. 9) 

was substantial by about 17 Ma and was followed by gradual subsidence in the head area. The uplift 

migrated northeastward to the present Yellowstone crescent of high terrain (fig. 9) that is associated with 

the current plume tail (Pierce and Morgan, 1992). The geoid anomaly that now centers on the Yellowstone 

Plateau (Pierce and Morgan, 1992; Pierce and others, 1992; Smith and Braille, 1993, fig. 7) and, more 

importantly, the larger plume-head uplift might provide a late Cenozoic broad uplift that could affect 

climate patterns in the ways modeled by Kutzbach and others (1989). Evidence (table 1) for uplift, 

migrating drainage divides, and an orographic precipitation shadow suggest uplift associated with the 

plume head and ensuing tail may have occurred. Much more complete analysis of existing and new 

evidence is needed to evaluate plume-head uplift and possible relations to climatic patterns, but we hope 

this brief outline suggests avenues for future research. 

CONCLUSIONS

We were led to the hypothesis of a 17-Ma Yellowstone plume head by tracing the Yellowstone 

hotspot track back to the area of the McDermitt volcanic field on the Nevada-Oregon border. We think the 

Yellowstone hotspot track is best explained by a mantle plume particularly because of the following 

evidence (Pierce and Morgan, 1992): (1) the volcanism that progressed from 10 Ma to 2 Ma is coincident 

with both the rate and orientation predicted by plate motion, (2) the belts of faulting and uplift are oriented about this volcanic progression like the bow wave of a boat with uplift occurring in advance (northeast) of volcanism, (3) a large geoid anomaly centers on Yellowstone, (4) 3

He/4He ratios near 16 suggest a 

deep mantle source, and, most important, (5) the scale of associated faulting and uplift is more than 400 

km across suggesting a deep, sub-lithospheric process. 

Upon backtracking the hotspot to its start near McDermitt, we find that a plume head (a sphere about 

400 km in diameter) explains the following observations (fig. 1): (1) the much wider dispersal of 14- to 

16-Ma rhyolitic volcanism than that after 10 Ma, (2) the association of the inferred plume head with the 

14- to 17-Ma Columbia River and Oregon Plateau basalts, (3) the plume-head location near the mid-point 

of the 1,100-km-long, 16- to 17-Ma, Nevada-Oregon rift zone, and (4) no continuous manifestation of a 

hotspot before 17 ma. 

If one accepts the Yellowstone plume head hypothesis, the following three topics merit attention. 

1. Westward displacement of the plume head upward along the inclined Vancouver slab (figs. 2 and 3). 

Plate tectonic reconstructions suggest the rising plume head rose upward into the Vancouver slab. We 

25

suggest its buoyant rise was deflected westward by the up-to-the-west rise of the Vancouver slab. This 

could explain an unresolved problem with the hotspot track between the 10-Ma Picabo volcanic field and 

the 16-Ma McDermitt area that calls for a rate too high (70 km/m.y) and a trend more easterly (about 20° 

clockwise) than indicated by both the post-10-Ma progression of volcanism and faulting (about 2.9 ± 0.5 

km/m.y to the N 54 ± 5° E) and the known rate and direction of North American plate motion (about 2.2 

± 0.8 km/m.y. to the S 56 ± 17° W; Pierce and Morgan, 1992, p. 6). 

2. The oblique transection of the craton margin by the plume head going from Mesozoic accreted oceanic 

lithosphere in the north to Precambrian craton in the south (figs. 5 and 6). Our best estimate is that the 

plume head contacted the lithosphere about 17 Ma near the Oregon-Nevada border beneath the 

McDermitt volcanic field. Reconstructions of crust and mantle lithosphere properties suggest a thinner, 

denser, oceanic lithosphere was present to the north in Oregon and Washington than in Nevada. This 

permitted the mantle plume head to migrate preferentially towards this "thinspot" and thus favor more 

decompression melting. The greater crustal density favored the eruption of the Columbia River and 

Oregon Plateau flood basalts. The 16-17-Ma Nevada-Oregon rift zone is defined by the N. 20° W. orientation of the following about 16-17 Ma features: (1) the northern Nevada rift, (2) the extension of this rift 

into southern Nevada, and (3) the feeder dikes to the Columbia River and Oregon Plateau flood basalts. 

Although flood basalts are north of the inferred plume center, the plume is at the midpoint of this 16-Ma 

rift zone. Composition of erupted or intruded magmatic material can be explained by differences in 

crustal composition: (1) flood basalts are restricted to the oceanic, mafic crust and (2) rhyolites are 

restricted to more continental crust where the partial melt of mantle basalt melted silicic crustal material 

which rose upward to form magma chambers that erupted to produce rhyolite flows and ignimbrites, 

leaving the heat-supplying basalt at depth. 

3. Plume-head uplift and associated climatic patterns (figs. 8 and 9). A plume head is expected to produce 

significant uplift. Based on several observations, we prefer roughly 1 km of uplift over an oblong northsouth ellipse of 1,100 km. Paleobotanical analysis of leaves suggest that at 15 Ma, central and northern 

Nevada was about 1 km higher than present. A precipitation shadow is expected to occur east of this 

postulated uplift; evidence of marked aridity about 15 Ma is found near the common boundaries of 

Montana, Wyoming, and Idaho. About 7 Ma, uplift associated with eastward hotspot migration is predicted to migrate to the area of the central SRP; paleo-mollusk studies suggest a drainage divide in this 

area near American Falls, Idaho, and other studies suggest aridity to the east of this uplift. The 

Yellowstone plume-head may provide a mechanism for regional uplift that is geophysically plausible. 

Notably, the plume-head mechanism can cause geologically rapid uplift without requiring crustal thickening. Modeling studies suggest this kind of uplift could contribute to late Cenozoic cooling leading to the 

ice ages (Kutzbach and others, 1989; Ruddiman and others, 1997). 

A more thorough evaluation of the plume head hypothesis and its implications regarding (1) westward offset up the inclined Vancouver slab; (2) lithospheric changes along the Nevada-Oregon rift zone, 

and (3) uplift and associated climatic changes all require much additional study. 

ACKNOWLEDGMENTS

We have discussed these ideas with many colleagues during the last few years. We have particularly benefited 

from discussions with George Thompson, Tom Parsons, Mike Coffin, Norm Sleep, Robert Duncan, Eugene 

Humphreys, Vic Camp, Alan Wallace, Anita Grunder, Scott Lundstrom, Margaret Hiza, Silvio Pezzopane, and Bob 

Thompson. We wish to thank Spence Wood, Tom Parsons, and Mike McCurry for technical review of an earlier 

version of this paper, and Roger Stewart for a thorough and quite helpful editorial review. 

26

REFERENCES CITED

Amundson, Ronald, Oliver Chadwick, Carol Kendall, Yang Wang, and Michael DeNiro, 1996, Isotopic 

evidence for shifts in atmospheric circulation patterns during the late Quaternary in mid-North 

America: Geology, v. 24, p. 23-36. 

Anders, M.H., 1994, Constraints on North American plate velocity from Yellowstone hotspot deformation 

field: Nature, v. 369, p. 53-55. 

Anders, M.H., J.W. Geissman, L.A. Piety, and J.T. Sullivan, 1989, Parabolic distribution of circumeastern Snake River Plain seismicity and latest Quaternary faulting: Migratory pattern and association with the Yellowstone hot spot: Journal of Geophysical Research, v. 94. No. B2, p. 1,589-1,621. 

Anderson, D.L., 1998, The edges of the mantle, in Michael Gurnis, M.E. Wysession, Elise Kittle, and 

B.A. Buffett, eds., The Core-Mantle Boundary Region, American Geophysical Union, Geodynamics 

Series, v. 28, p. 255-271. 

Arthur, M.A., W.E. Dean, and S.O. Schlanger, 1985, Variations in the global carbon cycle during the 

Cretaceous related to climate, volcanism, and changes in atmospheric CO2

, in E.T. Sundquist and 

W.H. Broecker, eds., The Carbon Cycle and Atmospheric CO2: Natural Variations Archean to 

Present, American Geophysical Union Geophysical Monograph 32, p. 504-529. 

Atwater, Tanya, 1989, Plate tectonic history of the northeast Pacific and western North America, in E.L. 

Winterer, D.M. Hussong, and R.W. Decker, eds., The Eastern Pacific Ocean and Hawaii: The 

Geology of North America, Geological Society of America, v. N, p. 21-72. 

Baksi, A.K., 1989, Reevaluation of the timing and duration of extrusion of the Imnaha, Picture Gorge, and 

Grande Ronde Basalts, Columbia River Group, in S.P. Reidel, and P.R. Hooper, Volcanism and 

Tectonism in the Columbia River Flood Basalt Province: Geological Society of America Special 

Paper 239, p. 105-111. 

Barnosky, A.D., and W.J. Labar, 1989, Mid-Miocene (Barstovian) environmental and tectonic setting near 

Yellowstone National Park, Wyoming and Montana: Geological Society of America Bulletin, v. 101, 

p. 1,448-1,456. 

Barron, J.A., and J.G. Baldauf, 1989, Tertiary cooling steps and paleoproductivity as reflected by diatoms 

and biosiliceous sediments, in Productivity of the Ocean: Present and Past, W.H. Berger, V.S. 

Smetacck, and G. Wefer, eds., John Wiley and Sons, . p. 341-354. 

Barron, J.A., and Gerta Keller, 1982, Widespread Miocene deep-sea hiatuses: Coincidence with periods of 

global cooling: Geology, v. 10, p. 577-581. 

Blackwell, D.D., 1989, Regional implications of heat flow of the Snake River Plain, northwestern United 

States: Tectonophysics, v. 164, p. 323-343. 

Blakely, R.J., and R.C. Jachens, 1991, Regional study of mineral resources in Nevada: Insights from 

three-dimensional analysis of gravity and magnetic anomalies: Geological Society of America 

Bulletin, v. 103, p. 795-803. 

Bonnichsen, Bill, 1982, The Bruneau-Jarbidge eruptive center, southwestern Idaho, in Bill Bonnichsen, 

and R.M. Breckenridge, eds., Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology 

Bulletin 26, p. 237-254. 

Braile, L.W., W.J. Hinze, R.R.B. von Frese, and G.R. Keller, 1989, Seismic properties of the crust and 

uppermost mantle of the conterminous United States and adjacent Canada, in L.C. Pakiser, and W.D. 

Mooney, Geophysical Framework of the Continental United States: Geological Society of America 

Memoir 172, p. 655-680. 

27

Camp, V.E., 1995, Mid-Miocene propagation of the Yellowstone mantle plume head beneath the Columbia River basalt source region: Geology, v. 23, no. 5, p. 435-438. 

Campbell, I.H., 1998, The mantle's chemical structure: Insights from melting products of mantle plumes, 

in Ian Jackson, ed., The Earth's Mantle: Cambridge University Press, Cambridge, p. 259-310. 

Carlson, R.W., and W.K. Hart, 1987, Crustal genesis of the Oregon Plateau: Journal of Geophysical 

Research, v. 92, p. 6,191-6,206. 

______ 1988, Flood basalt volcanism in the northwestern United States, in J.D. Macdougall, ed., Continental Flood Basalts: Kluwer Academic Publishers, Dordrecht, The Netherlands, p. 35-61. 

Christiansen, R.L., 1984, Yellowstone magmatic evolution: Its bearing on understanding large-volume 

explosive volcanism, in Explosive Volcanism: Inception, Evolution, and Hazards, Studies in Geophysics: National Academy Press, p. 84-95. 

Christiansen, R.L., and E.H. McKee, 1978, Late Cenozoic volcanic and tectonic evolution of the Great 

Basin and Columbia intermontane regions, in R.B. Smith, and G.P. Eaton, eds., Cenozoic Tectonics 

and Regional Geophysics of the Western Cordillera: Geological Society of America Memoir 152, p. 

283-311. 

Christiansen, R.L, and R.S. Yeats, 1992, Post-Laramide geology of the U.S. Cordilleran region, in B.C. 

Burchfiel, P.W. Lipman, and M.L. Zoback, eds., The Cordilleran Orogen: Conterminous U.S.: The 

Geology of North America: Geological Society of America, v. G-3, p. 261-406. 

Coffin, M.F., and Olav Eldholm, 1994, Large igneous provinces: Crustal structure, dimensions, and 

external consequences: Reviews of Geophysics, v. 32, p. 1-36. 

Covington, H.R., 1983, Structural evolution of the Raft River basin, Idaho, in D.M. Miller, V.R. Todd, and 

K.A. Howard, eds., Tectonic and Stratigraphic Studies in the Eastern Great Basin: Geological 

Society of America Memoir 157, p. 229-237. 

Crough, S.T., 1978, Thermal origin of mid-plate hot-spot swells: Geophysical Journal of the Royal 

Astronomical Society, v. 55, p. 451-469. 

Davies, G.F., 1998, Plates, plumes, mantle convection, and mantle evolution, in Ian Jackson, ed., The 

Earth's Mantle: Cambridge University Press, Cambridge, p. 228-258. 

Draper, D.S., 1991, Late Cenozoic bimodal magmatism in the northern Basin and Range Province of 

southeastern Oregon: Journal of Volcanology and Geothermal Research, v. 47, p. 299-328. 

Drummond, C.N., B.H. Wilkinson, K.C. Lohmann, and G.R. Smith, 1993, Effect of regional topography 

and hydrology on the lacustrine isotopic record of Miocene paleoclimate in the Rocky Mountains: 

Palaeogeography, Palaeoclimatology, Palaeoecology, v. 101, p. 67-79. 

Duncan, R.A., 1982, A captured island chain in the Coast Range of Oregon and Washington: Journal of 

Geophysical Research, v. 87, p. 10,827-10,837. 

Duncan, R.A., and M.A. Richards, 1991, Hotspots, mantle plumes, flood basalts, and true polar wander: 

Reviews in Geophysics, v. 29, p. 31-50. 

Elison, M.W., R.C. Speed, and R.W. Kistler, 1990, Geologic and isotopic constraints on the crustal 

structure of the northern Great Basin: Geological Society of America Bulletin, v. 102, p. 1,077-1,092. 

Fields, R.W., D.L. Rasmussen, A.R. Tabrum, and R. Nichol, 1985, Cenozoic rocks of the intermontane 

basins of western Montana and eastern Idaho, in R.M. Flores and S.S. Kaplan, eds., Cenozoic 

Paleogeography of the West-Central United States: Society of Economic Paleontologists and Mineralogists, Rocky Mountain Section, p. 9-36. 

28

Friedman, Irving, Jim Glean, R.A. Shepherd, and A.J. Guide, III, 1993, Deuterium fractionation as water 

diffuses into silicic volcanic ash: Geophysical Monograph 78, p. 321-323. 

Garnero, E.J., Justin Revenaugh, Quentin Williams, Thorne Lay, and L.H. Kellogg, 1998, Ultralow 

velocity zone at the core-mantle boundary in Michael Gurnis, M.E. Wysession, Elise Kittle, and B.A. 

Buffett, eds., the Core-mantle boundary region: American Geophysical Union, Geodynamics Series, 

v. 28, p. 319-334 

Geist, Dennis, and Mark Richards, 1993, Origin of the Columbia Plateau and Snake River Plain: Deflection of the Yellowstone plume: Geology, v. 21, p. 789-792. 

Gregory, K.M., and C.G. Chase, 1992, Tectonic significance of paleobotanically estimated climate and 

altitude of the late Eocene erosion surface, Colorado: Geology, v. 20, p. 581-585. 

Griffiths, R.W., and J.S. Turner, 1998, Understanding mantle dynamics through mathematical models and 

laboratory experiments, in Ian Jackson, ed., The Earth's Mantle: Cambridge University Press, 

Cambridge, p. 191-227. 

Hill, D.P., 1972, Crustal and upper-mantle structure of the Columbia Plateau from long-range seismicrefraction measurements: Geological Society of America Bulletin, v. 83, p. 1639-1648. 

Hill, R.I., I.H. Campbell, G.F. Davies, and R.W. Griffiths, 1992, Mantle plumes and continental tectonics: 

Science, v. 256, p. 186-193. 

Hittleman, A.M., D.T. Dater, R.W. Buhmann, and S.D. Racey, 1994, Gravity-1994 edition (CD-ROM): 

National Geophysical Data Center, Boulder, CO (request flier SE-0703). 

Hodell, D.A., and Fay Woodruff, 1994, Variations in the strontium isotopic ratio of seawater during the 

Miocene: Stratigrapic and geochemical implications: Paleoceanography, v. 9, p. 405-426. 

Hooper, P.R., 1990, The timing of crustal extension and the eruption of continental flood basalts: Nature, 

v. 345, p. 246-249. 

______ 1997, The Columbia River flood basalt province: Current status, in J.J. Mahoney and M.F. Coffin, 

eds., Large Igneous Provinces: Continental, Oceanic, and Planetary Flood Volcanism: American 

Geophysical Monograph 100, p. 1-27. 

Hooper, P.R., and C.J. Hawkesworth, 1993, Isotopic and geochemical constraints on the origin and 

evolution of the Columbia River basalt: Is there a Ba/La ratio argument for the Vancouver slab?: 

Journal of Petrology, v. 34, p. 1,203-1,246. 

House, M.A., B.P. Warnicke, and K.A., Farley, 1998, Dating topography of the Sierra Nevada, California, 

using apatite (U/Th)/He ages: Nature, v. 396, p. 66-69. 

Jackson, Ian, and S.M. Rigden, 1998, Composition and temperature of the earth's mantle: Seismological 

models interpreted through experimental studies of earth materials, in Ian Jackson, ed., The Earth's 

Mantle: Cambridge University Press, Cambridge, p. 405- 460. 

Johnston, S.T., P.J. Wynne, Don Francis, C.J.R. Hart, R.J. Enkin, and D.C. Engebretson, 1996, 

Yellowstone in Yukon: The late Cretaceous Carmaks Group: Geology, v. 24, p. 997-1,000. 

Katzman, Rafael, Li Zhao, and T.H. Jordan, 1998, High-resolution, two dimensional vertical tomography 

of the central Pacific mantle using ScS reverberations and frequency-dependent travel times: Journal 

of Geophysical Research, v. 103, p. 17,933-17,971. 

Kistler, R.W., 1983, Isotope geochemistry of plutons in the northern Great Basin: Geothermal Resources 

Council, Davis, California, Special Report 13, p. 3-8. 

29

Kutzbach, J.E., P.J. Guetter, W.F. Ruddiman, and W.L. Prell, 1989, Sensitivity of climate to late Cenozoic 

uplift in southern Asia and the American West: Numerical experiments: Journal of Geophysical 

Research, v. 94, p. 18,393-18,407. 

Lachenbruch, A.H., and J.H. Sass, 1978, Models of an extending lithosphere and heat flow in the Basin 

and Range province: Geological Society of America Memoir 152, p. 209-250. 

Leavitt, S.W., 1982, Annual volcanic carbon dioxide emission: an estimate from eruption chronologies: 

Environmental Geology, v. 4, p. 15-21. 

Leeman, W.P., 1982, Development of the Snake River Plain-Yellowstone Plateau province, Idaho and 

Wyoming: An overview and petrologic model, in Bill Bonnichsen, and R.M. Breckenridge, eds., 

Cenozoic Geology of Idaho: Idaho Bureau of Mines and Geology Bulletin 26, p. 155-178. 

Leeman, W.P., J.S. Oldow, and W.K. Hart, 1992, Lithosphere-scale thrusting in the western U.S. Cordillera as constrained by Sr and Nd isotopic transitions in Neogene volcanic rocks: Geology, v. 20, p. 

63-66. 

Leopold, E.B., and M.F. Denton, 1987, Comparative age of grassland and steppe east and west of the 

northern Rocky Mountains: Annals of the Missouri Botanical Garden, v. 74, p. 841-867. 

Leopold, E.B., and V.C. Wright, 1985, Pollen profiles of the Plio-Pleistocene transition in the Snake River 

Plain, Idaho: in Late Cenozoic History of the Pacific Northwest, American Association for the 

Advancement of Science, Pacific Division, San Francisco, p. 323-348. 

Link, P.K., N. Christie-Blick, J.H. Stewart, J.M.G. Miller, W.J. Devlin, and M. Levy, 1993, Late Proterozoic strata of the United States Cordillera in P.K. Link and others, Middle and Late Proterozoic 

Stratified Rocks of the Western U.S. Cordillera, Colorado Plateau, and Basin and Range Province, in 

J.C. Reed, M.E. Bickford, R.S. Houston, P.K. Link, D.W. Rankin, P.K. Sims, and W.R. Van Schmus, 

eds., PreCambrian: Conterminous U.S.: Geological Society of America, The Geology of North 

America, Boulder, v. C-2, p. 536-558. 

Lipman, P.W., 1992, Magmatism in the Cordilleran United States; Progress and problems: in B.C. 

Burchfiel, P.W. Lipman, and M.L. Zoback, eds., The Cordilleran Orogen: Conterminous U.S.: The 

Geology of North America: Geological Society of America, v. G-3, p. 481-514. 

Love, J.D., L.A. Morgan, and W.C. McIntosh, 1997, The Teewinot Formation: Evidence for Late Miocene 

basin formation in response to volcanism, faulting, and uplift associated with the Yellowstone 

hotspot?: Geological Society of America Abstracts with Programs, v. 29, no. 6, p. A-365. 

Luedke, R.G., and R.L. Smith, 1981, Map showing distribution and age of late Cenozoic volcanic centers 

in California and Nevada: U.S. Geological Survey Miscellaneous Geologic Investigations Map, I1091-C, scale 1:1,000,000. 

______ 1982, Map showing distribution and age of late Cenozoic volcanic centers in Oregon and Washington: U.S. Geological Survey Miscellaneous Geologic Investigations Map, I-1091-D, scale 

1:1,000,000. 

______ 1983, Map showing distribution and age of late Cenozoic volcanic centers in Idaho, western 

Montana, west-central South Dakota, and northwestern Wyoming: U.S. Geological Survey Miscellaneous Geologic Investigations Map, I-1091-E, scale 1:1,000,000. 

MacLeod, N.S., G.W. Walker, and E.H. McKee, 1976, Geothermal significance of eastward increase in 

age of upper Cenozoic rhyolitic domes in southeastern Oregon: Proceedings of the Second United 

Nations Symposium on the Development and Use of Geothermal Resources, v. 1, Washington, D.C., 

U.S. Government Printing Office, p. 456-474. 

30

Malde, H.E., 1991, Quaternary geology and structural history of the Snake River Plain, Idaho and Oregon, in R.B. Morrison, ed., Quaternary Nonglacial History of the Conterminous U.S.: Geological 

Society of America, The Geology of North America, v. K-2, ch. 4, p. 251-282. 

McClean, D.M., 1985, Mantle degassing induced dead ocean in the Cretaceous-Tertiary transition, in E.T. 

Sundquist and W.H. Broecker, eds., The Carbon Cycle and Atmospheric CO2

: Natural Variations 

Archean to Present, American Geophysical Union Geophysical Monograph 32, p. 493-503. 

Miller, K., and R.G. Fairbanks, and G.S. Mountain, 1987, Tertiary isotope synthesis, sea level history, and 

continental margin erosion: Paleooceanography, v. 2, p. 1-19. 

Molnar, Peter, and Philip England, 1990a, Surface uplift, uplift of rocks, and exhumation of rocks: 

Geology, v. 18, p. 1,173-1,177. 

Molnar, Peter, and Philip England, 1990b, Late Cenozoic uplift of mountain ranges and global climate 

change: Chicken or egg?: Nature, v. 346, p. 29-34. 

Mooney, W.D., and L.W. Braile, 1989, The seismic structure of the continental crust and upper mantle of 

North America, in A.W. Bally, and A.R. Palmer, eds., The Geology of North America—An Overview: Geological Society of America, The Geology of North America, v. A, p. 39-52. 

Mooney, W.D, and C.S. Weaver, 1989, Regional crustal structure and tectonics of the Pacific coastal 

states; California, Oregon, and Washington, in L.C. Pakiser, and W.D. Mooney, eds., Geophysical 

Framework of the Continental United States: Geological Society of America Memoir 172, p. 129-

161. 

Morgan, J.P., W.J. Morgan, and Evelyn Price, 1995, Hotspot melting generates both hotspot volcanism 

and a hotspot swell?: Journal of Geophysical Research, v. 100, p. 8045-8062. 

Morgan, L.A. and W.C. McIntosh, 2001, 40Ar/39Ar ages of silicic volcanic rocks in the Heise volcanic 

field, eastern Snake River Plain, Idaho: Timing of volcanism and tectonism, in Bill Bonnichsen, 

C.M. White, and Michael McCurry, eds., Tectonic and Magmatic Evolution of the Snake River Plain 

Volcanic Province: Idaho Geological Survey Bulletin. 

Morgan, L.A., D.J. Doherty, and W.P. Leeman, 1984, Ignimbrites of the eastern Snake River Plain: 

Evidence for major caldera forming eruptions: Journal of Geophysical Research, v. 89, no. B10, p. 

8,665-8,678. 

Morgan, L.A., W.C. McIntosh, and K.L. Pierce, 1997, Inferences for changes in plume dynamics from 

stratigraphic framework studies of ignimbrites, central Snake River Plain, Idaho: Geological Society 

of America Abstracts with Programs, v. 29, no. 6, p. A-299. 

Morgan, Paul, and W.D. Gosnold, 1989, Heat flow and thermal regimes of the continental United States, 

in L.C. Pakiser, and W.D. Mooney, eds., Geophysical Framework of the Continental United States: 

Geological Society of America Memoir 172, p. 493-522. 

Murphy, J.B., G.L. Oppliger, and G.H. Brimhall, Jr., 1998, Plume-modified orogeny: An example from 

the western United States: Geology, v. 26, p. 731-734. 

Oppliger, G.L., J.B Murphy, and G.H. Brimhall, Jr., 1997, Is the ancestral Yellowstone hotspot responsible for the Tertiary Carlin mineralization in the Great Basin of Nevada?: Geology, v. 25, no. 7, p. 

627-630. 

Page, W.D., T.L. Sawyer, and P, R. Renne, 1995, Quaternary geology along the boundary between the 

Modoc Plateau, southern Cascades, and the northern Sierra; Friends of the Pleistocene, Pacific Cell 

Field Trip. 

31

Pakiser, L.C., 1989, Geophysics of the Intermontane system, in L.C. Pakiser, and W.D. Mooney, eds., 

Geophysical Framework of the Continental United States: Geological Society of America Memoir 

172, p. 235-247. 

Parsons, Tom, 1995, The Basin and Range Province, in K.H. Olsen, ed., Continental Rifts: Evolution, 

Structure, Tectonics, Developments: Geotectonics 25, Elsevier, Amsterdam, p. 277-324. 

Parsons, Tom, G.A. Thompson, and N.H. Sleep, 1994, Mantle plume influence on the Neogene uplift and 

extension of the U.S. western Cordillera?: Geology, v. 22, p. 83-86. 

Parsons, Tom, A.M. Trehu, J.H. Luetgert, Kate Miller, Fiona Kilbride, R.E. Wells, M.A. Fisher, Ernst 

Flueh, U.S. Ten Brink, and N.I. Christensen, 1998, A new view into the Cascade subduction zone and 

volcanic arc: Implications for earthquake hazards along the Washington margin: Geology, v. 26, p. 

199-202. 

Perkins, M.E., F.H. Brown, W.P. Nash, W. McIntosh, and S.K. Williams, 1998, Sequence, age, and source 

of silicic fallout tuffs in middle to late Miocene basins of the Basin and Range province: Geological 

Society of America Bulletin, v. 110, p. 344-360. 

Pezzopane, S.K., and R.J. Weldon, II, 1993, Tectonic role of active faulting in central Oregon: Tectonics, 

v. 12, p. 1,140-1,169. 

Pierce, K.L., and L.A. Morgan, 1990, The track of the Yellowstone hot spot: Volcanism, faulting, and 

uplift: U.S. Geological Survey Open-File Report 90-415, 49 p. 

______ 1992, The track of the Yellowstone hot spot: Volcanism, faulting, and uplift, in P.K. Link, M.A. 

Kuntz, and L.B. Platt, eds., Regional Geology of Eastern Idaho and Western Wyoming: Geological 

Society of America Memoir 179, p. 1-53. 

Pierce, K.L., D.G. Milbert, and R.W. Saltus, 1992, Geoid dome culminates on Yellowstone: Yellowstone 

hotspot fed by a thermal mantle plume?: EOS, Transactions of the American Geophysical Union, v. 

73, p. 284. 

Pyle, D.G., B.B. Hanan, D.W. Graham, and R.A. Duncan, 1997, Siletzia-Geochemistry and geochronology of Yellowstone hot spot volcanism in a suboceanic setting: Geological Society of America 

Abstracts with Programs, v. 29, no. 6, p. A-298. 

Rasmussen, J., and E. Humphreys, 1988, Tomographic image of the Juan de Fuca plate beneath Washington and western Oregon using teleseismic P-wave travel times: Geophysical Research Letters, v. 15, 

p. 1,417-1,420. 

Rampino, M.R., 1991, Volcanism, climate change, and the geologic record, Society of Economic Paleontologists and Mineralogists, Special Publication 45, p. 

Raymo, M.W., and W.F. Ruddiman, 1992, Tectonic forcing of late Cenozoic climate: Nature, v. 359, p. 

117-122. 

Raymo, M.W., W.F. Ruddiman, and P.N. Froelich, 1988, The influence of late Cenozoic mountain building on oceanic geochemical cycles, Geology, v. 16, p. 649-653. 

Reed., J.C., Jr., 1993, Map of the PreCambrian rocks of the conterminous United States, in J.C. Reed, 

M.E. Bickford, R.S. Houston, P.K. Link, D.W. Rankin, P.K. Sims, and W.R. Van Schmus, eds., 

Precambrian: Conterminous U.S.: Geological Society of America, The Geology of North America, 

Boulder, v. C-2, Plate 1. 

Reidel, S.P, N.P. Campbell, K.R. Fecht, and K.A. Lindwey, 1989, Late Cenozoic structure and stratigraphy of south central Washington, in Regional Geology of Washington State: Washington Department 

of Natural Resources Bulletin 80, p. 159-180. 

32 
Richards, M.A., R.A. Duncan, and V.E. Courtillot, 1989, Flood basalts and hot spot tracks: Plume heads 

and tails: Science, v. 246, p. 103-107. 

Rodgers, D.W., W.R. Hackett, and H.T. Ore, 1990, Extension of the Yellowstone Plateau, eastern Snake 

River Plain, and Owyhee Plateau: Geology, v. 18, p. 1,138-1,141. 

Ruddiman, W.F., and J.E. Kutzbach, 1989, Forcing of late Cenozoic Northern Hemisphere climate by 

plateau uplift in southern Asia and the American West: Journal of Geophysical Research, v. 94, p. 

18,409-18,427. 

Ruddiman, W.F., J.E. Kutzbach, and I.C. Prentice, 1997, Testing the climatic effects of orography and 

CO2

 with general circulation and Biome models, in W.F. Ruddiman, ed., Tectonic Uplift and Climate 

Change, Plenum Press, New York, p. 204-235. 

Ruddiman, W.F., W.L. Prell, and M.E. Raymo, 1989, Late Cenozoic uplift in southern Asia and the 

American West: Rationale for general circulation modeling experiments: Journal of Geophysical 

Research, v. 94, p. 18,379-18,391. 

Ruddiman, W.F., M.E. Raymo, W.L. Prell, and J.E. Kutzbach., 1997, The uplift-climate connection: a 

synthesis, in W.F. Ruddiman, ed., Tectonic Uplift and Climate Change, Plenum Press, New York, p. 

471-515. 

Rytuba, J.J. and E.H. McKee, 1984, Peralkaline ash flow tuffs and calderas of the McDermitt volcanic 

field, southeast Oregon and north central Nevada: Journal of Geophysical Research, v. 89, no. B10, 

p. 8,616-8,628. 

Saltus, R.W., and G.A. Thompson, 1995, Why is it downhill from Tonopah to Las Vegas?: A case for a 

mantle plume support of the high northern Basin and Range: Tectonics, v. 14, p. 1,235-1,244. 

Self, Stephen, Thorvaldur Thordarson, and Laszlo Keszthelyi, 1997, Emplacement of continental flood 

basalt lava flows, in J.J. Mahoney and M.F. Coffin, eds., Large Igneous Provinces: Continental, 

Oceanic, and Planetary Flood Volcanism: American Geophysical Monograph 100, p. 381-410. 

Severinghaus, Jeff, and Tanya Atwater, 1990, Cenozoic geometry and thermal state of the subducting 

slabs beneath western North America, in B.P. Wernicke, ed., Basin and Range Extensional Tectonics 

Near the Latitude of Las Vegas, Nevada: Geological Society of America Memoir 176, p. 1-22. 

Shervais, John, Scott Vetter, and Barry Hanan, 1997, Shaking the plume's tail: Basaltic volcanism in the 

central Snake River Plain, Idaho: Geological Society of America Abstracts with Programs, v. 29, no. 

6, p. A-300. 

Sleep, N.H., 1990, Hotspots and mantle plumes: Some phenomenology: Journal of Geophysical Research, 

v. 95, no. B5, p. 6,715-6,736. 

______ 1997, Lateral flow and ponding of starting plume material: Journal of Geophysical Research, v. 

102, no. B5, p. 10,001-10,012. 

Smith, G.R., and W.P. Patterson, 1994, Mio-Pliocene seasonality on the Snake River Plain: Comparison 

of faunal and oxygen isotopic evidence: Palaeogeography, Palaeoclimatology, Palaeoecology, v. 107, 

p. 291-302. 

Smith, R.B., and L.W. Braille, 1993, Topographic signature, space-time evolution, and physical properties 

of the Yellowstone-Snake River Plain volcanic system: The Yellowstone hotspot, in A.W. Snoke, J.R. 

Steidtmann, and S.M. Roberts, eds., Geology of Wyoming: Geological Survey of Wyoming Memoir 

No. 5, p. 694-754. 

Snee, L.W., Karen Lund, J.F. Sutter, D.E. Balcer, and K.V. Evans, 1995, An 40Ar/39Ar chronicle of the 

33

tectonic development of the Salmon River suture zone, western Idaho, in T.L. Vallier and H.C. 

Brooks, eds., Geology of the Blue Mountains Region of Oregon, Idaho, and Washington: Petrology 

and Tectonic evolution of Pre-Tertiary Rocks of the Blue Mountain Region, U.S. Geological Survey 

Professional Paper 1438, p. 359-414. 

Takahahshi, Eiichi, Katsuji Nakajima, and T.L. Wright, 1998, Origin of the Columbia River basalts; 

Melting model of a heterogeneous plume head: Earth and Planetary Science letters, v. 162 (1-4), pp. 

63-80. 

Taylor, D.W., and R.C. Bright, 1987, Drainage history of the Bonneville Basin, in Cenozoic Geology of 

Western Utah: Utah Geological Association Publication 16, p. 239-256. 

Thelin, Gail, and R.J. Pike, 1991, Landforms of the conterminous United States-a digital shaded relief 

portrayal: U.S. Geological Survey Miscellaneous Geologic Investigations Map I-2206, scale 

1:3,500,000. 

Thompson, G.A., 1998, Deep mantle plumes and geoscience vision, Geological Society of America, and 

presidential address: GSA Today, April 1996, p. 17-25. 

Thompson, G.A., R. Catchings, E. Goodwin, S. Holbrook, C. Jarchow, C. Mann, J. McCarthy, and D. 

Okaya, 1989, Geophysics of the western Basin and Range Province, in L.C. Pakiser, and W.D. 

Mooney, eds., Geophysical Framework of the Continental United States: Geological Society of 

America Memoir 172, p. 177-203. 

Thompson, G.R., R.W. Fields, and David Alt, 1982, Land-based evidence for Tertiary climatic variations: 

Northern Rockies: Geology, v. 10, p. 413-417. 

Thompson, R.N., and S.A. Gibson, 1991, Subcontinental mantle plumes, hotspots, and pre-existing 

thinspots: Journal of the Geological Society, London, v. 148, p. 973-977. 

Vallier, T.L., 1995, Petrology of pre-Tertiary igneous rocks in the Blue Mountains province of Oregon, 

Idaho, and Washington: Implications for the geologic evolution of a complex island arc: U.S. Geological Survey Professional Paper 1438, p. 125-209. 

Van der Lee, Susan, and Guust Nolet, 1997, Seismic image of the subducted trailing fragments of the 

Vancouver plate: Nature, v. 386, p. 266-269. 

VanDecar, J.C., 1991, Upper-mantle structure of the Cascadia Subduction Zone from non-linear 

teleseismic travel-time inversion: University of Washington, Ph.D. dissertation. 

Wagner, D.L., and G.J. Saucedo, 1990, Age and stratigraphic relationships of Miocene volcanic rocks 

along the eastern margin of the Sacramento Valley, California, in R.V. Ingersoll and T.H. Nilsen, 

eds., Sacramento Valley Symposium and Guidebook, Pacific Section S.E.P.M., v. 65, p. 143-151. 

Walker, G.W., and Bruce Nolf, 1981, High lava plains, Brothers fault zone to Harvey Basin, Oregon, in 

D.A. Johnston, and Julie Donnelly-Nolan, eds., Guides to Some Volcanic Terranes in Washington, 

Idaho, Oregon, and Northern California: U.S. Geological Survey Circular 838, , p. 105-111. 

White, R.S. and D.P. McKenzie, 1989, Magmatism at rift zones: The generation of volcanic continental 

margins and flood basalts: Journal of Geophysical Research, v. 94, no. B6, p. 7,685-7,729. 

Wolfe, J.A., 1993, A method for obtaining climate parameters from leaf assemblages: U.S. Geological 

Survey Bulletin 2040, 71 p. 

Wolfe, J.A., C.E. Forest, Peter Molnar, 1998, Paleobotanical evidence on Eocene and Oligocene 

paleoaltitudes in mid-latitude western North America: Geological Society of America Bulletin, v. 

110, p. 664-678. 

34

Wolfe, J.A., H.E. Schorn, C.E. Forest, and Peter Molnar, 1997, Paleobotanical evidence for high altitudes 

in Nevada during the Miocene: Science, v. 276, p. 1,672-1,675. 

Wood, S.H., and Clemens, D.M., 2001, Geologic and tectonic history of the western Snake River Plain, 

Idaho and Oregon, in Bonnichsen, Bill, ed., Tectonic and Magmatic Evolution of the Snake River 

Plain Volcanic Province: Idaho Geological Survey Bulletin, v. 30. 

Wright, J.D., and K.G. Miller, 1992, Miocene stable isotope stratigraphy, site 747, Kerguelen Plateau: 

Proceedings Ocean Drilling Program, Scientific Results, v. 120, p. 855-866. 

Zoback, M.L., and G.A. Thompson, 1978, Basin and Range rifting in northern Nevada: Clues from a midMiocene rift and its subsequent offsets: Geology, v. 6, p. 111-116. 

Zoback, M.L., R.E. Anderson, and G.A. Thompson, 1981, Cainozoic evolution of the state of stress and 

style of tectonism of the Basin and Range Province of the western United States: Philosophical 

Transactions of the Royal Society of London, v. A 300, p. 407-434. 

Zoback, M.L., E.H. McKee, R.J. Blakely, and G.A. Thompson, 1994, The northern Nevada rift: Regional 

tectono-magmatic relations and middle Miocene stress direction: Geological Society of America Bulletin, 

v. 106, p. 371-382. 

35

Table Captions 

Table 1. A sampling of possible indicators of uplift, altitude, and precipitation shadow that may be related 

to Yellowstone hotspot plume uplift. 

Figure Captions 

Figure 1. Map of western United States showing the track of the Yellowstone hotspot (after Pierce and 

Morgan, 1992, fig. 23). The 10- to-2-Ma track is compatible with plate motion, whereas the 10- to 

16-Ma track has apparent rates three times faster than predicted. Projecting back from 10 Ma, we 

infer the plume was located at the site of the hexagon. As shown by the arrowed line, the geologic 

location of the inferred plume head was about 260 km further west at about the McDermitt caldera. 

Assuming the Vancouver slab was in the subsurface (figs. 2 and 3), buoyant rise up the slab may 

have deflected the plume head westward the 260 km shown by the arrowed line. Strontium lines 

after Reed (1993) with modifications in Washington from Robert Fleck (written commun., 1998) and 

southeastern Idaho from Leeman and others (1992). Transect A-A' is portrayed in Figure 5. 

Figure 2. Location of the Yellowstone plume relative to the North American Plate and the Vancouver slab 

(or Juan de Fuca slab; from Severinghaus and Atwater, 1990). The plume is assumed to be fixed in 

the mantle and now 50-km northeast of the 2.0-Ma caldera that started the Yellowstone Plateau 

volcanic field. Hotspot migration of 29 km/m.y. from 2 to 10 Ma is plotted as a circled star on the 

line of section. This rate of 29 km/m.y. is close to the plate motion rate of 22 ± 8 km/m.y. and 

includes tectonic extension that probably accounts for the apparent difference of 7 km/m.y. Cross 

sections are drawn to go through the McDermitt caldera and are nearly parallel to plate motion and to 

the inclination of the Vancouver slab. The reconstruction of Severinghaus and Atwater (1990) 

accounts for deformation, and the numbers on the Vancouver slab indicate the thermal state of the 

slab with 1 meaning solid enough to have earthquakes and 10 meaning nearly the same as the 

surrounding mantle. 

Figure 3. Postulated westward deflection of the Yellowstone plume head by buoyant rise up the inclined 

Vancouver slab. Time sequence from 30 Ma (top) to present (bottom) with position of plume held 

fixed in the mantle. At 20 Ma, the plume head has intersected the inclined Vancouver slab and is 

being displaced westward, and by 16 Ma the plume center is beneath McDermitt (along the OR/NV 

rift), about 260 km west of its feeding plume tail. As the plume flattens, the area of the plume in this 

cross section diminishes to about half because of spreading in the third dimension. The inclination of 

the Vancouver slab is shown to increase from 30 Ma to present (see text under heading "Reconciliation Based..."). Numbers on slab (1, 3, 5, and 10) are the thermal state of slab from Severinghaus 

and Atwater (1990); see also Figure 2. This cartoon does not account for interaction with the 660-

and 410-km discontinuities. 

Figure 4. Some positions for the Yellowstone hotspot at different times as postulated by others and by us. 

The location between 50 and 17 Ma presents a problem that is avoided if the Yellowstone plume 

starts at 17 Ma. 

Figure 5. Geological and geophysical characteristics along section A-A' (see fig. 1). From McDermitt, this 

section goes south, parallel to the Nevada-Oregon rift zone and north parallel to the feeder dikes of the 

36

Columbia River and Oregon Plateau basalts. The Strontium isotopic boundaries(87Sr/86Sr) define crustal 

changes with Precambrian silicic continental crust inside (continent-ward) the 706 line and Mesozoic and 

younger mafic oceanic crust outside (oceanward) the 704 line (87Sr/86Sr 0.704 and 0.706 shown here as 

704 and 706). 

A. 14-17-Ma magmatism near and along the Nevada-Oregon rift zone, which is about 1,100-km 

long (from fig. 1). The rhyolite overlaps both the southern part of the flood basalts and the 

mafic, rift-filling dikes of the Northern Nevada rift. The 14-17-Ma rhyolites fade out roughly in 

the area where 21-37-Ma magmatism had already occurred, perhaps related to crustal depletion. 

B. Crustal ages and origin. Strontium boundaries from Reed (1993, Plate 1, except where updated in 

Washington by Robert Fleck, written commun., 1998, and in southwest Idaho by Leeman and 

others, 1992). Precambrian areas from Link and others (1993, fig. 34 and p. 552-557). Northern 

Nevada from divisions by Elison and others (1990). Oregon after Draper (1991, p. 315) suggesting 2-4 km volcanic and volcaniclastic rock over 15-20 km of oceanic crustal material. 

Carlson and Hart (1987, fig. 8) show this area as new crust over depleted upper mantle. 

C. Heat flow (in mWm-2) from Morgan and Gosnold (1989). 

D. Regional topography from averaging of five parallel profiles spaced 15 km apart and centered on 

section A-A'. 

E. Regional complete Bouguer gravity anomaly from averaging of five parallel profiles spaced 15 

km apart and centered on section A-A'. The data are from the Decade of North American 

Geology (DNAG) data compilation for the conterminous United States (available on CD-ROM 

from National Geophysical Data Center, Boulder, Colorado; Hittelman and others, 1994). Thin 

line shows fit to model shown in F (below). 

F. Gravity model of lithosphere and upper asthenosphere. Crust/mantle boundary taken from an 

average of several Moho maps (Mooney and Weaver, 1989; Pakiser, 1989; Braile and others, 

1989). Southern boundary of "hot asthenosphere" taken from Saltus and Thompson (1995). 

Northern boundary of "hot asthenosphere" interpreted here to coincide with the gravity/topographic step centered at about -380 km on A-A'. The boundary between the upper (felsic) and 

lower (mafic) crust was allowed to vary between sea level and the Moho to match the Bouguer 

anomalies. This simple model shows that the gravity data are consistent with a crust that is 

largely mafic to the north and felsic to the south. 

Figure 6. Geologic cartoon along line of section A-A' showing emplacement of postulated Yellowstone 

plume head. The plume head spreads out (pancakes, mushrooms) upon rising into the lithosphere, 

but it can rise further beneath the thinner lithosphere (young thin mafic crust and lithospheric mantle) 

beneath Oregon. Decompression melting (squiggly vertical lines) produces basalt magma that rises 

upward and interacts differently with the crust: (A) in Oregon, magma rises through the dense, more 

oceanic crust, and surfaces as the Columbia River and Oregon Plateau flood basalts, and (B) in 

Nevada and southern Oregon, basalt magma melts silicic lower crustal material that then rises into 

upper crust forming magma chambers (oval pods) and rhyolitic ignimbrite eruptions, shown as 

volcanoes, and (C) in northernmost Nevada and southern Oregon, both processes have operated. The 

restite from decompression melting remains in the mantle but is buoyant and contributes to longlasting uplift (Morgan and others, 1995). At the time of emplacement, doming above the plume head 

may be 1-2 km (exaggerated in drawing). The northward gravitational push from this doming may 

have produced the Yakima fold belt near the northern end of the section. 

37

Figure 7. Age progression of silicic lava domes of the southern Oregon rhyolite belt. This figure shows 

upper Miocene and younger volcanic rocks and is copied from Christiansen and Yeats (1992, fig. 

55). The rhyolite domes and their ages were shown by MacLeod and others (1976) to have such a 

systematic northwest decrease in age that ages could be contoured in one-million-year increments. 

This northwest age progression is a key argument against a plume origin for the Yellowstone hotspot 

(Lipman, 1992; Christiansen and Yeats, 1992). See Figure 9 for our suggestions about how the 

southern Oregon rhyolite belt might be reconciled with a thermal plume origin of the Yellowstone 

hotspot. Both the southern Oregon rhyolite belt and the Yellowstone hotspot track start in the region 

near McDermitt, but they become younger in nearly opposite directions at the same time. The silicic 

domes (dark) form two belts that trend N. 75o W. The northern belt ends at Newberry caldera and 

parallels the Brothers fault zone (wide gray line) which is actually a complex en echelon pattern of 

normal faults whose trend of about 10o clockwise to the overall pattern (Walker and Nolf, 1981) 

suggests a right lateral component (Pezzopane and Weldon, 1993). Between 10 and 17 Ma, ignimbrites and other silicic volcanism are widespread in the southern Oregon, northern Nevada and 

southwestern Oregon area (fig. 1; Pierce and Morgan, 1992;) but have no systematic progression 

with age in southeast Oregon (MacLeod and others, 1976, p. 470). 

Figure 8. Geologic sketch (after Draper, 1991) along the southern Oregon rhyolite belt parallel to the 

trend of the Brothers fault zone. The mounds numbered 10 to 1 represent contours on the age or 

rhyolitic volcanism in Ma (MacLeod and others, 1976). Following Draper (1991), we suggest several 

factors may have acted separately or in combination to produce penetration of mantle lithosphere by 

basaltic melts to form rhyolite magma from crust: (1) cracking of mantle lithosphere by westward 

migration of activity on Brothers fault zone, (2) westward spreading of plume head, and (3) thermal 

erosion of mantle lithosphere associated with counterflow. Fracturing of the lithosphere along the 

Brothers fault zone (MacLeod and others, 1976; Carlson and Hart, 1987; Pezzopane and Weldon, 

1993) appears to have provided the conduits for volcanism as well as its overall northwest trend. The 

most reasonable mechanism for the northwest progression of rhyolitic volcanism may be a northwest 

progression of faulting to form the Brothers fault zone. Such a northwest progression of faulting 

might have been accompanied by a northwestward migration of plume head material aided by both 

general outward migration of the plume head, and by counter flow between the Vancouver slab and 

North American plate (Draper, 1991). If the two proceeded together, the faulting may have caused 

fracturing of the mantle lithosphere permitting the erosion and stooping of the mantle lithosphere (a 

mechanism suggested by George Thompson, oral commun., 1998) and thus forming an inverted 

channel enhancing both outward (northwestward) plume head flow and decompression melting by 

thinning of the lithosphere. Although Draper (1991) drew a similar diagram, we include this version 

because many have argued that this northwestward volcanic progression negates the Yellowstone 

hotspot as a mantle plume, and we think these factors offer a mantle-plume compatible explanation. 

The geometry of the subducting slab is based on a fairly well controlled section by Parsons and 

others (1998) near the Oregon-Washington border. 

Figure 9. Uplift and climatic patterns that may relate to Yellowstone plume (both head and tail). The 

dashed line shows the pancaked plume with a diameter suggested by Parsons and others (1994) of 

about 400 km but elongated north south as suggested by Camp (1995) parallel to the Nevada-Oregon 

rift zone. Table 1 describes the various localities, areas, and transects shown here. One of the main 

climatic parameters listed is orographic dryness (light shading) in the lee of the inferred Yellowstone 

plume-head uplift. In general, observations of uplift and climate are compatible with uplift above the 

plume head followed by uplift migrating to the east associated with the plume tail (chimney) to the 

38

"Yellowstone crescent of high terrain" associated with the present plume tail (Pierce and Morgan, 

1992) or a similar 650-m-high altitude anomaly that Smith and Braille (1993) associated with the 

current Yellowstone hotspot. 

Figure 10. Temporal relations among Cenozoic climate proxies, "large igneous provinces" (flood basalts) 

and postulated Yellowstone plume-head uplift. Atlantic deep waters (after Miller and others, 1987, as 

shown in Barron and Baldauf, 1989). The about 14- to 17-Ma middle Miocene "climatic optimum" 

(expanded plot, right side from Barron and Keller, 1982) has been associated with the Columbia 

River-Oregon Plateau flood basalts (CRB, for example, Hodell and Woodruff, 1994). After the 

Miocene “climatic optimum,” cooling continued eventually to the Pliocene-Pleistocene ice ages. 

Uplift associated with the Yellowstone plume head is a candidate for a "plateau uplift" whose plumehead mechanism could produce significant late Cenozoic uplift (fig. 9). Global climatic modeling by 

Kutzbach and others (1989) and Ruddiman and others, (1997) shows that “plateau uplifts” could 

produce late Cenozoic climatic cooling in the Northern Hemisphere that might lead to the PlioPleistocene ice ages. Late Cenozoic Yellowstone plume uplift (head and tail) could also diminish 

atmospheric CO2

 leading to cooling, by the mechanism of Raymo and Ruddiman (1992) and 

Ruddiman and others (1997). Plume-head uplift does not have the geophysical problem of a plausible mechanism of how the isostatic balance of the lithosphere could rapidly change that Molnar and 

England (1990a; b) used to challenge the validity of late Cenozoic "plateau uplifts". 

39