The Great Basin Altiplano during the middle Cenozoic ...geology.byu.edu/home/sites/default/files/best_09_altiplano.pdf · Great Basin and the modern crust in the central Andes, and

The Great Basin Altiplano during the middle Cenozoic ignimbriteflareup: insights from volcanic rocks

Myron G. Besta, Deborah L. Barra†, Eric H. Christiansena*, Sherman Grommeb,

Alan L. Deinoc and David G. Tingeya

aDepartment of Geological Sciences, Brigham Young University, Provo, UT 84602-4606, USAb420 Chaucer Street, Palo Alto, CA 94301, USA; cBerkeley Geochronology Center, Berkeley,

CA 94709, USA

(Accepted 26 February 2009)

Uncertainty surrounds the fate of the orogenic plateau in what is now the Great Basin inwestern Utah and Nevada, which resulted from the Mesozoic and earliest Cenozoiccontractile deformations and crustal thickening. Although there is some consensusregarding the gravitational collapse of the plateau by extensional faulting andconsequent crustal thinning, whether or not the plateau existed during the middleCenozoic Great Basin ignimbrite flareup – one of the grandest expressions ofcontinental volcanism in the geologic record – had remained in doubt. We usecompositions of contemporaneous calc-alkaline lava flows as well as configurations ofthe ignimbrite sheets to show that the Great Basin area during the middle Cenozoic wasa relatively smooth plateau underlain by unusually thick crust. We compare analyses of376 intermediate-composition lava flows in the Great Basin that were extruded at42–17 Ma with compositions of .6000 analyses of the late Cenozoic lava flows incontinental volcanic arcs that correlate roughly with known crustal thickness. Thiscomparison indicates that the middle Cenozoic Great Basin crust was much thickerthan the present ca. 30 km thickness, likely as much as 60–70 km. If isostaticequilibrium prevailed, this unusually thick continental crust must have supported hightopography. This high terrain in SE Nevada and SW Utah was progressively smoothedas successive ignimbrite outflow sheets were emplaced over areas currently as much astens of thousands of square kilometres to aggregate thicknesses of as much as hundredsof metres. The generally small between-site variations in the palaeomagnetic directionsof individual sheets lend further support for a relatively smooth landscape over whichthe sheets were draped. We conclude that during the middle Cenozoic, especiallytowards the close of the ignimbrite flareup, this Great Basin area was a relatively flatplateau, and because it was also high in elevation, we refer to it as an Altiplano. It wasnot unlike the present-day Altiplano-Puna in the tectonically similar central Andes,where an ignimbrite flareup comparable to that in the Great Basin occurred at ca.9–3 Ma. Outflow ignimbrite sheets that were deposited from 35 to 23 Ma on theprogressively smoothed Altiplano in south-eastern Nevada were derived from sourcecalderas to the west. Of the 12 major sheets from seven sources, nine are distributedunevenly east of their sources while the remaining three sheets are spread about as fareast as west of their sources. This eccentricity of sources near the western margin of75% of the sheets indicates the existence of a NS-trending topographic barrier incentral Nevada that restricted westward dispersal of ash flows. In a symmetric manner,eastward dispersal of ash flows from sources farther west seemed to have been impededby this same topographic barrier. The westward dispersal was controlled in part bywestward-draining stream valleys incised in the sloping flank of the Great Basin

ISSN 0020-6814 print/ISSN 1938-2839 online

q 2009 Taylor & Francis

DOI: 10.1080/00206810902867690

http://www.informaworld.com

*Corresponding author. Email: [email protected]

International Geology Review

Vol. 51, Nos. 7–8, July–August 2009, 589–633

Downloaded By: [Brigham Young University] At: 20:48 15 September 2009

Altiplano in western Nevada and adjacent California; at least one of these ash flowstravelled as far west as the western foothills of the Sierra Nevada. The nature and originof the implied topographic barrier are uncertain. It is possible that heavy orographicprecipitation on the western slope of the Altiplano and consequent focused denudationand isostatic uplift created a NS-trending topographic high at the crest of the westernslope and facing the smoothed Altiplano to the east. The barrier also lies near andessentially parallel to the buried western edge of the Precambrian basement and to azone of thermal-diapiric domes that were spawned in thickened crust as the basementedge was overrun by late Palaeozoic–Mesozoic thrust sheets.

Keywords: arc volcanic rocks; crustal thickness; great basin; ignimbrite flareup;Altiplano; orogenic plateau

Introduction

The middle Cenozoic ignimbrite flareup (Coney 1978) in the area that became the Great

Basin of the western USA ranks as one of the most voluminous productions of silicic

magma in the geologic record. At least a dozen very large volume (.1000 km3), or

‘supervolcanic’ (de Silva 2008; Miller and Wark 2008), eruptions as well as a greater

number of large volume (100s of km3) eruptions occurred from the latest Eocene to early

Miocene from 36 to 18 Ma (Best et al. 1989a, 1989b, 1993, 1995; Best and Christiansen

1991; Maughan et al. 2002; John et al. 2008). This flareup occurred in a remarkably brief

period of time relative to the span of some 200 million years through the Mesozoic and

into the middle Cenozoic when subduction of oceanic crust beneath the western North

American plate resulted in arc magmatism. In addition to the widespread rhyolite, four of

these supervolcanic eruptions were of crystal-rich dacite magma, the monotonous

intermediates of Hildreth (1981) and Maughan et al. (2002). Eruptions of these colossal

volumes of rhyolitic and hotter dacitic magma, accompanied by extrusion of only very

minor more mafic magma – and no true basalt until the waning stages of the ignimbrite

flareup – imply unusual magma generation in the volcanic arc, requiring a prodigious

amount of thermal energy. To augment the usual heat input from mantle-derived mafic

magmas in the arc system, and to provide the necessary volume of silicic source rock,

necessitates, in our view, an unusually thick crust that was already at near-solidus

temperatures in its deeper part.

Is there evidence for such an unusually thick crust in the Great Basin area during the

ignimbrite flareup?

Tectonic reconstructions and comparisons with active mountain belts together with

palaeobotanical and isotopic data suggest that, following Mesozoic–earliest Cenozoic

contractile deformations, the Great Basin area was a high orogenic plateau capping

unusually thick crust, not unlike the present-day Altiplano-Puna in the central Andes

Mountains. But this crustal welt may have collapsed and thinned before the middle

Cenozoic when the ignimbrite flareup occurred.

The purpose of this paper is to present independent evidence from the middle Cenozoic

volcanic rocks in the Great Basin that indicate the continued existence of an unusually

thick crust during the flareup and that shows it was similar to the Andean Altiplano. Our

plan is to, first, review current thinkings on the nature of the middle Cenozoic crust in the

Great Basin and the modern crust in the central Andes, and then present the compositional

data that indicates an unusually thick crust in the Great Basin. We then describe pertinent

information on the ignimbrite deposits that constrains the topographic character of the

Great Basin Altiplano.

M.G. Best et al.590


Previous thoughts on crustal thickness of the Great Basin during the middleCenozoic

In a widely cited paper, Coney and Harms (1984) reconstructed the crustal thickness in the

Great Basin area after late Palaeozoic and Mesozoic contractile orogenic deformations,

showing the crust to be as much as 50–60 km along a thickened welt, mostly in eastern

Nevada (Figure 1). Although admitting ‘circularity’ (p. 552) in their reconstruction, they

boldly asserted that the eastern Great Basin ‘ . . .was a vast Tibetan or Andean

Altiplanolike plateau prior to middle Cenozoic crustal extension’ (p. 553). Comparing the

active tectonics of the Himalaya–Tibet region and the Andes to the late Mesozoic–early

Cenozoic development of the western USA, Molnar and Lyon-Caen (1988, p. 202)

‘presumed [the existence of a] high plateau in western and central Nevada’ and suspected

that ‘the crust reached a thickness of 50–70 km’. Dilek and Moores (1999) considered the

western US Cordillera as a mature Tibetan Plateau, which has an average elevation of 5 km

underlain by crust of 60–85 km thickness. However, analogies between the western US

Cordillera and the Tibetan Plateau are flawed in that the latter lies inboard of a continent–

continent collision, whereas the US Cordillera resulted from subduction of oceanic

lithosphere studded with island arcs and oceanic plateaus. McQuarrie and Chase (2000)

referred to the elevated thick crust in the hinterland of the Sevier fold-thrust belt in western

Utah and eastern Nevada (Figure 1) as the ‘Sevier Plateau’. In a review of the Cordilleran

thrust belt, (DeCelles 2004, p. 147; see also DeCelles and Coogan 2006) referred to this

hinterland as the ‘Nevadaplano’ whose crustal thickness was 50–60 km and palaeo-elevation

was more than 3 km.

Front

of S

evie

r f o

ld a

nd

th

rust

bel

t

37˚

41˚

109˚111˚114˚117˚120˚

39˚ Ely

Rob

erts

Mou

ntai

nsGol

cond

a SaltLake City

Caliente

CedarCity

Utah

California

Wyoming

Arizona

Colorado

NewMexico

Nevada

Palaeogene crustalthickness (km)Coney and Harms (1984)

Thrust faults and folds

km25 200150100500

50

50

40

40

60

50

50

Las Vegas

Reno

Luni

ng F

ence

mak

er

Austin

Tonopah

Figure 1. Major thrust faults and fold belts in the Great Basin of Nevada and Utah (Oldow et al.1989; McQuarrie and Chase 2000; see also DeCelles 2004) and hypothetical contours (in km) ofearly Tertiary crustal thickness (Coney and Harms 1984).

International Geology Review 591


Several geologists have concluded that the thickened crust underpinning the orogenic

plateau in the Great Basin area became gravitationally unstable and experienced extensional

collapse and consequent thinning. Sonder et al. (1987) modelled these phenomena and

concluded that a time delay of as much as 100 million years can occur between the end of

compression and initiation of extension, depending on factors such as the thermal regime of

the lithosphere and particularly the Moho temperature. For a relatively hot Moho temperature

(500–7008C), the time delay is near zero. While there is no doubt that substantial extension

and crustal thinning has taken place since after the ignimbrite flareup – resulting in the Basin

and Range physiography – we question whether significant extension occurred on a regional

basis before this flareup. For example, in the northeastern Great Basin area, the Camilleri et al.

(1997) model states that the crust thinning was from 70 to 50 km by extensional unroofing

between the Cretaceous and Eocene–Oligocene, whereas Constenius (1996) concluded that

extensional basins formed from the middle Eocene to early Miocene. However, on the basis of

his study of interbedded ash-flow tuffs and sedimentary deposits in northeastern Nevada,

Henry (2008) concluded that Eocene extension was minor. The well-documented extreme

extension found locally in core complexes in the eastern Great Basin (e.g. MacCready et al.

1997) does not appear to be of regional character. Hudson et al. (2000) document greater than

100% extension at 24 Ma in the southern Stillwater Range in west-central Nevada (this and

other localities referred to below are shown in Figure 2), which they interpret to be a local

focusing of regional extension. But 110% extension in the northern Toiyabe and Shoshone

Ranges to the northeast occurred 16–10 Ma (Colgan et al. 2008). Humphreys (1995) points

out that the ignimbrite flareup followed soon after the demise of the Laramide deformation in

the western USA at about 45 Ma with the beginning of north to south rollback of the

subducting oceanic lithosphere from beneath the Great Basin area. Removal of the once ‘flat’

subducting slab brought hotter asthenosphere into contact with the overlying continental

lithosphere prompting voluminous magma generation in the crust. We concur with this

scenario but disagree with Humphrey’s contention that the crust thinned by extension during

the ignimbrite flareup. In our survey of stratigraphic sections of outflow ignimbrite sheets in

the central and eastern Great Basin, we found only limited local evidence for crustal extension

(and consequent thinning) during the flareup (Best and Christiansen 1991). As the flareup

waned, after the maximum production rate of silicic magma eruption at approximately

31–26 Ma, the Great Basin experienced widespread and profound EW crustal extension.

In some places this began as early as 24 Ma, such as in the Stillwater Range (Hudson et al.

2000) and in the Caliente area of southeastern Nevada (Rowley et al. 1995), but was

widespread and profound after about 18 Ma (McQuarrie and Wernicke 2005).

Another approach to the determination of the thickness of the middle Cenozoic crust is to

compensate for the amount of later extensional thinning, assuming plane strain in the EW

direction and using the observed province-wide crustal thickness today of a relatively uniform

30 ^ 5 km (Allmendinger et al. 1987; Mooney and Braile 1989; Gilbert and Sheehan 2004).

However, estimates of the amount of whole-province extensional thinning range widely, for

example, from as much as 100% (Hamilton 1989) to as little as 20–30% (Stewart 1980).

According to the detailed tectonic reconstruction of McQuarrie and Wernicke (2005), the

total extension, mostly after the ignimbrite flareup, oriented about N 788W across the Great

Basin (between longitude about 1128 and 1208 W and latitude 408 240 and 388 400 N), is

236 km or 50%. This value implies an initial crustal thickness before extension of 45 ^ 7 km.

However, this plane-strain calculation assumes the crust to be a closed system, which

is unlikely for two reasons. First, there have been additions to the crust in the form of

mantle-derived magma (Okaya and Thompson 1986; Mayer 1986; Gans 1987; Lachenbruch

and Morgan 1990) that provided heat and added mass to the middle Cenozoic silicic

M.G. Best et al.592


magma systems. The added mafic–ultramafic magmatic rock at the base of the open-system

crust could correspond to the anomalously low velocity (7.4 km/s) layer, several kilometres

thick, interpreted from a seismic profile in north-central Nevada (Catchings 1992) and the

7.5 km/s lens in western Utah and easternmost Nevada (Smith et al. 1989). Second, Bird

(1991; see also Gans 1987) claims that the puzzling flat Moho, despite spatially heterogeneous

extension, is the consequence of horizontal ductile flow of the hot lower crust between the

more rigid upper crust and mantle in the thickened (55? km), high (2? km) crust that resulted

from the Sevier orogeny. McQuarrie and Chase (2000) claimed that eastward flow from the

Sevier hinterland thickened and elevated the Colorado Plateau crust. Wernicke et al. (1988)

claimed flow towards the south into the highly extended Las Vegas–Death Valley corridor.

Palaeoaltitude of the Great Basin area

Palaeoaltitudes provide additional and independent insight into the crustal thickness of the

middle Cenozoic Great Basin. But because of an imperfect worldwide correlation between

42˚

41˚

40˚

39˚

38˚

37˚

36˚

35˚

120˚ 119˚ 118˚ 117˚ 116˚ 115˚ 114˚ 113˚ 112˚ 111˚ 110˚ 109˚

TQCC

CNCC

CCC

IPCC

Figure 2. Locations of 376 analysed samples of middle Tertiary lava flows in the Great Basin,which lies between the Sierra Nevada and the Colorado plateaus, are shown by red circles. Becauseof the small scale of this figure some sample sites overlap as a single point. Also shown are theoutlines of the CCC, Caliente caldera complex; CNCC, central Nevada caldera complex; IPCC,Indian Peak caldera complex; and TQCC, Toquima caldera complex as well as other localities citedin the text, as follows: A, Austin; BM, Battle Mountain; CA, Clan Alpine Mountains; CS, CarsonSink; DD, Diamond Mountains; DM, Dogskin Mountain; DS, Donner Summit; DV, Death Valley;FC, Fish Creek Mountains; HP, Haskell Peak; LV, Las Vegas; NP, New Pass Range; PA, PahrocRange; PR, Paradise Range; P, Provo, Utah; RR, Reese River Valley; SC, Schell Creek Range; S,Seaman Range; SH, Shoshone Range; ST, Stillwater Range; T, Tonopah; TQ, Toquima Range; TY,Toiyabe Range; W, Wah Wah Mountains; Y, Yerington; and YR, Yuba River drainage.



crustal thickness and topographic elevation, it is not feasible to infer quantitative values of

the former from the latter.

Despite inherent uncertainties, palaeoaltitudes derived from palaeobotanical studies

provide qualitative insight into crustal thickness. Such studies indicate that high altitudes

prevailed in the western USA during the Cenozoic (see review by Chase et al. 1998). But,

regrettably, there are only limited data for the Great Basin. Wolfe et al. (1997) list a dozen

sites in west-central Nevada for which mid-Miocene floras (mostly 15–16 Ma) indicate

palaeoelevations of about 3 km above sea level, 1–1.5 km above the present elevation,

thus implying thicker crust. Gregory-Wodzicki (1997) determined palaeoaltitude from a

late Eocene flora in western Utah of 2.9–3.6 km, compared to the present elevation of

1.7 km.

Palaeoaltitudes can also be determined isotopically. Using (U–Th)/He ages of apatites

in the Sierra Nevada, House et al. (2001) argue that the range may have occupied a

position at the western edge of an orogenic plateau of at least 3 km elevation that spanned

the Cordilleran interior during Late Cretaceous time. Deuterium values in Eocene

sediment deposits in the Sierra Nevada led Mulch et al. (2006) to basically the same

conclusion as those of House et al. (2001). Horton et al. (2004; see also Poage and

Chamberlain 2002) tracked changing palaeoelevations in the Great Basin using shifts in

oxygen and hydrogen isotope ratios in authigenic minerals precipitated from meteoric

waters in the Cenozoic basin deposits. They argue that the rain-shadow effect of a high

Sierra Nevada that caused precipitation of isotopically lighter isotopes of oxygen and

hydrogen farther inland in the Great Basin area prevailed throughout much of the

Cenozoic. Isotopic shifts corresponding to about 2 km of surface uplift of the Great Basin

between the middle Eocene and early Oligocene were followed by subsidence from the

middle Miocene to Pliocene. A similar pattern of elevation change was claimed by Wolfe

et al. (1997) using palaeobotanical data. Crowley et al. (2008) analysed oxygen isotopes in

mammalian tooth enamel and concluded that a Sierran rain-shadow had existed since at

least 16 Ma. Kent-Corson et al. (2006) conclude there is a growing body of evidence for a

migration of high surface elevation from the northern to the southern Great Basin from the

late Eocene to the Miocene. They relate this phenomenon to the southward rollback of

the subducting slab and upwelling of hotter and expanding asthenosphere beneath the

continental lithosphere (Davis et al. 2009).

These conclusions are at variance with the long-standing belief that significant uplift of

the Sierra Nevada to its present elevation had not occurred until the latest Cenozoic

(e.g. since ca. 5 Ma). For example, Wakabayashi and Sawyer (2001) constructed this

scenario from an extensive analysis of current topography coupled with thermochrono-

logic and geobarometric data. Clark et al. (2005) concluded that the southern Sierra

experienced two stages of uplift since 32 Ma that brought the elevation to 4 km that is

observed today from the roughly 1.5 km prevailing through the early Cenozoic.

A possible indication of the altitude of the Great Basin ignimbrite province before its

subsequent extensional collapse and subsidence lies in the altitude of the present Sierra

Madre Occidental ignimbrite plateau in Mexico (Swanson et al. 2006). This ignimbrite

plateau is readily interpreted to be a southward continuation of the Great Basin ignimbrite

province along the western North American continental margin with which it shares many

similarities (Best et al. 1989b, Table 1) but has essentially escaped late Cenozoic

extensional faulting after the ignimbrite flareup. The Sierra plateau lies mostly between

2.2 and 2.4 km elevation, roughly 0.5 km higher than the present average Great Basin.

Peaks in the plateau are as much as 1.8 km above and canyons 1 km below the 2.2–2.4 km

elevation.

M.G. Best et al.594


Tab

le1

.C

hem

ical

and

iso

top

icd

ata

for

con

tin

enta

lv

olc

anic

arcs

of

kn

ow

ncr

ust

alth

ick

nes

san

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lav

asin

the

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atB

asin

.

Reg

ion

Sy

mb

ol

Ch

emic

alan

aly

ses

(n)

Sr

iso

top

ean

aly

ses

(n)

K2O

(57

%S

iO2)

K2O

/Na 2

O(5

7–

63

%S

iO2)

Min

87S

r/8

6S

r(,

63

%S

iO2)

87S

r/8

6S

r(m

ean

all)

Th

ick

nes

scr

ust

(km

)R

efer

ence

An

des

cen

tral

A2

00

11

02

.20

0.8

00

.70

50

0.7

07

06

0L

eem

an(1

98

3)

An

des

no

rth

A5

43

91

.50

0.6

30

.70

38

0.7

04

35

5L

eem

an(1

98

3)

An

des

sou

thA

66

21

1.4

00

.70

34

0.7

04

23

5L

eem

an(1

98

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des

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65

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and

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98

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(19

93

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69

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28

25

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00

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06

30

.70

81

70

Fee

ley

and

Dav

idso

n(1

99

4)

An

des

Pay

ach

ata

A1

81

82

.75

0.7

40

.70

61

0.7

06

66

5D

avid

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al.

(19

90

)A

nd

esT

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11

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04

76

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and

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orb

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(19

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nd

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15

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67

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(19

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Alt

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22

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(19

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(19

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51

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40

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42

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04

35

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(19

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71

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and

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(19

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33

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04

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51

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98

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24

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48

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and

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rmey

etal.

(19

95

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Tab

le1

–continued

Reg

ion

Sy

mb

ol

Ch

emic

alan

aly

ses

(n)

Sr

iso

top

ean

aly

ses

(n)

K2O

(57

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/Na 2

O(5

7–

63

%S

iO2)

Min

87S

r/8

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r(,

63

%S

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87S

r/8

6S

r(m

ean

all)

Th

ick

nes

scr

ust

(km

)R

efer

ence

An

des

Des

cab

ez-G

ran

dA

66

1.6

20

.46

0.7

03

80

.70

39

45

Hil

dre

than

dM

oo

rbat

h(1

98

8)

An

des

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roA

zul

A4

32

51

.70

0.4

60

.07

36

0.7

03

94

5H

ild

reth

and

Mo

orb

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(19

88

)A

nd

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a-S

anP

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21

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0.5

44

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uso

net

al.

(19

92

)A

nd

esL

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na

del

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14

92

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20

.70

39

0.7

04

14

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reyet

al.

(19

84

)A

nd

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anP

edro

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lad

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23

91

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0.6

10

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37

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03

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al.

(19

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esC

ord

on

El

Gu

adal

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48

1.9

00

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0.7

03

90

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40

41

Fee

leyet

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(19

98

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nd

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gav

iA

11

21

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0.7

03

94

0H

ild

reth

and

Mo

orb

ath

(19

88

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nd

esA

ntu

coA

22

61

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0.7

03

74

0H

ild

reth

and

Mo

orb

ath

(19

88

)A

nd

esC

hil

lan

A2

41

.29

37

Hil

dre

than

dM

oo

rbat

h(1

98

8)

An

des

Pu

yeh

ue

A6

43

21

.40

0.3

70

.70

38

0.7

04

03

2G

erla

chet

al.

(19

88

)A

ndes

Moch

o-C

hosh

uen

coA

20

20

1.0

00

.70

40

0.7

04

13

0M

cMil

lanet

al.

(19

89

)A

nd

esC

alb

uco

A2

29

0.7

00

.18

0.7

03

70

.70

40

30

Lo

pez

-Esc

ob

aret

al.

(19

95

)A

leu

tian

sK

atm

aiK

36

16

1.1

00

.35

0.7

03

00

.70

33

33

Hil

dre

thet

al.

(20

04

)E

ast

Ale

uti

ans

K2

55

24

1.2

00

.44

0.7

02

80

.70

32

38

Lee

man

(19

83

)C

asca

des

C1

90

1.3

00

.43

0.7

02

64

0L

eem

an(1

98

3)

Cas

cad

esM

tH

oo

dC

19

01

.20

41

Lee

man

(19

83

)C

asca

des

Cra

ter

Lak

eC

57

1.1

00

.28

41

Bac

on

and

Dru

itt

(19

88

)C

asca

des

Med

icin

eL

ake

C3

51

.40

0.3

23

9M

ertz

man

(19

77

)C

asca

des

Mt

St

Hel

ens

C3

29

1.3

00

.27

0.7

03

00

.70

36

41

Hal

lid

ayet

al.

(19

83

)C

asca

des

Mt

Sh

asta

C1

12

32

1.0

00

.31

0.7

02

80

.70

34

38

Gro

veet

al.

(20

05

)S

anJu

anp

reca

lder

aC

on

ejo

sS

56

14

2.5

00

.82

0.7

04

60

.70

51

48

Co

lucc

iet

al.

(19

91

)

M.G. Best et al.596


Tab

le1

–continued

Reg

ion

Sy

mb

ol

Ch

emic

alan

aly

ses

(n)

Sr

iso

top

ean

aly

ses

(n)

K2O

(57

%S

iO2)

K2O

/Na 2

O(5

7–

63

%S

iO2)

Min

87S

r/8

6S

r(,

63

%S

iO2)

87S

r/8

6S

r(m

ean

all)

Th

ick

nes

scr

ust

(km

)R

efer

ence

San

Juan

po

stca

lder

aH

uer

toS

20

72

.30

0.6

80

.70

43

0.7

04

74

8P

arat

etal.

(20

05

)

San

Juan

syn

cald

era

S1

80

.70

45

0.7

06

34

8R

icip

utiet

al.

(19

95

)M

exic

oP

26

14

91

.50

0.7

03

20

.70

39

40

Lee

man

(19

83

)M

exic

o,

Val

ley

of

P7

31

.50

0.4

04

3W

alla

cean

dC

arm

ich

ael

(19

99

)M

exic

oN

WM

ex.

vo

lc.

bel

tP

16

16

1.5

50

.41

0.7

03

70

.70

40

30

Ver

ma

and

Nel

son

(19

89

)M

exic

oS

ang

ang

uey

P1

41

.50

0.4

73

0N

elso

nan

dL

ivie

res

(19

86

)M

exic

oP

aric

uti

nP

27

12

1.3

00

.40

0.7

03

70

.70

40

33

Wil

cox

(19

54

)an

dM

cBir

ney

etal.

(19

87

)C

entr

alA

mer

ica

M3

37

40

1.5

00

.45

0.7

03

10

.70

40

33

Lee

man

(19

83

)G

uat

emal

aA

M8

71

.50

0.5

35

0C

arret

al.

(20

03

)G

uat

emal

aB

M3

30

61

.40

0.4

20

.70

38

0.7

03

94

6C

arret

al.

(20

03

)G

uat

emal

a-E

lS

alv

ado

rM

11

62

21

.60

0.4

60

.70

32

0.7

03

63

8C

arret

al.

(20

03

)H

on

du

ras,

Yo

ho

aM

13

11

0.7

02

90

.70

30

37

Car

ret

al.

(20

03

)E

lS

alv

ado

r,H

on

d,

Nic

arag

ua

M2

46

69

1.5

00

.48

0.7

03

60

.70

40

34

Car

ret

al.

(20

03

)

Co

sta

Ric

aR

inco

nM

19

1.6

00

.52

36

Car

ret

al.

(20

03

)C

ost

aR

ica

oth

erM

76

71

.30

0.4

40

.70

38

0.7

03

83

7C

arret

al.

(20

03

)C

ost

aR

ica

Pla

tan

arM

50

1.7

00

.62

0.7

03

60

.70

37

39

Car

ret

al.

(20

03

)C

ost

aR

ica

Po

asM

21

41

.80

0.5

70

.70

37

0.7

03

74

1C

arret

al.

(20

03

)C

ost

aR

ica

Bar

ba

M2

13

2.6

00

.77

0.7

03

60

.70

37

43

Car

ret

al.

(20

03

)C

ost

aR

ica

Iraz

uM

50

2.2

04

5C

arret

al.

(20

03

)L

esse

rA

nti

lles

Do

min

ica

L2

80

23

0.9

00

.70

41

0.7

04

53

3L

eem

an(1

98

3)

Les

ser

An

till

esG

ren

ada

L2

64

52

1.2

00

.70

39

0.7

04

93

3L

eem

an(1

98

3)

Les

ser

An

till

esS

aba

L4

21

.10

0.3

53

3D

efan

tet

al.

(20

01

)



Tab

le1

–continued

Reg

ion

Sy

mb

ol

Ch

emic

alan

aly

ses

(n)

Sr

iso

top

ean

aly

ses

(n)

K2O

(57

%S

iO2)

K2O

/Na 2

O(5

7–

63

%S

iO2)

Min

87S

r/8

6S

r(,

63

%S

iO2)

87S

r/8

6S

r(m

ean

all)

Th

ick

nes

scr

ust

(km

)R

efer

ence

Les

ser

An

till

esS

tK

itts

L1

73

14

0.6

00

.70

36

0.7

03

73

3L

eem

an(1

98

3)

Kam

chat

ka

H3

30

12

1.3

00

.37

0.7

03

10

.70

33

35

Lee

man

(19

83

)T

urk

ey-I

ran

T1

39

2.3

00

.84

0.7

03

60

.70

50

50

Lee

man

(19

83

)N

ewZ

eala

nd

Z2

34

16

1.5

00

.42

0.7

04

10

.70

52

35

Lee

man

(19

83

)S

um

atra

Jav

aU

26

21

08

1.7

00

.70

39

0.7

04

73

0L

eem

an(1

98

3)

Jap

anK

yu

shu

J5

41

?1

.50

0.7

03

70

.70

50

28

Lee

man

(19

83

)Ja

pan

Ho

nsh

uJ

59

57

81

.00

0.7

02

60

.70

40

30

Lee

man

(19

83

)Ja

pan

Ho

kk

aid

oJ

12

54

61

.10

0.7

02

60

.70

37

28

Lee

man

(19

83

)N

evad

aE

gan

Ran

ge

26

21

3.2

01

.30

0.7

08

40

.71

25

Fee

ley

and

Gru

nd

er(1

99

1)

and

Gru

nd

er(1

99

2)

Cen

tral

Nev

ada

cald

era

com

ple

x7

22

.30

1.2

60

.70

93

0.7

10

1A

skre

n(1

99

2)

Ind

ian

Pea

kca

lder

aco

mp

lex

10

72

.90

1.1

90

.70

71

0.7

09

1A

skre

n(1

99

2)

and

Har

tet

al.

(19

98

)A

llG

reat

Bas

in3

76

2.9

01

.01

Bar

r(1

99

3)

M.G. Best et al.598


Geology of the central Andean plateau in brief

Because the tectonic setting of the modern central Andean continental margin appears to be

an appropriate model for the middle Cenozoic Great Basin area, it is useful to examine

pertinent aspects of Andean geology (Allmendinger et al. 1997; de Silva et al. 2006).

With an average elevation of 4 km and volcano peaks of more than 6 km, the relatively

arid 350–400 by 1800 km central Andean plateau is the highest region in the world

associated with voluminous arc volcanism. The internally draining Altiplano, at an

average elevation of 3.8 km, lies in the northern two-thirds of the plateau. Attesting to its

low relief as an enclosed basin during the Pleistocene, much of the Altiplano was covered

by Lake Ballivian whose remnants include Lakes Titicaca and Poopo and several salars

(playas). In the southern plateau, in northwest Argentina and a small part of adjoining

Chile, Puna has an average elevation of 5 km.

The 60–70 km thick crust of the Altiplano-Puna, which lies above the 308-east-dipping

Nazca plate, is predominantly felsic (Beck and Zandt 2002), but the lower part may be

mafic. A decrease in crustal thickness in the topographically higher Puna implies a

relatively low-density uppermost mantle, perhaps resulting from recent lithospheric

delaminations and invasion of hotter asthenosphere beneath the crust. Although

delamination may have contributed to the plateau elevation, workers generally agree that

most of the uplift resulted from tectonic shortening of the crust through the Palaeogene to

essentially its present state by the mid-Miocene (Hartley et al. 2007).

After a long period of thrust faulting, the southern and northern margins of the central

Andean plateau are now experiencing NS crustal extension with development of stretching

faults perpendicular to the mountain belt. Tibet is also currently experiencing similar

extension. This extension appears to be a result of gravitational buoyancy forces exceeding

compressive tectonic forces in the elevated plateaus (Molnar and Lyon-Caen 1988; Dilek

and Moores 1999, p. 932). Rare, but widespread, EW-striking dikes and local grabens in

the Great Basin indicate a similar state of stress during the middle Cenozoic (Best 1988;

Best et al. 1998).

The central Andean plateau experienced an ignimbrite flareup from the late Miocene

through the Pliocene that apparently accompanied a steepening of the dip of the subducting

Nazca plate. The increased volume of the asthenospheric wedge promoted melting in the

crust, thereby generating copious volumes of silicic magmas that created the largest young

ignimbrite province on Earth, covering more than 500,000 km2 (Allmendinger et al. 1997,

p. 157). The focus in space and time of this flareup lies in the Altiplano-Puna volcanic

complex (de Silva 1989), near the common borders of Argentina, Bolivia, and Chile, where

the ash-flows were erupted mostly from 9 to 3 Ma (Salisbury et al. 2008). During this flareup

about a dozen known calderas developed (e.g. Ort et al. 1996; Lindsay et al. 2001; Soler et al.

2006), the largest of which measures 65 £ 35 km. The ignimbrite deposits cover an area of

about 70,000 km2 with an aggregate volume of nearly 15,000 km3 (de Silva et al. 2006;

de Silva and Gosnold 2007). Seven of the individual deposits exceed 1000 km3, thus

qualifying as supervolcanic, and typically are of crystal-rich, calc-alkaline high-K dacite.

Another of this genre, but located 200 km south of the Altiplano-Puna volcanic complex, is the

Cerro Galan ignimbrite (Francis et al. 1989); its outflow sheet is upwards of 200 m thick and

extends as much as 100 km in all directions from its 50 £ 30 km resurgent caldera source.

Essentially coextensive with the Altiplano-Puna volcanic complex is an underlying

subhorizontal slab of magma 1–2 km thick at a depth of 17–19 km (Zandt et al. 2003;

de Silva et al. 2006), which may be a remnant of the magma-generating activity during the

ignimbrite flareup.



The tectonic setting and ignimbrite flareup of the central Andean plateau during the

late Cenozoic bear significant similarities to the middle Cenozoic Great Basin area

(Best and Christiansen 1991, pp. 13,522–13,525, Figure 14; Maughan et al. 2002,

pp. 150–154, Figure 18a). Especially noteworthy is the occurrence in both of several

thousands of cubic kilometres of crystal-rich, calc-alkaline high-K dacite ignimbrite–the

monotonous intermediates of Hildreth (1981) and Maughan et al. (2002) – together with

lesser, but still prodigious, volumes of rhyolite ash-flow tuff. Such large scale,

supervolcano melting events require unusually large inputs of mantle-derived magma into

the continental crust to power the silicic magma-generating systems. But a compounding,

even necessary, factor in such large-scale melting is an unusually thick and pre-warmed

crust (de Silva et al. 2006).

Chemical composition of middle Cenozoic Great Basin lava flows: implications for

crustal thickness

The tectonic and geologic similarities between the Andean orogenic plateau underlain by

60–70 km-thick crust and the middle Cenozoic Great Basin area suggest that the latter was

also a high plateau underlain by unusually thick crust. However, the inherent uncertainties

in palaeo-elevation determinations and the complexity of the tectonic and erosional

history of the Great Basin and its inferred crustal thickness, as summarized above, provide

little unambiguous support for this suggestion. Obviously, there is a need for an

independent determination of the thickness of the Great Basin crust during the middle

Cenozoic. Did thick crust prevail through the ignimbrite flareup, or not? We seek an

answer to this question in the composition of the volcanic rocks.

Rationale and previous work

It is well known that the composition of calc-alkaline magmas extruded in subduction-

related or arc settings reflects the type and thickness of the crust through which they ascend

(e.g. Gill 1981; Leeman 1983). Generally, the greater the path length through relatively

less dense and thicker felsic continental crust the greater is the opportunity for ascending

mantle-derived magma to be modified by differentiation processes, yielding more evolved

felsic magma compositions to be extruded. Longer paths allow more opportunity for

fractional crystallization and open-system assimilation of felsic continental rock, as well

as mixing with incompatible-element-enriched partial melts, especially in lower levels of

thick and therefore hotter crusts. Consequently, extruded magmas that experienced a

longer ascent path through continental crust will have greater concentrations of elements

such as K, Rb, and Ba, and a greater initial 87Sr/86Sr. A smaller proportion of more

primitive low-silica, basaltic magma is expected to be extruded. But because of

differences in age of the continental crust and its composition around the world and

because of variable rates of ascent and amounts of open-system differentiation of the

primary mantle-derived magmas there can be only a crude correspondence between the

composition of extruded magmas and the thickness of the crust at continental margins.

An additional factor governing the composition of arc volcanic rocks stems from sub-

crustal factors, such as where the primitive basaltic magmas are generated and

contamination with sediment dragged down on the subducting oceanic plate. Nonetheless,

variations in the mantle component in ascending magmas can be swamped by the

thickness-related crustal contribution (Hildreth and Moorbath 1988).

M.G. Best et al.600


Leeman (1983) inventoried whole-rock chemical and isotopic parameters of volcanic

rocks in mostly active arcs around the world, where the thickness of the crust has been

documented. Hildreth and Moorbath (1988) examined analyses of 15 andesite–dacite

composite volcano suites in the Andes that are sited on crust ranging in thickness from

about 30 to 60 km. We have expanded and updated these two compilations to include

locales, where the thickness of the crust is constrained to within approximately 3 km. Our

resulting ‘global’ data base in Table 1 is based on more than 6000 individual analyses

comprising 72 suites, including 26 individual volcanic suites in the Andes, suites in 10

areas of western North America, 11 suites in segments of the central America arc, 4 suites

from the Caribbean, and 6 suites from the western Pacific. The only few analyses included

from the Mediterranean arc are from Turkey–Iran, because of the complexity of plates and

plate interactions and the common occurrence of highly potassic volcanic rocks, such as in

the Eolian Islands. Minor proportions of alkaline rocks occur together with the more

prominent and typical calc-alkaline compositions in some arc suites listed in our global

data base. These alkaline rocks, in the active Mexican volcanic belt and the Oligocene San

Juan volcanic field of Colorado, have not been included in the global data base. In the

northwestern Mexican volcanic belt the alkaline rocks appear to be associated with local

rifting and the more mafic compositions have an oceanic-island-like source (Verma and

Nelson 1989); alkaline basalts in the Valley of Mexico have a complex ancestry

(Wallace and Carmichael 1999). In the San Juan field, where we assume the present crustal

thickness of about 48 km also prevailed during volcanism, extrusion of minor amounts of

Rock Creek alkaline lavas of the Conejos Formation preceded caldera-forming ash-flow

eruptions while trachyandesitic lavas of the Huerto Andesite followed. These water-poor,

mostly high-Zr (to as much as 580 ppm) alkaline rocks are believed to have resulted from

crystal fractionation of mantle-derived alkalic basalt parent magmas, modified by

assimilation of prior sub-volcanic intrusions (Parat et al. 2005), in contrast to the more

open-system, wetter and more oxidizing conditions that dominate the evolution of normal

arc calc-alkaline magmas.

Our global data base in Table 1 lists essentially the same compositional parameters as

used by Leeman (1983) in his inventory, viz.: (1) The average K2O/Na2O ratio for all

analyses of andesites (57–63 wt. % SiO2) within a suite. Leeman (1983) also used the

average K2O/Na2O in basaltic andesites, but these more mafic lavas are uncommon in the

Great Basin and their ratios were not used in our investigation. (2) The mean value of K2O

at 57.5 wt. % silica in a K2O–SiO2 variation diagram for the suite. (3) The minimum,

mean, and maximum initial 87Sr/86Sr ratio for the rocks in the suite. Additional data

include the crustal thickness determined by geophysical methods as cited in published

works. For the Andean locales, crustal thicknesses are from Hildreth and Moorbath (1988)

and Allmendinger et al. (1997), and for the western US, from Mooney and Braile (1989),

Prodehl and Lipman (1989), and Gilbert and Sheehan (2004). Nd and Pb isotopes are not

listed in Table 1 as such data are not available in many publications, especially older ones.

Neither are Mg and Na concentrations that have been correlated with crustal thickness in

oceanic arc settings. Leeman (1983) used the percentage of andesite, dacite, and rhyolite in

suites, but because of a limited number of samples in some individual suites or because of

sampling biases, we found considerable scatter in this parameter; thus, it was not used.

Great Basin data set

We collected, and analysed by X-ray fluorescence spectrometry, 268 samples of

essentially unaltered middle Cenozoic lava flows in the Great Basin that were extruded



42–17 Ma. An additional 108 high-quality whole-rock analyses, by a variety of methods,

of rocks of the same time period were selected from reports published by other workers,

yielding a total of 376 analyses that form the Great Basin data set used here (Barr 1993; see

supplementary data at www.informaworld.com/tigr). The analyses are of lavas bracketing

the Great Basin ignimbrite flareup in time (36–18 Ma) and space (Figure 2).

Many of the samples have been dated by various isotopic methods (e.g. McKee et al.

1993). For others, an approximate age of varying uncertainty is provided from bracketing

dated ash-flow sheets and from latitudinal proximity to other isotopically dated lava flows.

The latitudinal proximity basis for age approximation follows from the general southward

sweep of inception of volcanism through the Great Basin during the Cenozoic (Best et al.

1989b, Figure 3; Best and Christiansen 1991, Figure 2). Though imperfect, the pattern of

inception provides a useful approximation, in most cases, for the age of an individual lava

sample, where no other data are available.

Rhyolites and quartz-rich high-silica dacites were excluded from our Great Basin data

set, because the intent of the sampling programme was to inventory the mantle

contribution in arc volcanic rocks, however much that might be. Even though we made a

special effort to collect the most mafic lava samples in the Great Basin, only 14% (53) of

our samples have less than 57 wt. % silica (Figures 3 and 4). Only five true basalts,

according to the IUGS chemical classification (Le Maitre 1989), were found and these are

17–20 Ma. Hence, the greater than 17 Ma cut-off effectively excludes basalts from our

data set. ‘Fundamentally basaltic’ rocks in the central and eastern Great Basin related to

extensional tectonism (Christiansen and Lipman 1972) are mostly less than 13 Ma and

have been dealt with in regional studies (e.g. Best et al. 1980; Fitton et al. 1991; Nelson

and Tingey 1997). Eight of our samples with less than 52 wt. % silica are alkaline and all

but one of these are 17–23 Ma, the other being 38 Ma (McKee et al. 1993). Alkaline

samples seen in Figure 4 constitute only 6% (24) of our data set; they have mostly less than

57 wt. % silica and lack the high-Zr content of the alkaline rocks of the San Juan field

described above. Of the remainder, 4.5% (16) are tholeiitic (ferroan) and the rest are calc-

alkalic (magnesian), according to the criteria of Miyashiro (1974).

Subduction-related lavas were extruded throughout the 42–17 Ma period in the Great

Basin. However, a fundamental transition in the composition of volcanic rocks occurred

0

1

2

3

4

5

6

7

45 50 55 60 65 70

SiO2 (wt%)

K2O

(w

t%)

Basalt Basalticandesite

Andesite Dacite

SHOSHONITIC

LOW K

MEDIUM K

HIGH K

Figure 3. K2O–SiO2 diagram for middle Cenozoic lava flows from the Great Basin. Fieldboundaries and rock-suite labels from Le Maitre (1989) except shoshonitic from Ewart (1982).

M.G. Best et al.602


about 21–17 Ma when the sweeping volcanic activity reached and stagnated in the

southern part of the province (Best and Christiansen 1991, Figure 2). Three of the more

mafic lava flows in our data set extruded after about 21 Ma possess higher Nb and Ti

concentrations, manifesting waning generation of arc magmas typified by negative Nb and

Ti anomalies in normalized element variation diagrams. Increasing proportions of true

basalt lacking an arc chemical signature were extruded after 20 Ma, along with high-silica

rhyolite, heralding the initiation of a bimodal volcanic association accompanying crustal

extension (Christiansen and Lipman 1972).

Our Great Basin data set is essentially a unimodal spectrum of intermediate-

composition arc lavas that are mostly high-K but with lesser medium-K and shoshonitic

compositions (Figures 3 and 4). Most samples are andesite with lesser dacite, trachydacite,

basaltic andesite and K-rich latite and shoshonite. The most widespread phenocryst

assemblage in the lavas includes plagioclase, two pyroxenes, and Fe–Ti oxides. In rocks

that contain greater than 55 wt. % silica, amphibole, biotite, sanidine, and quartz are

common whereas olivine is essentially restricted to rocks having less than 55 wt. % silica.

Comparative determination of crustal thickness

Values of compositional parameters in the global data base (Table 1) are plotted against

crustal thickness in Figures 5–12. Linear best-fit lines (not shown) have R2 values of

0.47–0.50. Parameter values for the four sets of data for the Great Basin (‘Nevada . . . ’ and

‘All Great Basin’ bottom of Table 1) are shown as arrows.

Average K2O/Na2O ratio in andesites

The global data show a positive correlation between the average ratios and crustal

thickness up to about 52 km, beyond which the ratios, all for the Andes, appear to flatten

out (Figure 5). The four Great Basin ratios, shown by arrows in the diagram, all lie above

the global ratios, suggesting the Great Basin crust was greater than about 50 km thick. In

0

2

4

6

8

10

12

45 50 55 60 65 70

Na 2

O +

K2O

(w

t %)

Basalt Basalticandesite

Andesite Dacite

TrachydaciteLatite

Shoshonite

Trachy-basalt

Basanite

ALKALINEALKALINE

SiO2 (wt%)

SUBALKALINESUBALKALINE

Figure 4. Total alkalies–silica diagram for middle Cenozoic lava flows from the Great Basin. Fieldboundaries and rock-type labels from Le Maitre (1989). Dotted line (from Le Bas et al. 1992)separates 23 alkaline (silica undersaturated) samples above from 353 subalkaline below.



A

AA

A

A

A

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A

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AA

A

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AA

A

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AAA

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KK

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P PPP

P

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U M

J

JJ

0.0

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2.0

2.5

3.0

3.5

20 30 40 50 60 70 80

Thickness of crust (km)

K2O

(w

t.%)

at 5

7.5

wt.%

SiO

2

Egan Range

CentralNevada

Indian PeakGreat Basin

Figure 6. Thickness of crust plotted against K2O at 57.5 wt% SiO2 on variation diagrams for thesuites of lava samples listed in Table 1. Arrows on left side of diagram are values of middle Cenozoiclava suites in the Great Basin.

Z

M

MM

T

HLC

M

MM M

M

M

MP

PP P

S

S

CC

C

CK

K

A

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0.0

0.2

0.4

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0.8

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1.4

20 30 40 50 60 70 80


Ave

rage

K2O

/Na 2

O (

wt%

) in

and

esite

s (5

7-63

wt%

SiO

2)

Egan RangeCentral NevadaIndian Peak andGreat Basin >30 Ma

Great Basin <30 Ma

Figure 5. Thickness of the crust plotted against average K2O/Na2O in andesites for the suites oflava samples listed in Table 1. Arrows on left side of diagram are values of middle Cenozoic lavasuites in the Great Basin.

M.G. Best et al.604


J J

J

Z

TMU

H

LLL

MM

C

MM

MM

M

M

PP

P

SSS

C

CK

K

AA

A AAAAA A A

AAA

AA

A

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A

0.7020

0.7030

0.7040

0.7050

0.7060

0.7070

0.7080

0.7090

20 30 40 50 60 70 80Thickness of crust (km)

Egan Range

Indian Peak

Central Nevada

Min

imum

initi

al 87

Sr/

86S

r

Figure 7. Thickness of crust plotted against the minimum initial 87Sr/86Sr in lavas with ,63 wt. %SiO2 for the suites of samples listed in Table 1. Arrows on left side of diagram are values of middleCenozoic lava suites in the Great Basin.

JJ

L

LJ

UZ T

H

L

MMMC

MM

M

MMMPP P

S

SS

CKK

AAA AAA AAA AAA

AA

A

AA

A

A

A

A

A A

A

0.7020

0.7040

0.7060

0.7080

0.7100

0.7120

0.7140

20 30 40 50 60 70 80


Ave

rage

initi

al 87

Sr/

86S

r

Egan Range

Indian Peak

Central Nevada

Figure 8. Thickness of crust plotted against the average initial 87Sr/86Sr in the suites of lavas listedin Table 1. Arrows on left side of diagram are values of middle Cenozoic lava suites in the GreatBasin.



the Great Basin data set, the average K2O/Na2O in andesites older than 30 Ma is clearly

greater than in younger andesites (not distinguished in Figure 5). The implied thinning of

the crust with time is consistent with removal of lower crustal rock by outward ductile

flow, or by extension, or both, as discussed above. The average K2O/Na2O is slightly

greater in lavas east of Longitude 1178 W (not distinguished in the Figure 5), suggesting a

slightly thicker crust to the east on the Precambrian basement (see below).

4 4

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0

400

800

1200

1600

2000

40 45 50 55 60 65 70 75

SiO2 (wt%)

Ba

(ppm

)

3

3333 3

333

3 3

3 33

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3 3333

333333

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3333

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3

33 33 3

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333

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3333333

333333

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333333333333

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3

Great Basin lavasOther western hemisphere(except Central America)

3 30-39 km Crustal thickness

4 40-49 km 50-59 km

6 60-69 km7 70-79 km

5

55

5

55

5 55

5 5

5

5

5

5

5 55

5

55

5

5

5

555

5 55 55

55

555

5

Figure 9. Ba versus SiO2 in Great Basin lavas compared to lavas from other locales listed in Table 1.‘Main trend’ extends from about 50 wt. % SiO2 and 200 ppm Ba to 75 wt. % SiO2 and 900 ppm Ba.

0

400

800

1200

1600

2000

Ba

(ppm

)

40 45 50 65 70 7555 60

SiO2 (wt%)

Central AmericaCrustal thickness

30 - 39 km40 - 49 km50 - 59 km

Figure 10. Ba versus SiO2 in lavas from Central America listed in Table 1 (see Carr et al. 2003).Line is ‘main trend’ from previous figure.

M.G. Best et al.606


In the Paradise Range, at about 1178 500 W (Figure 2), an older suite (about 26–24 Ma)

of andesite lavas has a ratio of about 0.8 whereas a younger suite (20–16 Ma) has a ratio of

0.6 (John 1992, 2001). These ratios suggest thinner crust than that of our data set, but the

lavas are relatively young, lie in western Nevada, and some are altered; nonetheless, they

imply crustal thinning through time.

0

40

80

120

160

200

240

40 45 50 55 60 65 70 75

Rb

(ppm

)

SiO2 (wt%)

Great Basin lavasOther western hemisphere(except Central America)

30-39 kmCrustal thickness

40-49 km50-59 km60-69 km70-79 km

Figure 11. Rb versus SiO2 in Great Basin lavas compared to lavas from other locales listed in Table 1.‘Main trend’ extends from about 50 wt. % SiO2 and 10 ppm Ba to 75 wt. % SiO2 and 70 ppm Ba.

0

40

80

120

160

200

240

Rb

(ppm

)

40 45 50 55 60 65 70 75

SiO2 (wt%)

Central AmericaCrustal thickness

30 - 39 km40 - 49 km50 - 59 km

Figure 12. Rb versus SiO2 in lava samples from Central America listed in Table 1 (see Carr et al.2003).



K2O at 57.5% SiO2

This diagram is similar to the previous one except the positive correlation of values

continues to thickest crust (Figure 6). Again, there is a slight tendency for the K2O

parameter to flatten above 42 km, but this tendency is a result of only two points, viz.

M and S in crust 40–50 km thick. The Great Basin K2O parameters correspond to a

crust thicker than about 45 km. However, if the parameter for the only seven central

Nevada analyses is excluded in preference to the parameters for the more abundant (26)

Egan Range analyses and for all (376) the Great Basin analyses, a thickness of more

than 70 km is suggested. No difference in the parameter can be seen as a function of

age or longitude of the lavas.

In the Paradise Range (see above), K2O at 57.5% SiO2 in the older lava suite is 2.9 but

2.0 in the younger, again implying crustal thinning through time.

Minimum and average initial 87Sr/86Sr

(Maximum 87Sr/86Sr ratios are not shown here because of a wide scatter.) In the global

compilation of Leeman (1983), the mean 87Sr/86Sr ratio has a relatively strong positive

correlation with crustal thickness whereas the minimum ratios are more scattered. Leeman

(1983) notes that ‘ . . . these correlations are surprisingly strong considering that the age of

continental crust and its composition differ considerably for these arcs’. In our expanded

global data base of 87Sr/86Sr ratios, the correlation with thickness is equally strong and, as

in Leeman’s compilation, the correlation coefficient is smaller for the minimum ratios than

for the average ratios, which have the highest R2 (0.50) value for a best-fit straight line of

any plotted parameter. Figure 7 indicates the crust beneath the area of the Indian Peak

caldera complex, based on seven samples (Hart et al. 1998), was 60–70 km in thickness

whereas Figure 8 suggests a greater thickness.

The 20 87Sr/86Sr ratios for the Egan Range lavas cited by Grunder (1992) and shown in

Figures 7 and 8 are much higher than those from the Indian Peak area, which would

suggest a much greater crustal thickness. However, the crust in this area may be

significantly older (Archean) and more radiogenic (Wooden et al. 1999) than other

circum-Pacific regions in our global data base.

We have no definitive explanation for what would seem to be unrealistic thicknesses of

70–80 km, or more, that are implied for the Great Basin crust in Figures 5–8. Such

extreme thicknesses are known, today, only in the Himalayan region, where collision of

the Indian and Asian continental crusts has stacked them atop one another. However, this

tectonic setting is unlike the continental arc regime of the Great Basin area from the

Mesozoic through middle Cenozoic. In the Great Basin, sparse exposures in horsts reveal

several kilometres of weakly metamorphosed Neoproterozoic and lower Cambrian

miogeoclinal sedimentary rocks that were deposited on the thinned older crust along the

passive rifted margin of the continent (Stewart 1980; Hintze and Kowallis 2009).

Contamination of magma with the pelitic rocks within this dominantly siliciclastic

sequence could have contributed to the high 87Sr/86Sr ratios and to the extreme implied

crustal thicknesses in Figures 5–8. But whether this underpinning of the Great Basin

differs significantly from the crust underlying arcs in the global data base is not readily

answered; extensive blankets of volcanic rocks hide direct clues to the nature of the

underlying crust.

Some compositional parameters are probably asymptotic with respect to crustal

thickness. Strontium isotope ratios might correlate more or less linearly with respect to

crustal thickness until some particular thickness is attained, beyond which the magma

M.G. Best et al.608


composition would then depend more on the proportion of crust assimilated and its

thermal regime. For example, a 100% crustal melt would have the composition of the fusible

portion of the crust, regardless of thickness. In the case of a continental margin with a

thick section of miogeoclinal rocks, 87Sr/86Sr ratios would then be governed more by the

age of the crust and its Rb/Sr ratio. Thus, it is that strongly peraluminous two-mica

granites of Cretaceous age with initial 87Sr/86Sr ratios more than 0.720 in the central Great

Basin appear to be partial melts of the miogeoclinal wedge (Best et al. 1974; Lee and

Christiansen 1983).

Ba and Rb

In unusually thick, thermally equilibrated continental crusts the temperature in the deeper

parts likely exceeds the solidus temperatures of felsic rocks (e.g. Sonder et al. 1987).

Parcels of partial melt too small in concentration to form buoyant diapirs can nonetheless

mix with ascending mantle-derived magmas. Because of the small partition coefficients of

Ba and Rb for the common plagioclase-dominated source rocks in deep continental crust,

these two incompatible elements are enriched in granitic partial melts and in any mantle-

derived magmas mixing with them. Therefore, relatively high Ba and Rb in extruded lavas

indicate long crustal path length (i.e. thick underlying crust). Moreover, these are two of

the most commonly analysed trace elements in the global data set, thus facilitating

comparisons. Because Ba and Rb concentrations in calc-alkaline arc volcanic rocks are a

function of silica content, we present plots against this oxide (Figures 9–12).

In the Ba plot (Figure 9), ‘Other western hemisphere’ individual rocks (not averages of

suites as in previous variation diagrams) are mostly situated on crust 30–59 km thick,

defining a ‘main trend’ that is believed to have resulted from crystal fractionation with

minimal crustal contamination and mixing. Most of the analyses of rocks sited on crust

50–70 km thick lie above the main trend, thus reflecting the enrichment in Ba from mixing

with Ba-rich anatectic partial melts. Significantly, virtually the entire population of Great

Basin rocks also lies above the main trend; the most Ba-rich rocks in the Great Basin

coincide with similarly enriched rocks in the Andes, where the crust is 60–70 km thick.

It should be noted, however, there is a small proportion of rocks sited on 30–59 km-thick

crust that also lie above the main trend. These are rocks from the San Juan field, where the

crust is probably (see below) 48 km thick, and a few, particularly those with greater than

65 wt. % SiO2, from Mexico, where the crust is 30–43 km thick. How are we to account

for these anomalous rocks? Do they weaken the thickness correlation?

Not plotted in Figure 9 are the Ba contents of rocks in the central American arc

(Carr et al. 2003), which are shown separately, for the sake of clarity, in Figure 10.

Comparing the two figures it is immediately obvious that virtually all of the central

America rocks lie above the main trend defined by other western hemisphere rocks. Rocks

sited on the thinnest crust, 30–39 km thick, have the highest Ba contents. This anomaly

may be explained by an unusually high input of Ba-rich sediment into the magma-

generating systems in the central America subduction zone (Plank and Langmuir 1993;

Patino et al. 2000). Ba is presumably removed from the sediments in the shallower parts of

the subduction zone, thus accounting for the negative correlation of Ba content and crustal

thickness. Or, the Ba-rich sediments are located where the crust is 30 km thick. The

anomalously high Ba content of the Mexican rocks may also be the result of contamination

by Ba-rich sediment.

However, it is unlikely that the relatively high Ba contents of the San Juan lavas can be

similarly explained because these rocks lie so far inland, about 500 km east of the eastern



margin of the Great Basin. One explanation is suggested by the fact that the San Juan lavas

plotted (S) in Figures 5–8 consistently lie slightly above the other global data points. This

shift could mean that the assigned crustal thickness of 48 km at the time of Oligocene

volcanism in the San Juan region, based on current geophysical information (Prodehl and

Lipman 1989), is too thin. If true, the Ba content of these lavas is less anomalous.

In the plot of Rb (Figure 11), ‘Other western hemisphere’ arc rocks again define a main

differentiation trend and rocks sited on 50–69 km crust lie above it. However, Andean

rocks on the thickest crust (70 km) are not the highest in Rb and many analyses for rocks

sited on crust 40–49 km thick also lie well above this trend. Because Rb is not a highly

concentrated element in subducted sediment, the anomalous Rb points cannot be

accounted for in the same way as for Ba. The considerable scatter of Rb values is probably

a result of the large variations of Rb in crustal rocks. For example, accreted ocean floor or

island arc terranes have less Rb (and other incompatible elements) than highly processed

upper continental crust of Proterozoic age (e.g. Taylor and McLennan 1985).

Summary

Compositions of middle Cenozoic lava flows indicate the Great Basin crust at this time

was at least 50 km thick and likely 60–70 km. The crust was probably somewhat thinner in

the west than in the east and thinner in the waning stages of the ignimbrite flareup when

crustal extension began.

Sr/Y ratios and depth of magma equilibration

If magmas are generated at depths sufficient to stabilize garnet rather than plagioclase in

the residue, then the Sr/Y ratio of the magma will be relatively high. The minimum depth

for this condition depends on the composition of the source rock and the temperature but is

likely satisfied at depths of perhaps greater than 50 km. A Sr/Y . 40 and Y , 18 ppm

are commonly used to distinguish adakitic magmas that appear to have equilibrated with

garnet (e.g. Castillo et al. 1999). Less than 10% of the intermediate composition rocks we

have analysed from the Great Basin have Sr/Y . 40 and only seven samples have

Y , 18 ppm. The average Sr/Y ratio is 25 ^ 11 and the average Y concentration is

29 ^ 6 ppm. Consequently, there is little chemical evidence that the middle Cenozoic

lavas equilibrated with or fractionated garnet at high pressure in thick continental crust.

However, the lack of a garnet signature cannot be used as an argument against an

unusually thick continental crust because regardless of where the middle Cenozoic lavas

were initially generated they now possess a middle to shallow crustal imprint. Their high

initial Sr ratios (Table 1) suggest contamination with middle to upper crustal materials,

while their position on pseudoternary phase diagrams of Baker (1987) suggest

crystallization pressures between 2 and 5 kb, or 7–17 km deep.

Spatial aspects of ash-flow sheets: Implications for an Altiplano

The unusually thick crust, together with the assumption of isostatic adjustment, implies the

existence of some sort of elevated terrain in the Great Basin area during the ignimbrite

flareup. Because thick crust can underlie rugged mountain belts as well as flat highlands,

we now seek evidence for the character of the topographic relief of the Great Basin area

during the middle Cenozoic ignimbrite flareup.

Ash flows spread outward from their sources as mobile avalanches whose lateral

dispersal is governed by the dynamics of the eruption as well as by the surface topography

M.G. Best et al.610


in the area of dispersal. Subsequent erosive and tectonic processes can modify the primary

shape of a sheet. Examination of large ignimbrite outflow sheets in the Great Basin

provides insights into the nature of the landscape on which the sheets were deposited.

Progressively smoothed terrain

By the late Eocene, the orogenic plateau in what is now the Great Basin that resulted from

contractile deformation during earlier orogenies had been affected by erosion

(e.g. DeCelles 2004). Valleys cut into the terrane of mostly Palaeozoic rocks are marked

by local sediment deposits overlain by the earliest ignimbrites. In the southern Wah Wah

Mountains of southwestern Utah, Abbott et al. (1983) mapped such a palaeovalley filled

with about 700 m lava flows and ash-flow tuffs ranging in age from at least 32 to about

30 Ma. In the same area and westward to the state line, the thickness and distribution of

slightly older ignimbrites suggests EW-trending palaeovalleys that are nearly as deep.

In east-central and central Nevada, Hagstrum and Gans (1989) and Radke (1992) found

evidence for smoothing of palaeotopography by the earliest ash-flow tuffs. In the northern

Toquima Range (Figure 2), McKee (1976a) found that the 35 Ma Pancake Summit Tuff is

200 m thick but within about 6.5 km to the north, west, and south it pinches out completely.

In northeastern Nevada, Henry (2008) documented palaeovalleys as much as 10 km wide

and at least 500 m deep, perhaps as much as 1.6 km.

Continued deposition of ash flows filled palaeovalleys, progressively smoothing the

landscape. As an example, one can compare the substantial dispersion of palaeomagnetic

directions for the Kalamazoo tuff with the tight cluster of directions for the thin tuff of

Clipper Gap (Gromme et al. 1972, Figure 8). The Kalamazoo was deposited at 35 Ma near

the initiation of the flareup on an irregular topography carved into Palaeozoic rocks over

an area of about 120 km2 in the Schell Creek Range (Figure 2; Hagstrum and Gans 1989),

whereas the Clipper Gap was deposited near the end of the flareup, partly in the area of the

Kalamazoo, but in a sheet less than 20 m thick.

The aggregate thickness of the ignimbrite outflow sheets surrounding the Indian Peak,

Caliente, and central Nevada caldera complexes (Figure 2) is of the order of several

hundred metres (Best and Christiansen 1991, Figures 5 and 6), thus effectively eliminating

any relief less than this.

Anderson and Rowley (1975, p. 9) noted that ‘The region of the high plateaus [of

central Utah] was not physiographically distinct from the adjacent Colorado plateaus and

Great Basin during deposition of the lower Cenozoic sequence’. The ignimbrites in this

sequence comprise the three widespread and very large volume monotonous intermediates

emplaced from 31 to 29 Ma. In the earliest investigation of ignimbrites in the southwest

Utah sector of the Great Basin, Mackin (1960, pp. 89, 97), observed that sheets emplaced

about 24 – 22 Ma ‘ . . .were spread over a surface of low relief . . . ’ and as

‘ . . . flattish sheets over a tolerably level plain . . . ’. However, compilation of thicknesses

of outflow sheets from published geologic maps reveals that one of these ignimbrites, the

24(?) Ma Swett Tuff Member of the Condor Canyon Formation, seems to be distributed in

two northeast-trending palaeo-valleys on either side of its eruptive source area.

The apparent absence of palaeovalleys in southwestern Utah, as manifest by ignimbrite

deposits, from about 30 Ma (see above) to about 24 Ma may be the result of flooding of the

landscape by ash flows from the Indian Peak caldera complex to the northwest during its

peak eruptive activity.

Three features of all but the oldest outflow tuff sheets in the central and eastern Great

Basin support the existence of a smoothed depositional landscape. First, thicknesses of



outflow sheets beyond their source calderas generally vary in a systematic manner such

that these variations can be contoured on small scale maps (e.g. Best et al. 1989a, 1995;

Sweetkind and du Bray 2008). Thicknesses of sheets diminish more or less systematically

outwards from their sources, but lobate and eccentric outflow patterns from sources are

common. Second, relatively thin ash-flow sheets such as the 25 Ma Clipper Gap and Nine

Hill tuffs (Best et al. 1989b, pp. 105–110; Deino 1989), crop out over present areas of

.10,000 km2 in eastern Nevada. Third, the generally small between-site variations in

palaeomagnetic direction in individual younger ignimbrite sheets indicates an absence of

substantial topography over which the ignimbrite was draped (e.g. Gromme et al. 1972,

unpublished data 1990–1998; Best et al. 1995). On a hilly depositional surface, non-

horizontal compaction foliations would have increased the between-site variations and

adversely affected the success of palaeomagnetic correlations. Evidence for this fortunate

circumstance consists of the positive results of the palaeomagnetic tilt tests – commonly

known as fold tests–for nearly all of the regionally extensive ignimbrites surrounding the

central Nevada caldera complex (Figure 2; see also Figure 14) with sufficient sites, as well

as for the Nine Hill Tuff, for the most part east of the Carson Sink.

The palaeomagnetic data are consistent with a mostly smooth landscape near the end

of the flareup extending from approximately longitude 1178 W eastward across central and

eastern Nevada and into the high plateaus of south-central Utah, a present area greater than

120,000 km2, equivalent to more than 80,000 km2 prior to Basin and Range extension.

Only approximately 12% of this area was occupied by caldera complexes and single

calderas, which possessed significant topographic relief.

Locally, this smoothed ignimbrite terrain was interrupted by numerous effusions of

lava that today are generally less than 1 km thick and range in composition from rhyolite to

andesite (Best et al. 1989a, Fig. 3). Few major composite volcanoes on the scale of the

modern Cascades formed in the smoothed terrain during the ignimbrite flareup. In the

compilation of 34–17 Ma volcanic rocks of Nevada, Stewart and Carlson (1976) show

three andesitic volcanoes that developed on the margin of the ignimbrite province. Within

the province, remnants of a small dacitic composite volcano now about 10 km in diameter

and 0.3 km high formed in the Seaman Range in southeastern Nevada at about 27 Ma

(du Bray and Hurtubise 1994). In the nearby Pahroc Range, an unusually large pile of

andesitic lava and volcanic debris flows emplaced at about 28 Ma is as much as 1.3 km

thick and extends along the range for about 15 km (Swadley et al. 1995). Major composite

volcanoes did develop in the Marysvale area of central Utah just to the east of the Great

Basin during the middle Cenozoic (Steven et al. 1979). The scarcity of large composite

volcanoes, and accompanying mountainous relief, in the middle Cenozoic ignimbrite

province contrasts with the character of the contemporaneous Southern Rocky Mountains

volcanic field in southwestern Colorado, where the volume of lava and volcanic debris

flows in such edifices exceeds that of ash-flow tuffs (Lipman and McIntosh 2008).

Summary statement: existence of the Great Basin Altiplano

We conclude that during, and especially in the latter stages of, the ignimbrite flareup, a

substantial part of the Great Basin area was a relatively smooth plain capping a sequence

several hundreds of metres thick of essentially conformable ignimbrite outflow sheets.

Locally, in the ‘outflow alley’ between calderas, the plain was dotted by piles of lava.

Together, with the evidence indicated above for an unusually thick crust, which very likely

resulted in an isostatically high elevation, we feel justified in concluding that this part of

the middle Cenozoic Great Basin area was a relatively flat, high plateau, i.e. an Altiplano.

M.G. Best et al.612


This orogenic plateau apparently had not experienced significant regional tectonic

extension and crustal thinning by the time of the ignimbrite flareup. Moreover, it did not

experience such extension concurrent with the ignimbrite flareup, in contrast to the

interpretation of Gans et al. (1989). Had there been significant concurrent extensional

faulting the success of palaeomagnetic correlations of ash-flow sheets and wide dispersion

from their sources described above would have been greatly hindered. The general

absence of sediments between the conformable outflow sheets (Best and Christiansen

1991, Figures 5 and 6) that would have been shed off fault-related uplifts into adjacent

basins is also inconsistent with contemporaneous extension. South and southwest of the

Caliente caldera complex (Figure 2), lacustrine limestone is intercalated within the lower

part of the ignimbrite sequence and lower in that sequence conglomerate of Palaeozoic-

rock clasts becomes more abundant than limestone (Scott et al. 1992; Anderson and

Hintze 1993); these stratigraphic relations imply progressive filling by tuff of a

depositional basin on the southern margin of the middle Cenozoic ignimbrite field and not

synvolcanic extensional faulting.

Eccentricity of source calderas within their corresponding outflow sheets

Rarely do outflow sheets possess anything close to axial symmetry surrounding their

source caldera. In principle, it seems unlikely that pyroclasts would vent uniformly around

the perimeter of the ring-fracture to create a radially uniform outflow sheet. Among others,

Gromme et al. (1997) have documented lobate emplacement of ash flows vented from

adjacent sectors of a caldera system. Topographic influences on ash-flow dispersal as well

as differential post-emplacement erosion of the outflow sheet can also create an eccentric

source-outflow sheet relationship.

The manner in which large to very large volume (100s to 1000s of km3) outflow tuff

cooling units in the central Nevada field were distributed from about 35 to 23 Ma around

their sources in the central Nevada caldera complex (Best et al. 1993, unpublished data

1990–2008) furnish significant insight into the topography of the Great Basin Altiplano.

The location of the source from which two large outflow cooling-unit members of the

35 Ma Stone Cabin Formation were derived has not been ascertained so we cannot tell

how these sheets are distributed around their source. Of the remaining 12 sheets for

which the locations of seven sources are known, nine are distributed asymmetrically,

more to the east than the west of their sources; that is, their eruptive sources are

eccentrically offset toward the west within their corresponding outflow sheets

(Figure 13). Three sheets are distributed essentially symmetrically east and west of

their sources, namely, one of four members of the 26 Ma Shingle Pass Formation and

two of three members of the 27 Ma Monotony Tuff. One of these two symmetrically

distributed Monotony tuffs occurs west of about 1178 W Longitude and is the only one

of the 14 ignimbrite sheets known to occur west of this meridian. Because the

occurrence of this one Monotony sheet so far west is anomalous we made a special

effort to be certain of the correlation of the Monotony Tuff to the east with the tuff of

Miller Mountain (Robinson and Stewart 1984) that is found about 75 km due west of

Tonopah. The latter has an identical precise 40Ar/39Ar age on sanidine as the average

(n ¼ 10) sanidine age of the Monotony and has a modal and whole-rock chemical

composition (30 elements) consistent with the Monotony. Microprobe analyses of biotite

and hornblende phenocrysts in the Miller Mountain and Monotony are indistinguishable

and are distinct from the compositions of these phenocrysts in other tuffs in the central

Nevada field.



It is unlikely that eruptive mechanisms alone – with the one possible exception of the

Monotony sheet just described – would have allowed 75% of the ash flows to spread more

to the east than to the west. Thus, topography and/or erosion are left as factors governing

the eccentricity of most of the tuffs. Erosion of the western parts of the outflow sheets is an

unlikely explanation, because ignimbrite sheets of much the same, but particularly

younger, age range are widely exposed to the west (e.g. Whitebread and Hardyman 1987;

John 1992) and they too should have been eroded away. Therefore, we believe that central

M

PS

StoneCabin

35 WindousButte 31WB

Nev

ada

Uta

h

Austin

Tonopah

117° 115°

39°

40°

37°

38°

113°

Ely

Monotony 27

PancakeSummit

35

(b) Tuffs older than 27 Ma

P

LunarCuesta

25LC

SP

Nev

ada

Uta

h

* ShinglePass26

Austin

Tonopah

Ely

117° 115°

39°

40°

37°

38°

113°

km25 200150100500

km25 200150100500

Pahranagat23

(a) Tuffs younger than 27 Ma

Clipper Gap25

Caliente

Caliente

Figure 13. Distributions of major outflow tuff sheets and their source calderas in the centralNevada volcanic field that surrounds the central Nevada caldera complex, based on our unpublisheddata (see Best et al. 1989b, 1995 for preliminary data). Calderas labelled with abbreviations of thename of the erupted tuff and with age in Ma. (a) ,27 Ma. Asterisk southeast of Austin is our inferredlocation of the source of the tuff of Clipper Gap; no fault-bounded caldera source has been found.(b) .27 Ma.

M.G. Best et al.614


Nevada ash flows could not always surmount a NS-trending topographic barrier positioned

just east of Tonopah and Austin (Figure 2).

In a symmetric manner, little or none of the outflow from sources west of the barrier

(Figure 14) is known for certain east of it. The 22 Ma outflow tuff of Toiyabe (John 1992),

whose caldera source is located about 20 km south of the south end of the Reese River

Valley, extends at least 200 km to the west near the Nevada-California state line. However,

Shawe and associates (Shawe 1998; Shawe and Byers 1999; Shawe et al. 2000) found

none of this distinctive titanite-bearing tuff in their detailed mapping of the Toquima

Range and we have found none farther east in the central Nevada volcanic field.

Nonetheless, the 24 Ma tuff of Arc Dome, whose source is apparently in the southernmost

Reese River Valley, may crop out in the western Toquima Range, where Shawe (1998)

called it the Diamond King Formation. About 50 km farther to the southeast, we found a

thin (10 m) tuff of similar composition as the Arc Dome, except for the presence in the

former of titanite, that has analytically the same 40Ar/39Ar age on sanidine. Without

further data we can only make a tentative correlation with the Arc Dome. None of the other

ignimbrites described by John (1992) west of the Toquima Range have been found in our

investigation of the tuffs in the central Nevada field.

The reconnaissance work of Garside et al. (2002, 2005; also C.D. Henry unpublished

data 2008) indicates some of the tuffs that have sources in the Toquima Range occur far

to the west, to at least Yerington. Thus, source calderas of the Toquima, Toiyabe,

*ADT MJ

Caliente

Tonopah

Ely

Austin

Toyabe Uplift

km25 200150100500

UTAH

CALIFORNIA

ARIZONA

NEVADA

0.706 Line

120°

37°

41°

114°117°

Ignimbrites from CNCC

Figure 14. Western edge of the Precambrian basement corresponding to the ISr ¼ 0.706 line(modified from Wooden et al. 1998) and ‘possible distribution of Toyabe uplift zone’ (darker grey)of (Speed et al. 1988, Figure 22-11). Calderas embedded eccentrically in the western part of themaximum distribution of outflow sheets from Figure 13. Calderas west of the Toquima Range fromJohn (1992) include Arc Dome (AD, 25 Ma) and Toiyabe (T, 22 Ma). Mt Jefferson caldera (MJ,27 Ma) in the Toquima caldera complex from Henry et al. (1996).



and possibly the Arc Dome tuffs are eccentrically positioned to the east within their

corresponding westward extending outflow sheets. Most of their westward dispersal

appears to have been governed by ancient drainage ways on the western sloping margin of

the Great Basin Altiplano (see below). But to the east their dispersal seems to have been

impeded by the same topographic barrier as the one controlling dispersal of ash flows from

calderas to the east in the central Nevada field.

The mirror-image eccentricity of ignimbrite sheets and sources on either side of the

topographic barrier is a fundamental feature of the middle Cenozoic Great Basin

ignimbrite province. Although our current understanding of the correlation and

distribution of ignimbrite sheets on which the existence of this topographic barrier is

based is not perfect, there is, in our opinion, sufficient evidence for it. This barrier was

actually recognized at least as early as the 1980s by Bart Ekren and Jack Stewart, who

suggested its existence at that time to the senior author (MGB).

The location of the barrier is constrained to lie essentially in Monitor Valley between

the Monitor and the Toquima Ranges but to the south swings west roughly through

Tonopah (Figures 2, 13, and 14). Lunar Cuesta and Big Ten Peak ash flows erupted about

25.6 Ma were dispersed eastward from their sources at the south end of the Monitor Range

and 35 Ma Pancake Summit ash flows eastward from their source at the north end of the

range. Ash flows from sources farther to the east were also dispersed eastward. Ash flows

erupted about 29–26 Ma from sources in the Toquima Range travelled westward.

The barrier was not strictly impassable and its effectiveness may have diminished with

time. We have noted that one cooling unit of the 27 Ma Monotony Tuff transgressed the

apparent barrier just north of Tonopah. The 25 Ma tuff of Arc Dome that erupted from a

source west of the barrier is tentatively correlated with tuff to the east of it. Comparison of

the distribution of tuffs that are less than 27 Ma with that of older tuffs (Figure 13) suggests

that with time the ash flows from eastern sources advanced farther west (one cooling unit

of the Monotony Tuff excepted), partially surmounting the topographic barrier. Although

no exposed source for the 24 Ma Clipper Gap ash-flow tuff has been found, thickest

sections of the tuff in the Toquima and Monitor Ranges suggest the source may lie

concealed beneath northern Monitor Valley. A thinner lobe extends almost to Utah,

implying a topographic barrier west of the Toquima Range or special eruption dynamics to

create the asymmetric tuff distribution.

An apparently more serious difficulty with the implied topographic barrier is the fact

that the 25 Ma Nine Hill Tuff, whose concealed source may be in the Carson Sink area,

crops out from the western foothills of the Sierra Nevada across most of Nevada to a little

west of Ely, in many places as a relatively thin, densely welded sheet (Figure 2; Deino

1985, 1989; Best et al. 1989b, Figure R3). Ash flows would not likely have been dispersed

so far to the east with the eruptive source apparently lying well down the western slope of

the Altiplano (see below), unless by this time the effectiveness of the barrier or the slope

reduced in some way. Possibly, crustal extension as documented in the Stillwater Range

just to the east of the apparent source area at about the time of eruption of the Nine Hill

(Hudson et al. 2000) may have modified the topography so as to allow dispersal of the ash

flows eastward. The unusual chemical composition of the Nine Hill magma (Deino 1985)

may have resulted in unusual eruption processes that allowed the ash flow to be so widely

dispersed.

The origin of the barrier is less certain than is its location and apparent effectiveness in

blocking the dispersal of most ash flows. A study of the topography of the Tibetan and

Andean Altiplano continental plateau margins by Masek et al. (1994) may furnish an

explanation for the apparent ash-flow barrier in the Great Basin Altiplano. They point out

M.G. Best et al.616


that the margin of a high continental plateau that faces prevailing storms experiences

substantial orographic precipitation and consequent high erosion, whereas the plateau

interior receives much less precipitation so that there is accordingly little denudation. (For

the Great Basin Altiplano its western margin was exposed to storms coming inland from

the Pacific.) Because of the focused denudation along the margin and consequent isostatic

uplift, a topographic high develops at the lip or break in slope between the plateau interior

and the sloping margin (Figure 15(a)). Numerical modelling of precipitation, erosion, and

isostatic uplift by Masek et al. (1994) simulated development of the topographic profile

observed in the Tibetan and Andean plateau margins (Figure 15(b)). If such a ridge had

evolved on the western margin of the Great Basin Altiplano, its position would have been

near that of our postulated topographic barrier to ash flow dispersal.

The topographic barrier lies near and parallels the western edge of the Precambrian

basement as well as the Toiyabe uplift zone (Figure 14) that is postulated by Speed et al.

(1988) to consist of a chain of domical uplifts, at least some of which are thermal and

diapiric. North of Austin, the zone is manifest by three domical uplifts that expose deep-

seated rocks of moderate metamorphic grade; the dome at Austin is cored by a Jurassic

Denudation

Final topography

Initial topography

Isostatic uplift

Ele

vatio

n (k

m)

0

2

4

6

Distance (km)500 6004003002001000

(b)

Andean plateau

Tibetan plateau

(a)

Ele

vatio

n (k

m)

Distance (km)500 6004003002001000

0

2

4

6

Figure 15. Topographic profiles across the edges of high orogenic plateaus from Masek et al.(1994). (a) Average profiles, based on 100-km-wide swaths (Masek et al. 1994, Figure 3), adjustedto the same horizontal scale, across the Tibetan (solid line) and Andean (dashed lines) plateaus andoriented so that storms approach from the left as they did for the western margin of the Great BasinAltiplano. The upper Andean profile is for the northeast (Beni) part of the Altiplano and the lowerfor the southeastern (Pilcomayo) part. Note the topographic high on the lip of the plateau betweenits interior on the right and its sloping margin. (b) Results of numerical modelling show arepresentative initial topography and the final topography that results from the interaction of heavyorographic precipitation on the initial plateau slope and consequent erosional denudation andisostatic uplift.



granitic pluton. However, Speed et al. (1988) do not specify how the zone is expressed

south of Austin and only note that the uplift zone as a whole lies along the Precambrian

basement edge that was overrun by Palaeozoic–Mesozoic thrust sheets (Figure 1).

Character of the western margin of the Altiplano

In western Nevada, west of the topographic barrier that impeded the distribution of ash

flows, the dispersal of ash-flow tuffs contrasts with that to the east and appears to have

been, at least in part, controlled by palaeovalleys, or ancient drainages, incised in the

western sloping margin of the Altiplano.

The 34 Ma tuff of Cove Mine is draped over a roughly NS-trending topographic high in the

northern Fish Creek Mts. and the southern and eastern parts of Battle Mountain (Figure 2;

Doebrich 1995, Figure 19; S. Gromme, H.C. Palmer, and W.D. MacDonald unpublished

data 1968–2006; John et al. 2008). To the east, the tuff of Cove Mine, which may have

erupted from the Caetano caldera to the south, evidently filled a broad NS palaeovalley.

Part of the outflow from the caldera, the Caetano Tuff itself, travelled westward through an

evident gap in the Toiyabe uplift at least as far as the west side of what is now the Tobin

Range (Burke 1977). In the Toiyabe Range, the 25 Ma Nine Hill Tuff (Deino 1989) can be

observed draped over an erosional surface, which retained appreciable topographic relief,

on the Jurassic granitic pluton exposed east of Austin along US Highway 50

(McKee 1976b).

Garside et al. (2005; see also Henry 2008) summarize evidence for the existence of

Eocene-Oligocene palaeovalleys that extended across the area of the present Sierra

Nevada and were headed in a highland in west-central Nevada. As one example of such

palaeovalleys, they (p. 217) cite the occurrence of the Guild Mine Member of the Mickey

Pass Tuff, which was first described in the Yerington district by Proffett and Proffett

(1976), in the ancestral Yuba River drainage in the Sierra Nevada and indicate its

correlation with the intracaldera lower tuff of Mt Jefferson in the Toquima Range

(Figures 2 and 14). The out-flow length in this palaeovalley is about 210 km, after

compensating for subsequent crustal extension (C.D. Henry Written Communication,

December 2008). Some of these ancient drainage ways were major canyons; for example,

in the Yerington district, Proffett and Proffett (1976) document a palaeovalley 1.6 km deep

filled with Oligocene tuffs and floored by conglomerate.

In describing the mid-Cenozoic topography west of the Precambrian continental

margin and especially west of the Stillwater Range, we can do no better than to quote

Larry Garside:

It is clear that the pre-tuff erosional surface had some relief . . . and there was a well-developedsystem of westward-flowing streams in western Nevada, which headed in a central Nevadahighland . . . These streams apparently flowed in broad palaeocanyons and palaeovalleysdeveloped on the basement. Locally, stream deposits are preserved in the central parts of thesevalleys below the rhyolitic ash-flow tuffs . . . Deposition of these tuffs has preserved thesepalaeovalleys. The source calderas of the outflow ash-flow tuffs . . . are believed to have beento the east, in western or central Nevada . . . The tuffs filled the channels and may havecovered the surrounding higher grounds as well. They were thicker in the channels and thusbecame more strongly welded there, possibly producing slightly lower surfaces directly abovethe channels and allowing subsequently developed stream channels to follow the original, pre-tuff drainages rather closely. (Garside et al. 2003)

Garside’s comment about channel overflow applies especially to the wide distribution of

remnants of the Nine Hill Tuff on the western slope of the Sierra Nevada between 388 and

398 450 N (Deino 1985).

M.G. Best et al.618


An EW-trending belt of complex middle Cenozoic volcanic activity that is

approximately 40 km wide extends across four mountain ranges from the east margin of

the Shoshone Range to the west margin of the Stillwater Range, a present distance of about

100 km. Mapping by Riehle et al. (1972) in the southern Clan Alpine Mountains and by

John (1995) and Hudson et al. (2000) in the southern Stillwater Range showed that this

belt, delineated on Sheet 2 of Stewart and Carlson (1976), is characterized in its western

part by faulted and tilted calderas, subjacent coeval plutons, and ash-flow tuffs, mostly

intracaldera. North of the belt and separated from it by a wide swath of presently exposed

Mesozoic rocks, which may have been a topographic high during middle Cenozoic time, a

wide band of a thick uninterrupted sequence of welded ash-flow tuffs extends from the

Shoshone Range at least to the western margin of the Stillwater Range. This area is

characterized by the complete absence of calderas but is dominated by outflow tuff sheets.

In the New Pass Range and western Shoshone Mountains it is about 57 km wide and

consists of up to 19 separate cooling units with an aggregate thickness of about 600 m in

which intercalated sedimentary rocks are essentially absent (McKee and Stewart 1971).

To the west in the Clan Alpine Mountains the zone of ash-flow tuffs is about 30 km wide

and consists of at least 12 separate tuffs, having a total thickness of 250 m or more

(Riehle et al. 1972; Hudson and Geissman 1991). Farther west in the Stillwater Range the

same zone is about 20 km wide and is abruptly terminated by the western frontal fault of

the Stillwater Range facing the Carson Sink (Hudson and Geissman 1991). Moreover, this

zone of outflow tuffs can be traced farther westward in progressively narrowing

palaeovalleys through Dogskin Mountain and the Diamond Mountains and into the Sierra

Nevada proper at Haskell Peak (Stewart and Carlson 1976, Sheet 2; Brooks et al. 2003;

Henry et al. 2004; Faulds et al. 2005a, 2005b). At Haskell Peak, a stratigraphic section of

nine ash-flow tuffs has been uniquely preserved within an ancient erosional channel cut

into bedrock (Brooks et al. 2003). At least three of these tuff units can be unambiguously

correlated throughout the area just described, which extends as far east as the western

Shoshone Range and spans a present-day EW distance of about 275 km. Notable among

these three tuffs is the Nine Hill, which near the present crest of the Sierra Nevada,

occupies at least two broad channels (Deino 1985). Exposures of the Nine Hill Tuff lie tens

of kilometres farther southwest in the foothills of the Sierra Nevada.

The fact that several mid-Cenozoic ash-flow tuffs spread into the Sierra Nevada from

sources farther east raises the question of the elevation of the east margin of the Sierra at

that time. It is well known that the ancient Pacific shoreline in Late Cretaceous through

Miocene time coincided approximately with the present eastern margin of the Sacramento

and San Joaquin valleys (i.e. the western margin of exposed Sierra Nevada bedrock).

In Figure 16, we show two schematics depicting the implications of two different models

for the history of the Sierra Nevada structural block (mostly a Cretaceous batholith). In the

traditional model (Figure 16(a)), the main uplift occurred in two stages. The first stage

occurred between about 85 Ma to about 50 Ma, represented by exhumation of the

batholithic rocks and terminating at the time of deposition of the Eocene Auriferous

Gravels. Negligible additional uplift occurred until after about 5 Ma when the present

elevation would have been obtained. This model dates back at least to the time of

Waldemar Lindgren (1911) and is based almost entirely on geomorphologic evidence,

reinforced lately by potassium-argon ages of overlying late Cenozoic volcanic

rocks (Wakabayashi and Sawyer 2001, and comprehensive references therein). Part A

of Figure 16 is an attempt to show how this particular uplift history might connect with the

western part of the Great Basin Altiplano; the figure involves a gross (more than sixfold)

linear extrapolation of the slopes of Eocene Auriferous Gravel calculated and projected by



New

Pas

s R

ange

Cla

n A

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e M

tsP

arad

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Ran

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tillw

ater

Ran

ge

Mop

ung

Hill

s

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Rid

ge

Nin

e H

ill

Don

ner

Sum

mit

Ele

vatio

n (1

000

feet

)

Ele

vatio

n (k

m)

Distance east from west margin of Sierra Nevada (km)

Reconstructed distance from west margin of Sierra Nevada (km)

Present elevations and distances

Reconstructed elevations and distances in mid-Cenozoic time

100

Bedrockcrests

Base of tuffat DonnerSummit

Slope of gravel +bedrock crests

11.8 +/–0.5 m/km

Slope of gravel alone11.8 +/–0.5 m/km

(Yeend, 1974)

4.5 +/– 0.5 m/km (Yeend, 1974)

(a)

50040030020000

1

2

3 10

5

0

4

Ele

vatio

n (1

000

feet

)

Rec

onst

ruct

ed e

leva

tion

(km

)

100

Bedrockcrests

Base of tuff at Donner Summit

Hypothetical elevationof altiplano

Mulch et al. (2006)21-22 m/km

(b)

50040030020000

1

2

3 10

5

0

4

Present elevations and distances

Present elevations; distances in mid-Cenozoic time

7.8 ± 0.5 m/km

M.G. Best et al.620


Yeend (1974). Shifting the longitudes of seven of the thick sections of welded ash-flow

tuffs mentioned above to compensate for post-Oligocene extension brings the present

elevations closer to the extrapolated gradient, but the extrapolated gradient would only

intersect the original (Oligocene) longitude of the New Pass Range at about 2.5 km

palaeoelevation, which is at least 0.5 km less than what we envisage for the Great Basin

Altiplano.

A more recent version of the Sierra Nevada uplift history has been derived from

palaeoelevations inferred from measurements of oxygen and hydrogen isotopes in

Cenozoic authigenic minerals, both within the Auriferous Gravels of the Sierra Nevada

(Mulch et al. 2006) and in Miocene and Pliocene basins to the east (Poage and

Chamberlain 2002). These workers concluded that nearly all the uplift of the Sierra

Nevada occurred between the solidification of the youngest Late Cretaceous batholithic

rocks (85–80 Ma) and the deposition of the Eocene Auriferous Gravels (about 50–40 Ma).

Many (U–Th)/He cooling ages measured in apatite from the granitic rocks by House et al.

(2001) also imply early uplift. Mulch et al. (2006) concluded that Eocene elevations in the

northern Sierra Nevada were 1.7–1.8 km, with highest bedrock peaks up to 2.2 km.

In constructing Part B of Figure 16, we have taken the 1.7–1.8 km elevation inferred by

Mulch et al. (2006) and have applied it to Donner Summit, again using a linear

extrapolation from ancient sea level. This leaves about 300–400 m between the Mulch

et al. (2006) estimate and the present elevations at Donner Summit to represent

post-Eocene uplift coupled with an unknown amount of erosion of the bedrock crests.

Quoting Mulch et al. (2006, p. 88): ‘Because relative surface displacements between the

western Basin and Range province and the northen Sierra Nevada are of Miocene and

younger age . . . , we speculate that the Eocene Sierra Nevada formed the western edge of a

high-elevation landscape that characterized large areas of the western United States . . . ’.

Following their speculation, we have taken two-dimensional liberties with the elevations

as well as the original longitudes of the sections of ash-flow tuffs described above. First,

we moved them proportionally westward according to 50% post-Oligocene extension

(McQuarrie and Wernicke 2005), and then increased their elevations so as to represent an

even slope from the Sierra Nevada eastward through the Clan Alpine Mountains to the

Figure 16. Schematic profiles of the western slope of Great Basin Altiplano projected eastward tothe Clan Alpine Mountains and New Pass Range in the western Great Basin. Points represent thebases of Oligocene–Miocene ash-flow tuff sections projected onto a WE profile at the latitude ofDonner Summit. The two elevations shown for the Clan Alpine Mountains correspond to the CCAand NCA sections of Hudson and Geissman (1991). The west margin of the Sierra Nevada structuralblock is placed at the present east margin of the Sacramento-San Joaquin Valley. (a) Estimatesassuming second stage of major uplift starting in late Cenozoic time (e.g. Wakabayashi and Sawyer2001). Slope of gravel with confidence limits extrapolated from Yeend (1974) using method ofLindgren (1911) for western sector of 4.5 ^ 0.5 m/km (20–25 feet/mile) and using Yeend’s estimateof 11.8 ^ 0.5 m/km (60–64 feet/mile) eastward. Distances and elevations shown for present day andreconstructed distances adjusted to 50% extension in the Great Basin. (b) Estimates assuming thatnearly all uplift occurred prior to deposition of Eocene Auriferous Gravels, so that present slopes inthe Sierra Nevada are nearly the same as existed in late Eocene time (Mulch et al. 2006). Overall,slopes of Auriferous Gravels to elevation of 1.7–1.8 km in northern Sierra (Mulch et al. 2006) arelinear interpolations to Donner Summit. Distances eastward from east margin of Sierra Nevadastructural block are adjusted assuming 50% extension in the present Carson Sink and Walker LaneBelt after emplacement of middle Cenozoic ignimbrites. Elevations of tuffs are adjusted to fit ahypothetical slope between Donner Summit and our assumed elevations of ignimbrite sections inStillwater, Paradise, Clan Alpine, and New Pass mountain ranges approaching 4 km.

R



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Elevation (km) Depth (km)

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M.G. Best et al.622


New Pass Range. These adjustments imply overall slopes less than that of the present

Sierra Nevada result and result in a rather even gradient from sea level to nearly 4 km

elevation in middle Cenozoic time. Although the construction of Figure 16(b) involves a

circular argument it is our intent to show the much greater plausibility of the newer model

for timing of uplift of the Sierra Nevada in the context of a middle Cenozoic Great Basin

Altiplano. Moreover, the form of Figure 16(b) is not entirely different from the Andean

Pilcomayo profile displayed in Figure 15(a), with due allowance for the difference

between a hypothetical profile and one based on real topography.

Figure 17 portrays our concept of the Great Basin Altiplano or orogenic plateau

in a west-east cross-section near 38.58 N Latitude. The topographic barrier to dispersal of

ash flows lies at the crest of and between the western slope of the Altiplano and

its smoothed interior. The present crest of the Sierra Nevada lies roughly midway in

that slope.

Demise of the Great Basin Altiplano

If the middle Cenozoic Great Basin crust was 60–70 km thick during the ignimbrite

flareup, as the composition of the lavas suggest, how was the transition made to the

30 ^ km thickness today? With an approximately 50% province-wide extension

(McQuarrie and Wernicke 2005) since the flareup the crust should now be in the order

of 40–47 km thick, assuming a closed system. If we assume an open system, as discussed

in the introduction, then how is it possible to eliminate about 10–17 km of crust?

McQuarrie and Chase (2000) suggest that the Colorado plateaus crust has been

thickened by intracrustal flow of hot ductile rock from the eastern Great Basin, causing it

to thin. They do not specify exactly how much crust was swept eastward, driven by the

topographic head of the orogenic plateau (Altiplano), but the lower part of the range of

10–17 km would appear more reasonable.

Another mechanism to thin the Great Basin crust since the middle Cenozoic is by

delamination of the lower part. Although the details of delamination are complex, Kay

and Kay (1993) conclude that, where the crust is more than about 50 km thick rocks of

basaltic composition in the lower crust consist of a dense eclogitic phase assemblage

that can readily separate from the overlying crust if the underlying lithospheric mantle

has delaminated, or delaminates with the crust. In the Great Basin area, a considerable

volume of mantle-derived basaltic magma must have been intruded into the lower crust

to power the silicic magma generating systems from which tens of thousands of

cubic kilometres of dacitic and rhyolitic ash were erupted. However, geophysical

studies of the eastern part of the Great Basin (Zandt et al. 1995; Hasterok et al. 2007)

reveal that a few tens of kilometres of mantle lithosphere currently underlie the crust,

thus precluding its delamination during the middle Cenozoic. Interestingly, the

subcrustal lithosphere beneath the Andean Altiplano, where the ignimbrite flareup

died within the past few million years, has not been completely removed (Beck and

Zandt 2002).

In a variation of the delamination model that overcomes the barrier imposed by the

mantle lithosphere, the mass of dense garnet-pyroxene-bearing mafic rocks and ultramafic

cumulates that remains after silicic magma generation in the lower crust develop Raleigh-

Taylor convective instabilities that allows it to ‘drip’ into the underlying less dense mantle

(Jull and Keleman 2001; Dufek and Bergantz 2005). Although many factors govern the

instability, reasonable boundary conditions, particularly for thicker crust, predict such

piecemeal thinning on a time scale of several millions of years.



Conclusions

Geologists over the past two decades have hypothesized from different lines of evidence

that the Great Basin area was a high orogenic plateau after contractile deformation and

thickening of the crust during Mesozoic–early Cenozoic mountain building. Although

many are of the opinion that the plateau experienced gravitational collapse and crustal

thinning soon after the deformation, that is, before and during the middle Cenozoic

ignimbrite flareup, we have reason to doubt this timing.

The prodigious ‘supervolcanic’ volume of silicic calc-alkaline magmas erupted during

the ignimbrite flareups in the Great Basin area and in the central Andes since the middle

Miocene are believed to be a direct consequence of an unusually thick crust (Maughan

et al. 2002; de Silva et al. 2006). Because the lower to middle crust was unusually hot, it

was a fertile site for silicic magma generation as vast amounts of mantle-derived magma

were intruded into it. The unique monotonous intermediates, or crystal-rich high-K dacite

tuffs, that constitute a significant proportion of the ignimbrites in both locales (Maughan

et al. 2002) are best explained by evolution in pancake-shaped magma chambers resulting

from magma accumulation within a thickened crust, where compressive tectonic and

gravitational stresses are more or less balanced.

To provide an independent assessment of the crustal thickness in the Great Basin we

examined chemical and Sr-isotopic parameters sensitive to crustal thickness in 376

samples of 42–17 Ma mafic lava flows in the Great Basin and compared these to .6000

analyses of similar but mostly late Cenozoic lavas extruded in continental arcs, where the

thickness of the crust has been geophysically determined. The compositional parameters

clearly indicate that the middle Cenozoic crust in the Great Basin was significantly thicker

than the present 30 km, likely 60–70 km. Isostatically, this thickened crust would have had

a high surface elevation, thus indirectly supporting palaeobotanical and stable isotopic

data claimed by many workers to indicate high elevation during the middle Cenozoic.

The widespread areal extent of ignimbrite sheets, the small variation in palaeomagnetic

directions between sample sites within an individual sheet, and the absence of intervening

sedimentary deposits between sheets imply the depositional surface, especially during the

closing of the flareup, was a smoothed landscape. Altogether, the character of volcanic

rocks is consistent with a high, relatively flat plateau, or Altiplano, in the Great Basin

during the middle Cenozoic. After the ignimbrite flareup the orogenic Altiplano

experienced significant collapse and crustal thinning to 30 ^ km by extensional faulting.

The deep crust could have been thinned by ‘dripping’ of dense residues from the silicic

magma generating systems that created the ignimbrite flareup into the underlying less

dense upper mantle and possibly by intracrustal ductile flowage into adjacent regions.

In south-central Nevada, the systematic eccentric position of source calderas within

most ignimbrite outflow sheets is consistent with a NS topographic barrier on the Altiplano

that extended southward from just east of Austin and governed the dispersal of pyroclastic

flows across the landscape. Ash flows erupted from. the central Nevada caldera complex

were dispersed eastward, whereas ash flows erupted from sources west of the barrier

flowed westward and at least some entered into stream valleys draining the western

margin of the Altiplano. The exact nature and origin of the topographic barrier are

uncertain. Possibilities include a nearby and essentially parallel chain of domical uplifts

and intrusions along the western edge of the Precambrian basement overrun by

Palaeozoic–Mesozoic thrust sheets and a topographic ridge between the western slope of

the Altiplano and its smooth interior that resulted from an interplay of heavy orographic

precipitation on the slope and consequent focused denudation and isostatic uplift.

M.G. Best et al.624


Acknowledgements

This work is an outcome of a larger project on the middle Cenozoic Great Basin ignimbrite provincefinancially supported by the National Science Foundation through grants EAR-8604195, -8618323,-8904245, -9104612, and -9706906 awarded to M.G. Best and E.H. Christiansen. The support ofBrigham Young University and, in the early stages of the project, the US Geological Survey is alsogratefully acknowledged. Olivier Bachmann brought the paper by Dufek and Bergantz (2005) to ourattention. We are indebted to David John and Chris Henry for very constructive reviews of an earlyversion of the manuscript that provided a sharper focus on its content.

Note

†Present address: US Department of Energy, Office of Civilian Radioactive Waste Management,1551 Hillshire Drive, Las Vegas, 89134 Nevada.

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