International Geology Review
Vol. 51, Nos. 7 – 8, July– August 2009, 589–633
The Great Basin Altiplano during the middle Cenozoic ignimbrite
flareup: insights from volcanic rocks
Myron G. Besta, Deborah L. Barra†, Eric H. Christiansena*, Sherman Grommeb,
Alan L. Deinoc and David G. Tingeya
a
Department of Geological Sciences, Brigham Young University, Provo, UT 84602-4606, USA
420 Chaucer Street, Palo Alto, CA 94301, USA; cBerkeley Geochronology Center, Berkeley,
CA 94709, USA
b
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(Accepted 26 February 2009)
Uncertainty surrounds the fate of the orogenic plateau in what is now the Great Basin in
western Utah and Nevada, which resulted from the Mesozoic and earliest Cenozoic
contractile deformations and crustal thickening. Although there is some consensus
regarding the gravitational collapse of the plateau by extensional faulting and
consequent crustal thinning, whether or not the plateau existed during the middle
Cenozoic Great Basin ignimbrite flareup – one of the grandest expressions of
continental volcanism in the geologic record – had remained in doubt. We use
compositions of contemporaneous calc-alkaline lava flows as well as configurations of
the ignimbrite sheets to show that the Great Basin area during the middle Cenozoic was
a relatively smooth plateau underlain by unusually thick crust. We compare analyses of
376 intermediate-composition lava flows in the Great Basin that were extruded at
42 –17 Ma with compositions of .6000 analyses of the late Cenozoic lava flows in
continental volcanic arcs that correlate roughly with known crustal thickness. This
comparison indicates that the middle Cenozoic Great Basin crust was much thicker
than the present ca. 30 km thickness, likely as much as 60 – 70 km. If isostatic
equilibrium prevailed, this unusually thick continental crust must have supported high
topography. This high terrain in SE Nevada and SW Utah was progressively smoothed
as successive ignimbrite outflow sheets were emplaced over areas currently as much as
tens of thousands of square kilometres to aggregate thicknesses of as much as hundreds
of metres. The generally small between-site variations in the palaeomagnetic directions
of individual sheets lend further support for a relatively smooth landscape over which
the sheets were draped. We conclude that during the middle Cenozoic, especially
towards the close of the ignimbrite flareup, this Great Basin area was a relatively flat
plateau, and because it was also high in elevation, we refer to it as an Altiplano. It was
not 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 the
progressively smoothed Altiplano in south-eastern Nevada were derived from source
calderas to the west. Of the 12 major sheets from seven sources, nine are distributed
unevenly east of their sources while the remaining three sheets are spread about as far
east as west of their sources. This eccentricity of sources near the western margin of
75% of the sheets indicates the existence of a NS-trending topographic barrier in
central 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 impeded
by this same topographic barrier. The westward dispersal was controlled in part by
westward-draining stream valleys incised in the sloping flank of the Great Basin
*Corresponding author. Email: eric_christiansen@byu.edu
ISSN 0020-6814 print/ISSN 1938-2839 online
q 2009 Taylor & Francis
DOI: 10.1080/00206810902867690
http://www.informaworld.com
590
M.G. Best et al.
Altiplano in western Nevada and adjacent California; at least one of these ash flows
travelled as far west as the western foothills of the Sierra Nevada. The nature and origin
of the implied topographic barrier are uncertain. It is possible that heavy orographic
precipitation on the western slope of the Altiplano and consequent focused denudation
and isostatic uplift created a NS-trending topographic high at the crest of the western
slope and facing the smoothed Altiplano to the east. The barrier also lies near and
essentially parallel to the buried western edge of the Precambrian basement and to a
zone of thermal-diapiric domes that were spawned in thickened crust as the basement
edge was overrun by late Palaeozoic –Mesozoic thrust sheets.
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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.
591
Previous thoughts on crustal thickness of the Great Basin during the middle
Cenozoic
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.
120˚
117˚
114˚
111˚
109˚
ema
ker
Nevada
Fen
c
Wyoming
41˚
Utah Colorado
er f o l d d t h r u st b el t
an
50
40
39˚
Se
vi
Ely
Tonopah
Caliente
Cedar
City
F
nt
of
Austin
Salt
Lake
City
Roberts M
ountains
Golconda
Reno
50
50
ing
40
Lun
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International Geology Review
ro
37˚
60
Las Vegas
California
50
Arizona
New
Mexico
Thrust faults and folds
Palaeogene crustal
thickness (km)
Coney and Harms (1984)
25 0
50 100 150 200 km
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) of
early Tertiary crustal thickness (Coney and Harms 1984).
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M.G. Best et al.
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
International Geology Review
593
42˚
41˚
40˚
39˚
TQCC
CNCC
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38˚
IPCC
CCC
37˚
36˚
35˚
120˚
119˚
118˚
117˚
116˚
115˚
114˚
113˚
112˚
111˚
110˚
109˚
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. Because
of the small scale of this figure some sample sites overlap as a single point. Also shown are the
outlines 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 cited
in the text, as follows: A, Austin; BM, Battle Mountain; CA, Clan Alpine Mountains; CS, Carson
Sink; 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, Pahroc
Range; 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.
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
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594
M.G. Best et al.
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 thermochronologic 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.
Symbol
A
A
A
A
A
A
A
A
A
A
A
A
A
A
A
A
A
Region
central
north
south
San Pedro
Andes
Andes
Andes
Andes
Andes Tata Sabaya
Andes Bolivia minor cntrs
Andes Ollague
Andes Payachata
Andes Tapungato
Andes Tapungatito
Andes Ojos del Salado
Andes Cerro Alto
Andes Marm-San Jose
Andes Palomo
Andes Tinguirrica
Andes Sordo Lucas
Andes Planch-Pet-Azuf
83
3
17
9
26
25
22
15
18
21
28
30
17
200
54
66
65
Chemical
analyses
(n)
24
3
12
5
16
4
8
5
18
11
25
4
16
110
39
21
Sr isotope
analyses
(n)
1.80
1.50
2.36
1.71
1.93
2.30
1.94
2.27
2.75
1.80
2.20
2.80
2.20
1.50
1.40
1.85
K2O
(57%
SiO2)
0.56
0.54
0.73
0.54
0.58
0.67
0.55
0.80
0.74
0.51
0.73
0.74
0.80
0.63
K2O/Na2O
(57 – 63%
SiO2)
0.7039
0.7042
0.7039
0.7042
0.7048
0.7065
0.7047
0.7049
0.7061
0.7046
0.7063
0.7052
0.7041
0.7050
0.7038
0.7034
87
Min
Sr/86Sr
(, 63%
SiO2)
0.7042
0.7042
0.7041
0.7043
0.7052
0.7067
0.7048
0.7049
0.7066
0.7047
0.7081
0.7060
0.7069
0.7070
0.7043
0.7042
Sr/86Sr
(mean
all)
87
48
51
53
54
59
60
60
61
65
62
70
70
70
60
55
35
70
Thickness
crust
(km)
Leeman (1983)
Leeman (1983)
Leeman (1983)
O’Callaghan and Francis
(1986)
de Silva et al. (1993)
Davidson and de Silva
(1992)
Feeley and Davidson
(1994)
Davidson et al. (1990)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Baker et al. (1987)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988) and Tormey et al.
(1995)
Reference
Table 1. Chemical and isotopic data for continental volcanic arcs of known crustal thickness and for middle Cenozoic lavas in the Great Basin.
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International Geology Review
595
A
A
A
A
A
A
Andes Cerro Azul
Andes
Andes
Andes
Andes
Andes
A
A
A
A
A
K
K
C
C
C
C
C
C
S
Andes Antuco
Andes Chillan
Andes Puyehue
Andes Mocho-Choshuenco
Andes Calbuco
Aleutians Katmai
East Aleutians
Cascades
Cascades Mt Hood
Cascades Crater Lake
Cascades Medicine Lake
Cascades Mt St Helens
Cascades Mt Shasta
San Juan precaldera
Conejos
Tatara-San Pedro
Laguna del Maule
San Pedro-Pellado
Cordon El Guadal
Nev Longavi
A
Symbol
Andes Descabez-Grand
Region
Table 1 – continued
36
255
190
190
57
35
32
112
56
64
20
22
24
22
22
14
23
34
11
43
6
Chemical
analyses
(n)
9
32
14
16
24
32
20
9
6
9
9
8
2
25
6
Sr isotope
analyses
(n)
1.10
1.20
1.30
1.20
1.10
1.40
1.30
1.00
2.50
1.40
1.00
0.70
1.29
1.34
1.60
2.00
1.70
1.90
1.02
1.70
1.62
K2O
(57%
SiO2)
0.7030
0.7028
0.7026
0.7030
0.7028
0.7046
0.28
0.32
0.27
0.31
0.82
0.7038
0.7040
0.7037
0.7037
0.7039
0.7037
0.7039
0.7039
0.0736
0.7038
Min
Sr/86Sr
(, 63%
SiO2)
87
0.35
0.44
0.43
0.18
0.37
0.54
0.62
0.61
0.56
0.46
0.46
K2O/Na2O
(57 – 63%
SiO2)
0.7036
0.7034
0.7051
0.7033
0.7032
0.7040
0.7041
0.7040
0.7041
0.7039
0.7040
0.7039
0.7039
Sr/86Sr
(mean
all)
87
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33
38
40
41
41
39
41
38
48
32
30
30
37
40
41
41
41
41
40
45
45
Thickness
crust
(km)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Ferguson et al. (1992)
Frey et al. (1984)
Davidson et al. (1988)
Feeley et al. (1998)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Hildreth and Moorbath
(1988)
Gerlach et al. (1988)
McMillan et al. (1989)
Lopez-Escobar et al.
(1995)
Hildreth et al. (2004)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Bacon and Druitt (1988)
Mertzman (1977)
Halliday et al. (1983)
Grove et al. (2005)
Colucci et al. (1991)
Reference
596
M.G. Best et al.
14
337
87
330
116
13
246
19
76
50
21
21
50
280
264
42
P
P
M
M
M
M
M
M
M
M
M
M
M
M
L
L
L
Mexico Paricutin
Central America
Guatemala A
Guatemala B
Guatemala-El Salvador
Honduras, Yohoa
El Salvador, Hond,
Nicaragua
Costa Rica Rincon
Costa Rica other
Costa Rica Platanar
Costa Rica Poas
Costa Rica Barba
Costa Rica Irazu
Lesser Antilles Dominica
Lesser Antilles Grenada
Lesser Antilles Saba
27
16
18
261
73
S
P
P
P
20
S
Symbol
Chemical
analyses
(n)
Mexico NW Mex. volc.
belt
Mexico Sanganguey
San Juan postcaldera
Huerto
San Juan syncaldera
Mexico
Mexico, Valley of
Region
Table 1 – continued
23
52
4
3
7
6
22
11
69
40
12
16
49
7
Sr isotope
analyses
(n)
0.35
0.52
0.44
0.62
0.57
0.77
0.48
1.50
1.60
1.30
1.70
1.80
2.60
2.20
0.90
1.20
1.10
0.45
0.53
0.42
0.46
0.40
0.47
0.41
0.40
0.68
K2O/Na2O
(57 – 63%
SiO2)
1.50
1.50
1.40
1.60
1.30
1.50
1.55
1.50
1.50
2.30
K2O
(57%
SiO2)
0.7041
0.7039
0.7038
0.7036
0.7037
0.7036
0.7038
0.7032
0.7029
0.7036
0.7031
0.7037
0.7037
0.7045
0.7032
0.7043
Min
Sr/86Sr
(, 63%
SiO2)
87
0.7045
0.7049
0.7038
0.7037
0.7037
0.7037
0.7039
0.7036
0.7030
0.7040
0.7040
0.7040
0.7040
0.7063
0.7039
0.7047
Sr/86Sr
(mean
all)
87
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36
37
39
41
43
45
33
33
33
33
50
46
38
37
34
33
30
30
48
40
43
48
Thickness
crust
(km)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Leeman (1983)
Leeman (1983)
Defant et al. (2001)
Riciputi et al. (1995)
Leeman (1983)
Wallace and Carmichael
(1999)
Verma and Nelson
(1989)
Nelson and Livieres
(1986)
Wilcox (1954) and
McBirney et al. (1987)
Leeman (1983)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Carr et al. (2003)
Parat et al. (2005)
Reference
International Geology Review
597
Central Nevada caldera
complex
Indian Peak caldera
complex
All Great Basin
Lesser Antilles St Kitts
Kamchatka
Turkey-Iran
New Zealand
Sumatra Java
Japan Kyushu
Japan Honshu
Japan Hokkaido
Nevada Egan Range
Region
Table 1 – continued
L
H
T
Z
U
J
J
J
Symbol
7
10
376
2
16
108
?
78
46
21
14
12
Sr isotope
analyses
(n)
7
173
330
139
234
262
541
595
125
26
Chemical
analyses
(n)
2.90
2.90
2.30
0.60
1.30
2.30
1.50
1.70
1.50
1.00
1.10
3.20
K2O
(57%
SiO2)
1.01
1.19
1.26
1.30
0.37
0.84
0.42
K2O/Na2O
(57 – 63%
SiO2)
0.7071
0.7093
0.7036
0.7031
0.7036
0.7041
0.7039
0.7037
0.7026
0.7026
0.7084
Min
Sr/86Sr
(, 63%
SiO2)
87
0.7091
0.7101
0.7037
0.7033
0.7050
0.7052
0.7047
0.7050
0.7040
0.7037
0.7125
Sr/86Sr
(mean
all)
87
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33
35
50
35
30
28
30
28
Thickness
crust
(km)
Askren (1992) and Hart
et al. (1998)
Barr (1993)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Leeman (1983)
Feeley and Grunder
(1991) and Grunder
(1992)
Askren (1992)
Reference
598
M.G. Best et al.
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International Geology Review
599
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.
600
M.G. Best et al.
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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 subductionrelated 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 subcrustal 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).
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International Geology Review
601
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
602
M.G. Best et al.
7
6
SHOSHONITIC
K2O (wt%)
5
4
HIGH K
3
2
1
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0
45
MEDIUM K
Basalt
Basaltic
andesite
50
55
Andesite
60
Dacite
65
LOW K
70
SiO2 (wt%)
Figure 3. K2O – SiO2 diagram for middle Cenozoic lava flows from the Great Basin. Field
boundaries and rock-suite labels from Le Maitre (1989) except shoshonitic from Ewart (1982).
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 calcalkalic (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
International Geology Review
603
12
ALKALINE
SUBALKALINE
10
Na2O + K2O (wt %)
Latite
8
Basanite
6
Shoshonite
Trachybasalt
4
2
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Trachydacite
0
45
Basalt
50
Basaltic
andesite
Andesite
55
60
SiO2 (wt%)
Dacite
65
70
Figure 4. Total alkalies –silica diagram for middle Cenozoic lava flows from the Great Basin. Field
boundaries 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.
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 intermediatecomposition 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
Average K2O/Na2O (wt%) in andesites (57-63 wt% SiO2)
M.G. Best et al.
1.4
Egan Range
Central Nevada
Indian Peak and
Great Basin >30 Ma
1.2
Great Basin <30 Ma
1.0
ST
0.8
AA
M
A
A
S
A
A
A
A
MA
A
M
A
A A A
M
M A
P MM M
A
M
K
C
P PZ
P M
AK
LH
CC
C
C
0.6
0.4
0.2
A
A
A
A
0.0
20
30
40
50
60
70
Thickness of crust (km)
80
Figure 5. Thickness of the crust plotted against average K2O/Na2O in andesites for the suites of
lava samples listed in Table 1. Arrows on left side of diagram are values of middle Cenozoic lava
suites in the Great Basin.
3.5
Egan Range
3.0
K2O (wt.%) at 57.5 wt.% SiO2
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604
Indian Peak
Great Basin
A
A
M
2.5
S
Central
Nevada
2.0
M
ST
A
AA
A
A
AA
A
M
A
A
A
U
MA A
MM A A
P
J P MMZ P P
MA A
M
A A C
AC
P HM
A C
L K C
J K
L
C
J
CA
A
L
1.5
1.0
A
0.5
A
A
L
0.0
20
30
40
50
60
70
80
Thickness of crust (km)
Figure 6. Thickness of crust plotted against K2O at 57.5 wt% SiO2 on variation diagrams for the
suites of lava samples listed in Table 1. Arrows on left side of diagram are values of middle Cenozoic
lava suites in the Great Basin.
International Geology Review
605
Central Nevada
0.7090
Egan Range
Indian Peak
0.7070
A
A
0.7060
0.7050
S
S
S A A
Z
L
A L
A
A
A
AA
U
M A AM
P AP
JA
T
LM MAM
M
A
MP
M
K HM C
C
K
C
JJ
0.7040
0.7030
A
A
A
AA
AA
A
0.7020
20
30
40
50
60
70
Thickness of crust (km)
80
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 middle
Cenozoic lava suites in the Great Basin.
0.7140
Egan Range
0.7120
Average initial 87Sr/ 86Sr
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Minimum initial 87Sr/ 86Sr
0.7080
Central Nevada
0.7100
Indian Peak
A
0.7080
A
A
S
0.7060
J
Z
L
L
A
A
A
P
PM
J AM
A
PA
A AM
J L MMMM
CM
C
KH K
M
0.7040
U
A
A
A
A
ST
AAA
S
A
A
A AA
0.7020
20
30
40
50
60
70
80
Thickness of crust (km)
Figure 8. Thickness of crust plotted against the average initial 87Sr/86Sr in the suites of lavas listed
in Table 1. Arrows on left side of diagram are values of middle Cenozoic lava suites in the Great
Basin.
606
M.G. Best et al.
2000
Great Basin lavas
Other western hemisphere
(except Central America)
Crustal thickness
30-39 km
40-49 km
50-59 km
60-69 km
70-79 km
Ba (ppm)
1600
7
3
4
5
6
7
1200
6
6 6
6
7
7
66 6
6
800
0
40
7
7
6
6
6
3
6
77
6
7 6
6
3
3
7
4 4
7
7
3
3
3
43
3
3
3
4
44
3
66 4
4
44 44
7
34 3 733443 347 774
4
44 4
4 7
4 44
4
7
3
3
4
4
6
7
4
6
4 47 6
4 44 4444
7 77
774 4 74 4 7 77
4
6
3
3
4
4
6
7
7
6
3 43333
76 6 6 67
6 3 467 4
4
3
4
7
6
3
4 4 5
6 66 7 6 466 66 4 444 4 55 3 4
3 33
43344343333
6 33
6
6
7
6 4343 4 3 34
6
66
3
44
6
66 4453 6 4 66 6 444 43333
4 4
44 4
7
4445445
4 444 4
4
334
4
4 6 3
4
3
6
4
4
4
3
4
5
4
3 4 4 664 46 4
4
5
4
3
3
5
4
3444 3 5
43
44444
4
6
3
4
5
4
4
46 4
5
3
4
6
434 4444 3444
5465 5 5435
44345
4 64
6 44 4
4444
545
45 3
34
6
5
4
53
4 44 4434
44
4443
3 4433
43 5
5344443
4544 33 4 43 33
35
3443
3
443 44
34 5
444
5
54 44
545445434443
433
4447 44433
4
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3
400
45
50
55
60
SiO2 (wt%)
65
70
75
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.
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).
2000
Central America
Crustal thickness
30 - 39 km
40 - 49 km
50 - 59 km
1600
Ba (ppm)
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7 6
6
1200
800
400
0
40
45
50
55
60
SiO2 (wt%)
65
70
75
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.
International Geology Review
607
240
200
Rb (ppm)
160
Great Basin lavas
Other western hemisphere
(except Central America)
Crustal thickness
30-39 km
40-49 km
50-59 km
60-69 km
70-79 km
120
80
0
40
45
50
55
60
65
70
75
SiO2 (wt%)
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.
240
Central America
Crustal thickness
30 - 39 km
40 - 49 km
50 - 59 km
200
160
Rb (ppm)
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40
120
80
40
0
40
45
50
55
60
65
70
75
SiO2 (wt%)
Figure 12. Rb versus SiO2 in lava samples from Central America listed in Table 1 (see Carr et al.
2003).
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.
608
M.G. Best et al.
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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
International Geology Review
609
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).
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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 mantlederived 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 magmagenerating 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
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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
International Geology Review
611
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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
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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, nonhorizontal 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.
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International Geology Review
613
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 Palaeozoicrock 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.
614
M.G. Best et al.
Tuffs younger than 27 Ma
(a)
117°
115°
113°
40°
Austin
Clipper Gap
25
*
Ely
Shingle
Pass
26
39°
SP
38°
Lunar
LC Cuesta
25
Tonopah
P
Pahranagat
23
Nevada
Utah
25 0
(b)
50
100
150
37°
200 km
Tuffs older than 27 Ma
117°
113°
40°
115°
Austin
PS
Ely
Pancake
Summit Stone
35
Cabin
35 Windous
WB
Butte 31
M
39°
38°
Tonopah
Monotony 27
25 0
50
100
150
Utah
Caliente
Nevada
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Caliente
37°
200 km
Figure 13. Distributions of major outflow tuff sheets and their source calderas in the central
Nevada volcanic field that surrounds the central Nevada caldera complex, based on our unpublished
data (see Best et al. 1989b, 1995 for preliminary data). Calderas labelled with abbreviations of the
name of the erupted tuff and with age in Ma. (a) , 27 Ma. Asterisk southeast of Austin is our inferred
location of the source of the tuff of Clipper Gap; no fault-bounded caldera source has been found.
(b) .27 Ma.
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
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120°
117°
615
114°
NEVADA
ine
6L
0
0.7
41°
Toyabe Uplift
Austin
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AD
T
*
MJ
Ely
Tonopah
CALIFORNIA
Caliente
UTAH
Ignimbrites from CNCC
25 0
50
ARIZONA
37°
100 150 200 km
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 the
maximum distribution of outflow sheets from Figure 13. Calderas west of the Toquima Range from
John (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).
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,
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M.G. Best et al.
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
International Geology Review
617
(a) 6
Elevation (km)
Tibetan plateau
Andean plateau
4
2
0
0
100
200
300
400
Distance (km)
500
600
Elevation (km)
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(b) 6
Final topography
4
Initial topography
Denudation
2
Isostatic uplift
0
0
100
200
300
400
Distance (km)
500
600
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), adjusted
to the same horizontal scale, across the Tibetan (solid line) and Andean (dashed lines) plateaus and
oriented so that storms approach from the left as they did for the western margin of the Great Basin
Altiplano. The upper Andean profile is for the northeast (Beni) part of the Altiplano and the lower
for the southeastern (Pilcomayo) part. Note the topographic high on the lip of the plateau between
its interior on the right and its sloping margin. (b) Results of numerical modelling show a
representative initial topography and the final topography that results from the interaction of heavy
orographic precipitation on the initial plateau slope and consequent erosional denudation and
isostatic uplift.
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 deepseated rocks of moderate metamorphic grade; the dome at Austin is cored by a Jurassic
618
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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-developed
system of westward-flowing streams in western Nevada, which headed in a central Nevada
highland . . . These streams apparently flowed in broad palaeocanyons and palaeovalleys
developed on the basement. Locally, stream deposits are preserved in the central parts of these
valleys below the rhyolitic ash-flow tuffs . . . Deposition of these tuffs has preserved these
palaeovalleys. The source calderas of the outflow ash-flow tuffs . . . are believed to have been
to the east, in western or central Nevada . . . The tuffs filled the channels and may have
covered the surrounding higher grounds as well. They were thicker in the channels and thus
became more strongly welded there, possibly producing slightly lower surfaces directly above
the channels and allowing subsequently developed stream channels to follow the original, pretuff 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).
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International Geology Review
619
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
Stillwater Range
Paradaise Range
Clan Alpine Mts
New Pass Range
Mopung Hills
Fireball Ridge
(a)
Nine Hill
M.G. Best et al.
Donner Summit
620
4
Elevation (km)
Bedrock
crests
10
Slope of gravel alone
11.8 +/–0.5 m/km
(Yeend, 1974)
2
Base of tuff
at Donner
Summit
5
1
7.8 ± 0.5 m/km
4.5 +/– 0.5 m/km (Yeend, 1974)
0
0
0
100
200
300
400
500
Distance east from west margin of Sierra Nevada (km)
Present elevations and distances
Present elevations; distances in mid-Cenozoic time
(b) 4
Hypothetical elevation
of altiplano
3
2
10
Bedrock
crests
Base of tuff at Donner Summit
Mulch et al. (2006)
21-22 m/km
5
1
0
100
0
200
300
400
500
Reconstructed distance from west margin of Sierra Nevada (km)
0
Present elevations and distances
Reconstructed elevations and distances in mid-Cenozoic time
Elevation (1000 feet)
Reconstructed elevation (km)
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3
Elevation (1000 feet)
Slope of gravel +
bedrock crests
11.8 +/–0.5 m/km
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International Geology Review
621
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
R
Figure 16. Schematic profiles of the western slope of Great Basin Altiplano projected eastward to
the Clan Alpine Mountains and New Pass Range in the western Great Basin. Points represent the
bases of Oligocene – Miocene ash-flow tuff sections projected onto a WE profile at the latitude of
Donner Summit. The two elevations shown for the Clan Alpine Mountains correspond to the CCA
and NCA sections of Hudson and Geissman (1991). The west margin of the Sierra Nevada structural
block is placed at the present east margin of the Sacramento-San Joaquin Valley. (a) Estimates
assuming second stage of major uplift starting in late Cenozoic time (e.g. Wakabayashi and Sawyer
2001). Slope of gravel with confidence limits extrapolated from Yeend (1974) using method of
Lindgren (1911) for western sector of 4.5 ^ 0.5 m/km (20 – 25 feet/mile) and using Yeend’s estimate
of 11.8 ^ 0.5 m/km (60 – 64 feet/mile) eastward. Distances and elevations shown for present day and
reconstructed distances adjusted to 50% extension in the Great Basin. (b) Estimates assuming that
nearly all uplift occurred prior to deposition of Eocene Auriferous Gravels, so that present slopes in
the 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) are
linear interpolations to Donner Summit. Distances eastward from east margin of Sierra Nevada
structural block are adjusted assuming 50% extension in the present Carson Sink and Walker Lane
Belt after emplacement of middle Cenozoic ignimbrites. Elevations of tuffs are adjusted to fit a
hypothetical slope between Donner Summit and our assumed elevations of ignimbrite sections in
Stillwater, Paradise, Clan Alpine, and New Pass mountain ranges approaching 4 km.
60
40
20
4
2
0
Figure 17.
Depth (km) Elevation (km)
? ? ?
Sierra Nevada
Phanerozoic
batholith
accreted terranes
Proterozoic crust
Utah
Subcontinental lithospheric mantle
Topographic
Ignimbrite flattened surface
barrier CNCC
IPCC
Poorly integrated drainage
Deeply incised, tuff-filled valleys
Conceptual WE cross-section through the middle Cenozoic Great Basin Altiplano at approximately 38.58 N latitude.
Oceanic
lithosphere
Line of section
Dry plateau (3-4 km high)
G r e a t B a s i n “ a l t i pl ano”
Nevada
Wet western slope
California
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M.G. Best et al.
International Geology Review
623
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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 RaleighTaylor 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.
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624
M.G. Best et al.
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.
International Geology Review
625
Acknowledgements
This work is an outcome of a larger project on the middle Cenozoic Great Basin ignimbrite province
financially 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 of
Brigham Young University and, in the early stages of the project, the US Geological Survey is also
gratefully acknowledged. Olivier Bachmann brought the paper by Dufek and Bergantz (2005) to our
attention. We are indebted to David John and Chris Henry for very constructive reviews of an early
version of the manuscript that provided a sharper focus on its content.
Note
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†Present address: US Department of Energy, Office of Civilian Radioactive Waste Management,
1551 Hillshire Drive, Las Vegas, 89134 Nevada.
References
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