The Geology and Geochemistry of Cenozoic Topaz Rhyolites from

Transcription

The Geology and Geochemistry of
Cenozoic Topaz Rhyolites
from the Western United States
Eric H. Christiansen
Department of Geology
University of Iowa
Iowa City, Iowa 52242
Michael F. Sheridan
Donald M. Burt
Arizona State University
Tempe, Arizona 85287
SFEE'It':' FAFE.,
205
© 1986 The Geological Society of America, Inc.
All rights reserved.
All materials subject to this copyright and included
in this volume may be photocopied for the noncommercial
purpose of scientific or educational advancement.
Published by The Geological Society of America, Inc.
3300 Penrose Place, P.O. Box 9140, Boulder, Colorado 80301
GSA Books Science Editor Campbell Craddock
Printed in U.S.A.
Library of Congress Cataloging-in-Publication Data
Christiansen, Eric H
The geology and geochemistry of Cenozoic topaz
rhyolites from the western United States.
(Special paper; 205)
Bibliography: p.
1. Rhyolite-West (U.S.) 2. Topaz. 3. Oredeposits-West (U.S.) 4. Geology, StratigraphicCrenozoic. 5. Geology-West (U.S.) I. Sheridan,
Michael F. II. Burt, Donald M., 1943. m. Title.
IV. Series: Special paper (Geological Society of
America); 205.
QE462.R4C48
1986
552'.2
86-273
ISBN 0-8137-2205-5
Contents
Acknowledgments
v
Abstract
I
Introduction
;................................................ 3
Cenozoic topaz rhyoUtes from the western United States
1. Thomas Range, west-central Utah
2. Spor Mountain, west-central Utah
3. Honeycomb Hills, west-central Utah ........................•.........
4. Smelter Knolls, west-central Utah
5. Keg Mountain, west-central Utah
6. Mineral Mountains, western Utah
7. Wah Wah Mountains and vicinity, southwestern Utah and
southeastern Nevada
8. Wilson Creek Range, southeastern Nevada
9. Kane Springs Wash, southeastern Nevada.........
.
.
Topaz rhyolites in the eastern Great Basin: A summary
10. Cortez Mountains, north-central Nevada
11. Sheep Creek Range, north-central Nevada
12. Jarbidge, northern Nevada
'"
13. Blackfoot lava field, southeastern Idaho
14. Elkhorn Mountains, western Montana
15. Little Belt Mountains, central Montana
16. Specimen Mountain, north-central Colorado
17. Chalk Mountain, central Colorado
18. Nathrop, central Colorado
19. Silver Cliff-Rosita, central Colorado
20. Tomichi Dome, central Colorado
21. Boston Peak, central Colorado. . . . . . . . . . . . . . . . . .. . . . . . . . . . . . . . . . . . . ..
22. Lake City, southwestern Colorado
Topaz rhyolites in Colorado: A summary
23. East Grants Ridge, west-central New Mexico ....•......................
24. Black Range, southwestern New Mexico
25. Saddle Mountain, eastern Arizona
26. Burro Creek, western Arizona
Other "topaz rhyolite" occurrences
3
3
10
13
14
15
15
17
19
19
21
21
23
24
25
26
27
29
30
31
32
34
35
36
37
37
39
41
41
42
Other Cenozoic occurrences, western United States
42
Mexican topaz rhyolites
'. . . . . . . . . . . . . . . . . . . . . . . . . . . .. 42
Precambrian topaz rhyolites
-. . . . . . . . . . . . . . . . . . .. 42
iii
Contents
iv
Principal characteristics of topaz rhyolites
Distribution and ages .....................................•.........
Mode of emplacement
Mineralogy
Fe-Ti oxides and titanite
Feldspar
Mafic silicates
Geochemistry and differentiation trends
Isotopic composition
Magma-tectonic setting
Ore deposits
Beryllium
Climax-type molybdenum deposits
Tin
'
Uranium
Fluorite
Comparison with other types ofrhyolitic rocks
43
43
44
46
46
47
48
50
59
59
61
61
62
63
64
64
64
64
Calc-alkaline rhyolites
Peralkaline rhyolites
Aluminous bimodal rhyolites
Ongonites
66
67
67
Petrogenetic modelfor topaz rhyolites
69
References cited
74
Acknowledgments
. This work was partially supported by U.S. DOE Subcontract #79-270-E from Bendix
Field Engineering Corporation. Additional support was provided by Arizona· State University, the University of Iowa, the U.S. Geological Survey, and the National Aeronautics and
Space Administration (grant NAGW-537). A large number of people have helped with the
new analytical work presented in this report. They include D. McRoberts, M. Druecker, J.
Edie, J. V. Bikun, B. Correa, K. Evans, A. Yates, R. Satkin, K. Hon, D. Lambert, C. E. Hedge,
K. Futa, A. Bartel, D. R. Shawe, J. S. Stuckless, L. Jones, R. T. Wilson, W. Rehrig, G. Goles,
and G. Pine. The technical reviews by W. Nash and W. Hildreth, and editorial assistance of C.
Craddock and L. Gregonis are greatly appreciated. We are also indebted to the authors of
many of the articles cited herein for helpful discussions and for recording the presence of topaz
in the rhyolites they have studied.
v
Geological Society of America
Special Paper 205
1986
The Geology and Geochemistry of Cenozoic Topaz Rhyolites
from the Western United States
ABSTRACT
High-silica, topaz-bearing rhyolites of Cenozoic age are widely distributed across
the western United States and Mexico. Topaz rhyolites are characteristically enriched in
fluorine (>0.2 wt%) and contain topaz crystallized during post-magmatic vapor-phase
alteration. In the United States, their ages span much of the Cenozoic Era (50 to 0.06
Ma). Their emplacement followed or was contemporaneous with calc-alkaline and basaltic magmatism in the Basin and Range province, along the Rio Grande rift, and in
Montana, and coincided with episodes of extensional tectonism in these regions.
Nearly all topaz rhyolites extruded as small, endogenous lava domes with or without lava flows; no topaz-bearing ash-flow tuffs have yet been identified with certainty in
the western United States. Most domes are underlain by a precursory blanket of nonwelded tephra. A few are small, shallowly emplaced intrusive plugs. Volumes of rock
«1 to 100 km 3) in individual complexes composed of 1 to many separate extrusions
suggest that the lavas were erupted from small to medium sized magma bodies.
In addition to topaz, these rhyolites also contain garnet, bixbyite, pseudobrookite,
hematite, and fluorite in cavities or in their devitrified groundmasses. All of these phases
may form during vapor-phase crystallization. Magmatic phenocrysts include sanidine
(ca. Orso), quartz, sodic plagioclase (usually oligoclase), and F- and Fe-rich biotite in
order of usual abundance. Fe-rich hornblende or clinopyroxene occur in a few lavas.
Common magmatic accessory minerals include magnetite, ilmenite, zircon, apatite, allanite, and fluorite. Titanite and REE-rich phosphates have been identified in a few lavas.
The rhyolites crystallized over a wide temperature interval (850 to 600°C, with most at
the lower end of this range) and at variable oxygen fugacities. Titanite-bearing lavas
crystallized above the NNO buffer under oxidizing conditions. Most others appear to
have crystallized near the QFM oxygen buffer. For individual complexes, temperatures
correlate negatively with F-content.
All topaz rhyolites are high-SiOz rhyolites with elevated F, Na, K, Fe/Mg and low
Ti, Mg, Ca, and P. Samples with F concentrations of about 1% have notably lower Si
and higher AI and Na than other topaz rhyolite glasses. Most glasses from topaz rhyolites are metaluminous, but many appear to be slightly peraluminous. Fluorine concentrations in glasses range from slightly less than 0.2 to more than 1.0 wt%, and F / Cl ratios
are high (3 to 10) compared to F-rich peralkaline glasses «3). Topaz rhyolites are
characteristically enriched in incompatible lithophile elements (Rb, U, Th, Ta, Nb, Y, Be,
Li, and Cs). Elements compatible in feldspars (Sr, Eu, Ba), ferromagnesian minerals (Ti,
. Co, Ni, Cr), and zircon (Zr, Hi) are depleted. The REE patterns of most topaz rhyolites
are almost flat (La/YbN = 1 to 3) and have pronounced negative Eu anomalies (Eu/Eu*
= 0.01 to 0.02). Both of these parameters decrease with differentiation as indicated by
increasing F, U, Cs, and other incompatible elements. Titanite-bearing rhyolites have
prominent middle REE depletions. Initial Sr-isotope ratios range from 0.705 to over
0.710.
Geochemical trends at individual complexes are interpreted as arising from fractional crystallization of an initially more "mafic" rhyolite with about 0.2% fluorine.
Extensive fractionation of sanidine, quartz, plagioclase, biotite, and Fe-Ti oxides (in
1
2
Christiansen, Sheridan, and Burt
proportions consistent with their modes) produced much of the trace element patterns.
Zircon, apatite, and a REE-rich phase (allanite, monazite, or titanite) were minor but
important fractionating phases. No liquid-state fractionation is required to explain the
geochemical trends. The high F content and FICI ratios of topaz rhyolite melts may have
modified phase relations so as to produce Na and AI enrichments for highly evolved
magmas.
Topaz rhyolites are intimately related to economic deposits of lithophile elements
(i.e. Be, U, F, Li, and Sn). The volcanic rocks were probably ore- and, in some cases,
fluid-sources for these mineral deposits. In their age, tectonic setting, mineralogy, chemistry, and style of emplacement, topaz rhyolites bear resemblance to the rhyolitic stocks
associated with Climax-type Mo deposits, and some may be surface manifestations of
such deposits.
In their chemical composition and mineralogy, topaz rhyolites are distinct from both
peralkaline rhyolites and calc-alkaline rhyolites with which they may be spatially and
temporally associated. Some of the compositional differences between topaz rhyolites
and peralkaline rhyolites may be attributed to the relative effects of F and CI, on melt
structure and phase relationships in their parental magmas. The F/CI ratios of the melt
or its source may determine the alumina saturation of the magma series. Topaz rhyolites
are distinguishable from calc-alkaline rhyolites by lower Sr, Ba, and Eu, and higher F,
Rb, U; and Th. The usually low La/Yb ratios of topaz rhyolites distinguish them from
both peralkaline and calc-alkaline rhyolite suites. Topaz rhyolites are similar to other
aluminous rhyolites erupted in bimodal associations with basalt in the western United
States. They may be the equivalent of the topaz-bearing ongonites of central Asia.
Topaz rhyolites from the western United States are not the eruptive equivalents of
S-type granites; we liken them to the highly evolved, non~peralkaline,and F-rich anorogenic grnnites. Topaz rhyolites appear to have evolved from partial melts of a residual
felsic granulite source in the lower or middle crust of the Precambrian continent. According to the proposed model, the passage of contemporaneous mafic magmas through
the crust produced necessarily small volumes of partial melts as a result of the decomposition of small amounts of F.;rich biotite that persisted in a high-grade metamorphic
protolith. An extensional tectonic setting allowed these small batches of magma to rise
without substantial mixing with contemporaneous mafic magmas. Subsequent fractionation led to their extreme trace element characteristics.
Topaz Rhyolites
INTRODUCTION
For decades petrologists have been concerned with the role
of volatiles (principally HzO, cOz, sOz, HzS, BZ03, HCI, and
HF) in the genesis and evolution of igneous rocks. Even in fluidundersaturated magmas, volatiles playa key role in determining
the physical properties, crystallization histories, and emplacement
mechanisms of magmas. Studies of their role can be pursued
through theoretical, experimental, and analytical methods. Rhyolites that contain topaz (AlzSi04FZ) appear to form a distinctive
group of silicic lavas with high fluorine concentrations. The occurrence of fluorine-rich volcanic rocks provides the opportunity
to examine the effect of fluorine on the mineralogy, geochemical
evolution, and physical nature of natural rhyolitic magmas. The
intent of this report is to document these geologic and petrologic
characteristics as a basis for ongoing efforts to determine the
origin and evolution of fluorine-rich silicic magmas (e.g., Christiansen et al. 1983a; Ruiz et al. 1985; Kovalenko and Kovalenko
1984; Pichavant and Manning 1984; Dingwell et al. 1985) and to
determine the nature of the ore deposits associated with them
(e.g., Burt et al. 1982; Burt and Sheridan 1981).
The occurrence of topaz lining vugs and cavities in rhyolitic
lavas from Colorado and Utah was first reported in the nineteenth century (Smith 1883; Simpson 1876). More recent investigations, summarized here, have shown that topaz-bearing lavas
are widespread in the western United States and that they contain
other minerals uncommon in silicic volcanic rocks (e.g., beryl,
gamet, pseudobrookite, and bixbyite) that reflect the unique
chemistry and origin of these rhyolites. The lavas also contain
unusually high concentrations of incompatible lithophile elem.ents (e.g., Be, Li, U, Th, Sn, Ta, Rb) and fluorine. Information
about these distinctive rocks is scattered in the literature on the
geology of the western United States.
During a study of uranium mineralization associated with
fluorine-rich volcanic rocks (Burt et al. 1980), it became obvious
that topaz rhyolites are surprisingly similar to one another in their
mode of emplacement, mineralogy, major and trace element
chemistry, and tectonic setting. These features are summarized
here.
CENOZOIC TOPAZ RHYOLITES FROM THE
WESTERN UNITED STATES
The distribution of Cenozoic topaz rhyolites in the western
United States is shown in Figure 1 where the occurrences are
numbered in their order of discussion (generally clockwise
around the Colorado Plateau, starting in west-central Utah). We
have visited most of the localities described in this report (all
Utah occurrences; Sheep Creek Range, Jarbidge, and Kane
Springs Wash, Nevada; Burro Creek, Arizona; both New Mexico
occurrences; Nathrop, Chalk Mountain, and Tomichi Dome,
Colorado; Blackfoot lava field, Idaho; and the Elkhorn Mountains, Montana). Complete results of our new findings are presented here. For each locality, we have summarized pertinent
information about 1) the geologic setting and emplacement of the
3
rhyolites; 2) their petrography and mineralogy; 3) the majorelement, trace-element, and isotopic composition of the lavas; 4)
the nature of ore deposits associated with them; and 5) where
possible, the volcano-tectonic setting as revealed by contemporaneous magmatism and tectonism.
Many of the data on mineralogy, elemental and isotopic
composition, and mineralization are summarized in the figures
and tables in the last part of the report. The reader is referred to
these summaries in the descriptions of each occurrence. Volcanic .
rock classification in this report follows that of the lUGS and is
based on KzO plus NazO and SiOz concentrations (TAS diagram; LeMaitre 1984). Where informative, in parentheses we
have also included the original rock name used by the authors.
1. Thomas Range, west-central Utah
The best-known topaz rhyolites are those from the Thomas
Range in west-central Utah (Figure 2). The occurrence of topaz
in rhyolitic lavas from the Thomas Range has been known for
more than a century (Simpson 1876: 325-326). Because of the
occurrence of topaz in the lavas and the presence of U, Be, and F
deposits in the vicinity, these rhyolites have received considerable
attention in the literature. The most recent comprehensive study
of the area is that by Lindsey (1979, 1982). Turley and Nash
(1980), Bikun (1980), and Christiansen et al. (1984) have examined the petrology of the lavas.
The Thomas Range consists of a group of coalesced lava
flows and domes that were erupted from at least 12 separate vents
6 to 7 Ma (Lindsey 1979). Eruptive episodes, as described by
Bikun (1980), commenced with the emplacement of a series of
pyroclastic flows, minor air-fall sheets, and pyroclastic surge
units, and were terminated by the effusion of rhyolite lavas.
Welded ash-flows occur within the tuffs, but more commonly the
ignimbrite units are thin (3 to 4 m) and unwelded. Fused tuffs
(Christiansen and Lipman 1966) 1 to 2 m thick occur in the
tephra immediately below some lava flows. Flow breccias, consisting mostly of vitrophyre blocks up to 2 m in diameter, are
usually found at the base of the lavas. The breccia grades upward
into flow-banded rhyolite, commonly with numerous lithophysae. The volume of rhyolitic eruptives in the Thomas Range is
about 50 km 3.
Rhyolites from the Thomas Range (the Topaz Mountain
Rhyolite; Lindsey 1982) contain up to 20% phenocrysts, but most
samples are crystal-poor felsites or obsidians. (The mineralogy of
the rhyolites is summarized in Tables 8 and 9). Sanidine (Or4s to
Or6S), quartz and plagioclase (AnlO to Anzs) occur in almost all
samples. Biotite of variable Fe/Mg occurs in most (Figures 31
and 32), whereas spessartine-almandine gamet, ferro-augite, and
Fe-rich hornblende occur as magmatic minerals in a few samples.
Accessory minerals include zircon, fluorite, allanite, Fe-Ti oxides,
and, in at least one instance, fluorine-bearing titanite (Turley and
Nash 1980; Christiansen et al. 1984). Fe-Ti oxide and twofeldspar geothermometry indicate that the Topaz Mountain
Rhyolite crystallized at temperatures between 630 and 790°C at
4
r-·-·
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Figure 1. Locations of known Cenozoic topaz rhyolites in the western United States, The numbers refer
to the localities listed below and described in the text. Open circles without numbers show locations of
some of the peralkaline rhyolites that are approximately contempciraneons with the topaz rhyolites
(Noble and Parker 1974). Also shown are several approximations of the western edge of the Precambrian craton in the western United States. The solid line represents the outcrop limit of Precambrian
rocks (King 1977), the dashed line represents the edge of the craton inferred from Sr-isotope composition ofMesozoic granitoids (Kistler et al. 1981; Armstrong et al. 1977), and the dash-dot line as inferred
by Nd-isotope composition of Mesozoic and Cenozoic granitoids (epsilon Nd = -7; Farmer and
DePaolo 1983, 1984).
1.
2.
3.
4.
5.
6,
7.
8.
9.
Thomas Range, Utah
Spor Mountain, Utah
Honeycomb Hills, Utah
Smelter Knolls, Utah
Keg Mountains, Utah
MineralMountains, Utah
Wah Wah Mountains, Utah
Wilson Creek Range, Nevada
Kane Springs Wash, Nevada
10.
11.
12.
13.
14.
15.
16.
17.
18.
19.
Cortez, Nevada
Sheep Creek Range, Nevada
Jarbidge, Nevada
Blackfoot lava field, Idaho
Elkhorn Mtns, Montana
Little Belt Mtns, Montana
Specimen Mtn, Colorado
Chalk Mountain, Colorado
Nathrop, Colorado
Silver Cliff, Colorado
20.
21.
22.
23.
24.
25.
26.
Tomichi Dome, Colorado
Boston Peak, Colorado
Lake City, Colorado
Grants Ridge, New Mexico
Black Range, New Mexico
Saddle Mountain, Arizona
Burro Creek, Arizona
Topaz Rhyolites
5
LEGEND
D
Quaternary alluvium
Topaz Mtn Rhyolite
(6 Ma)
G
~
Spor Mtn Formation
(21 Ma)
Older volcanic rocks
(30-42 Ma)
Sedimentary rocks
Location in Utah
oI
3km
I
I
I
Scale
!
Figure 2. Generalized geologic map of the southern part of the Thomas Range, Utah (after Lindsey
1979; Christiansen et al. 1984a). Both the Spor Mountain Formation and the Topaz Mountain Rhyolite
contain topaz in rhyolitic lavas. Numbers indicate samples analyzed by Christiansen et aI. (1984).
-I
fairly low oxygen fugacities (QFM; Figure 30; Turley and Nash
1980; Christiansen et al. 1980). Topaz occurs in lithophysal cavities and in the devitrified groundmass of many lava flows. No
magmatic topaz (e.g., in glass) has been identified. Other vaporphase minerals in lavas from the Thomas Range include quartz,
alkali feldspar, beryl, bixbyite, pseudobrookite, hematite, spessartine garnet, and cassiterite.
The average compositions of samples from the Thomas
Range are given in Table 1. The compositions of felsites and
vitrophyres are similar but felsites have higher KINa ratios than
their corresponding vitrophyres. The analyses show high Si, K,
and Na and low Ti, Mg, Ca, and P typical. of topaz rhyolites.
Fluorine ranges from 0.2 to 0.5% in vitrophyres. Most of the lavas
are diopside-normative if calculated on a fluorine-free basis. The
trace element geochemistry of the lavas is typical of topaz rhyolites with generally high and covarying concentrations of U:, Th,
Rb, Li, Be, and Ta (Table 2 and Figure 3). The rare earth element
(REB) distributions in the vitrophyres are similar to other topaz
rhyolites with relatively large negative Eu-anomalies and heavy
RBB (HRBB) enrichments that are correlated with F content and
other chemical indexes of differentiation (REB patterns are illustrated in Figure 40a). Light REE (LREE) abundances decline
with increasing evolution. Sr-isotope ratios (0.707 to 0.712;
Table 3) show that the Thomas Range lavas are moderately
radiogenic, which is consistent with a crustal origin for the parental magmas.
Christiansen et al. (1984) presented a quantitative model for
the geochemical evolution of these lavas based on the fractionation of observed phenocrysts from rhyolitic magmas. Major and
trace element geochemistry demand an interpretation that involves about 70% crystallization of the most mafic rhyolite analyzed to produce the most evolved rhyolite, even though the Si02
content increases by only 2.5% across the series. Both majorelement mass-balance calculations and Rayleigh fractionation
models, using the distribution coefficients of Hildreth (1977) and
Crecraft et al. (1981), suggested that fractionation of sanidine (45
to 50%), quartz (30%), plagioclase (15 to 20%), biotite (3%), and
Fe-Ti oxides (1%) were the principal fractionating phases. In
addition, the observed changes in trace elements (La, Hf, Zr, and
Lu) led to estimates of 0.04% each of allanite and zircon in the
removed mineral phases. The observed P depletion implies that
0.06% of the cumulate mineral assemblage was apatite. Minor
discrepancies for Y, Nb, Ta, and Th could be explained by the
fractionation of extremely small quantities of REB-rich phosphates (not yet observed in the vitrophyres) and titanite.
Crystallization near the minimum in the simple ternary granite system should produce differentiates whose major element
chemistry is not dramatically different from their parent magmas.
6
TABLE 1. MAJOR ELEMENT COMPOSITION OF TOPAZ RHYOLITES FROM THE WESTERN UNITED STATES (IN WEIGHT %)
Thomas Range
ave.
S.D.
ave.
S.D.
ave.
S.D.
ave.
S.D.
Spor Mountain
5
6
S.D.
ave.
75.9
0.10
12.7
0.33
0.03
0.17
76.28
0.17
12.42
1.14
0.04
0.21
76.6
0.10
12.4
0.22
0.01
0.21
76.46
0.13
12.53
1. 03
0.04
1. 09
74.2
0.05
13.5
0.19
1
Si02
Ti02
A1203
2
3
4
Honeycomb
Hills
7
8
ave.
0.82
0.01
0.48
73.66
0.03
14.34
0.25
0.43
0.03
0.06
63.6
tr.
11.1
75.0
0.04
13.6
Fe203
FeO
MnO
1. 07*
0.33
0.28
0.00
0.82
0.29
0.05
0.09
0.07
0.01
0.91
0.24
0.04
0.37
0.12
0.01
0.24
0.01
0.47
0.46
0.04
1. 29*
0.06
0.06
0.02
0.34
1. 90
0.07
MgO
CaO
Na20
0.14
0.80
3.78
0.07
0.42
0.27
0.08
0.77
3.48
0.00
0.07
0.19
0.16
0.85
3.33
0.08
0.08
0.19
0.18
0.96
3.34
0.11
0.35
0.41
0.11
0.61
3.95
0.06
0.10
0.56
0.13
0.34
3.86
0.05
11.1
3.64
0.07
0.62
4.60
K20
P205
F
Cl
4.92
0.00
0.28
0.27
0.00
0.08
4.95
0.00
0.21
0.06
0.49
0.00
0.04
0.01
5.10
0.01
0.29
0.11
0.01
4.91
0.02
0.29
0.38
0.02
4.86
0.00
1.14
0.14
0.52
0.00
0.35
4.76
0.03
0.77
4.00
0.03
8.00
0.01
4.46
0.00
0.95
0.07
----
----
----
----
-------
----
----
----
----
----
----
----
----
0.98*
----
1. Average of 11 rhyolites (Christiansen et al. 1984). 6. Rhyolitic lava (Staatz and Carr 1964).
2. Average of 4 rhyolite lavas (Turley and Nash 1980). 7. Low-silica phase of Honeycomb Hills rhyolite
3. Average of 3 analyses representing 5 rhyolites
(Turley and Nash 1980).
(Shawe 1966).
8. Average of 2 rhyolite lavas (Christiansen et
4. Average of 7 rhyolite lavas (Staatz and Carr 1964).
al 1980).
5. Average of 11 rhyolites (Christiansen et al. 1984).
TABLE 1. (CONTINUED)
Smelter
Knolls
9
ave.
S.D.
Mineral
Mountains
10
ave. S.D.
ave.
Wah Wah vicinity
12
11
S.D.
S.D.
ave.
Si02
Ti02
A1203
75.84
0.04
12.56
loll
0.01
0.39
76.5
0.08
12.7
0.29
0.02
0.11
Fe203
FeO
MnO
0.12
0.99
0.04
0.09
0.09
0.01
0.35
0.28
0.08
0.16
0.11
0.02
1.13*
0.19
0.08
0.03
MgO
CaO
Na20
0.08
0.96
3.79
0.04
0.61
0.20
0.16
0.45
4.30
0.12
0.04
0.16
0.10
0.52
3.90
0.04
0.18
0.56
K20
P205
F
Cl
4.77
0.00
0.72
0.10
0.-08
0.00
0.07
0.03
4.77
0.02
0.41
0.14
0.01
4.83
0.26
----
----
----
----
----
76.1
0.07
12.7
----
0.32
1. 05
0.04
0.29
----
-------
----
9. Average of 4 rhyolites (Turley and Nash 1980).
10. Average of 5 rhyolites from domes (Evans and
Nash 1978).
11. Average of 7 early Miocene rhyolites
(Christiansen 1980; Best et al. 1981).
12. Average of 8 Pliocene rhyolites (Christiansen
1980; Best et al. 1981).
76.2
0.08
12.3
Wilson
Creek
Range
13
ave. S.D.
75.4
0.04
13.2
1.16*
0.27
1.28*
0.19
0.02
0.04
----
0.89*
----
0.09
0.01
----
---0.04
0.09
0.74
3.76
0.06
0.35
0.25
0.04
0.43
4.75
0.02
0.15
0.12
0.12
0.42
3.84
0.09
0.52
3.00
0.20
0.34
4.32
4.53
0.60
4.70
0.15
4.60
<0.05
0.49
0.05
5.20
0.02
0.28
-------
----
0.42
0.12
----
----
----
----
----
----
----
----
----------
76.7
<0.2
13.2
Sheep
Creek
Range Jarbidg e
15
16
0.90
0.02
0.70
----
0.50
0.02
0.22
Kane
Springs
Wash
14
77.6
0.12
12.5
1.56*
----
75.3
0.16
12.9
1. 60*
----
0.02
5.44
----
13. Average of 5 rhyolite lavas (Barrott 1984;
written communication 1985).
14. Kane Spring Wash topaz rhyolite (Novak 1984).
15. Rhyolite lava (Christiansen et al. 1980).
16. Rhyolite lava (Christiansen, unpublished
analysis.)
Topaz Rhyolites
TABLE 1.
Elkhorn
Mountains
18
ave.
S.D.
China Cap
17
ave.
S.D.
Si02
Ti02
A1203
76.4
0.12
12.9
0.53
0.016
0.42
77.3
0.07
13.63
Fe203
FeO
MnO
0.46
0.42
0.06
0.133
0.125
0.003
1. 00
0.49
0.14
MgO
CaO
Na20
0.2
0.52
4.21
---0.048
0.118
K20
P205
F
Cl
4.50
0.01
0.45
0.04
17.
18.
19.
20.
21.
22.
(CONTINUED)
Specimen
Chalk
Mountain
Mountain
22
23
S.D.
ave.
Little Belt Mtns.
20
21
19
76.2
0.02
13.7
76.51
0.03
13.81
0.45
0.16
0.19
0.51
0.27
0.25
0.23
0.20
0.25
0.56*
0.64*
-------
----
----
----
----
----
0.18
0.10
0.03
0.34
3.59
0.04
0.29
0.42
0.06
loll
3.50
0.10
0.50
4.55
0.03
0.29
4.61
0.05
0.42
4.04
0.03
0.37
0.42
0.37
0.84
4.00
4.70
0.01
0.66
0.01
5.00
0.03
4.24
0.00
----
4.56
0.01
-------
4.14
0.01
0.34
0.35
----
0.073
0.001
----
----
-------
0.43*
-------
TABLE 1.
28
0.37
0.05
0.35
75.6
0.08
13.6
0.70
0.30
0.17
0.44
0.27
0.20
0.27
0.06
0.04
1. 61
0.14
0.i4
0.22
3.96
0.15
0.78
3.99
0.10
0.65
3.45
0.07
0.12
0.51
4.10
0.02
0.18
4.30
-------
5.14
0.04
0.14
0.69
0.02
0.06
77.0
0.06
13.0
Fe203
FeO
MnO
0.78
0.15
0.14
MgO
CaO
Na20
K20
P205
F
Cl
----
75.7
----
13.9
----
----
Burro
Topaz
Creek
Rhyoli te
36
37
ave. S:D.
Si02
Ti02
A1203
75.6
0.04
12.7
0.42
0.01
0.16
76.0
0.06
13.0
Fe203
FeO
MnO
0.79* 0.12
1. 0*
----
----
0.09
0.02
----
0.06
MgO
CaO
Na20
0.09
0.71
4.25
0.06
0.12
0.32
0.06
0.60
4.00
K20
P205
F
Cl
4.47
0.01
0.18
0.04
0.37
0.01
0.03
0.01
4.80
0.01
0.30
----
----
----
,
0.97
0.30
0.38
0.10
4.55
74.94
----
14.82
----
----
75.3
0.09
13.1'
-------
----
27
76.6
0.08
12.9
77.5
0.07
12.5
0.06
0.40
0.23
0.01
0.35
0.25
0.07
0.22
0.61
4.26
0.05
0.41
4.35
0.05
0.43
4.20
0.04
0.43
4.5
4.97
0.01
0.55
----
4.54
0.01
4.70
0.00
0.21
----
----
----
75.8
0.08
12.7
0.76*
4.4
0.00
-------
"Effusive" rhyolite (Cross 1886).
Rhyolite vitrophyre (Christiansen et al. 1980).
Devitrified rhyolite (Christiansen et al. 1980).
Average of 2 vitrophyres (Van Alstine 1969).
Devitrified groundmass of rhyolite (Carmichael
1963) .
(CONTINUED)
Tomichi
Dome
31
S.D.
ave.
Silver Cliff
30
29
S.D.
ave.
75.9
0.08
13.6
Si02
Ti02
A1203
23.
24.
25.
26.
27.
0.81
0.11
0.55
-------
----
Average of 6 analyses (Dayvault et al 1984).
Average of 3 rhyolite lavas (Smedes 1966).
Rhyolite sill at Yogo Peak (Pirsson 1900).
Rhyolite stock at Granite Mountain (Witkind 1973).
Rhyolite stock at Granite Mountain (Rupp 1980).
Average of 4 rhyolitic lavas (Wahlstrom 1944).
77.0
0.06
12.6
Nathrop
25
26
24
74.7
0.08
14.5
0.105
1.4
0.0
1.2
7
Boston
Peak
32
S.D.
ave.
0.32 76.2
0.006 0.07
0.25 13.4
0.13
0.01
0.22
0.26
0.04
0.08
1.00*
----
----
Lake
City
33
Grants
Ridge
34
ave.
S.D.
76.2
0.19
13.8
1.22
0.09
0.72
74.7
' 0.07
13.7
1. 29*
0.43
0.07
----
0.87
0.48
0.06
1.18*
----
----
----
0.06
0.01
0.37
0.30
4.96
0.07
0.53
3.30
0.06
0.19
0.17
4.50
0.01
4.67
0.03
0.38
0.07
0.01
----
0.12
0.03
0.09
0.02
0.13
0.34
4.05
0.09
0.05
0.39
0.06
0.38
4.15
0.04
0.10
0.26
----
----
0.78
2.34
0.60
0.99
4.38
0.06
0.14
0.04
0.06
0.05
4.53
0.02
0.52
0.10
0.005
4.85
0.05
0.10
0.25
0.03
0.05
----
----
----
-------
Black
Range
35
S.D.
ave.
----
----
-------
77.7
0.18
12.0
----
0.88
0.02
0.81
----
----
Average of 2 rhyolites (Phair and Jenkins 1975) •
"Pitchstone" (Cross 1896) •
Average of 5 rhyolite glasses (Mutshler et a1. 1985) •
Average of 3 rhyolites (Ernst 1980) •
Average of rhyolite pumice and glassy lava and (Baker
and Ridley 1970) •
35. Average of 3 rhyolitic lavas (Correa 1980) •
36. Average of 9 rhyolite vitrophyres (Moyer 1982) •
37. Modal values of histograms in Figure 35.
28.
29.
30.
31.
32.
33.
34.
Note: All analyses recalculated H20 and CO2 free.
Fluorine an d
chlorine concentrations only reported for vitrophyres.
* FeTotal reported as Fe203.
---- Not reported.
S.D. - 1 standard deviation reported for averages of more than
two samples.
8
TABLE 2. TRACE ELEMENT COMPOSITION (IN PPM) OF TOPAZ RHYOLITES FROM THE WESTERN UNITED STATES
Thomas
Range
Spor
Mountain
3
4
1
2
Li
Rb
Cs
50
423
11.3
37
369
80
1010
56
Be
Sr
Ba
6.5
28
10
22
41
63
Cr
Co
Cu
Ga
2.0
nd
Sc
Y
Zr
2.0
58
129
5
336
201
602
24.4
16*
27*
63*
6
49
126
134
1051
19.6
11
200
nd
3.4
0.4
34
Honeycomb
Hills
7
6
13
15
nd
3.2
48
53*
2.6
116
110
85*
97*
8
590
1400
245
1025
6
28
30
12
115
441
21.3
Mineral
Mtns
10
333
13
5
2
2.6
3.7
120
50
Smelter
Knolls
9
nd
nd
3.2
3.9
105
90
80
65
42
46
90
60
2.4
175
110
Wah Wah
Mountains
12
11
65
634
50
9
11
1.7
0.3
25
45
18
87
1.1
74
140
Wilson
Creek
Range Cortez
14
13
564
21
617
665
71*
156
14
19
32
13
11.1
50
1.4
214* 134
151* 172
75
95
------------------------------------------------------------~--------------------------------------------
Nb
Mo
Sn
109
122*
30
Hf
Ta
Pb
5.5
Th
54.8
21.6
U
64
53
67
37.1
2025
631
56*
40
61. 8
12.4
26.8
151
10000
1370
40
70
30.4
16.7
19000
80
Cs
92
125*
14
90
40
12
7.6
30
79700
75
SM-29-206
SM-61a
Be
52
3.8
65
1. Average of 5 rhyolite vitrophyres (Christiansen
et a1. 1984).
2. Average rhyolite (Turley and Nash 1980).
3. Average of 3 rhyolite vitrophyres (Christiansen
et a1. 1984).
4. Average devitrified rhyolite. Analyses with *
are semi-quantitative emission spectrometry
analyses (Lindsey 1979).
5. Low-silica phase of Honeycomb Hills rhyolite
(Turley and Nash 1980).
6. Rhyolite (Turley and Nash 1980).
4
45
6.1
5.5
31
49
11
145
80
3.4
25
6.7
26
5.6
F 4150
Cl
50
24
8..1
7.7
50
42
37
26
22
57.4
15.0
37
7200
957
4100
16.3
50
59
20.4
47.3
11. 7
4600
1230
Pegmatitic inclusion (Christiansen et al. 1980).
Rhyolite vitrophyre (Christiansen et al. 1980).
Average vitrophyre (Turley and Nash 1980).
Average dome-related obsidian (Evans and Nash 1978).
Average rhyolite vitrophyre (Christiansen et al.
1980; Christiansen 1980).
12. Average rhyolite. Analyses with * are semiquantitative (Keith 1980).
13. Average of 5 rhyolite lavas (Barrott 1984).
14. Average of 2 devitrified rhyolites.
7.
8.
9.
10.
11.
U
Lu
3
Li
Vb
Rb
2
Ta
Th
Tb"
Mn
1
Fe
0.5
Ce
Mg
0.1
Hf
"
p
Nd
La
Tl
Co
Sr
Eu
Figure 3. Enrichment factor diagram showing evolutionary trends in the
rhyolites of the Thomas Range, Utah, (thick line) derived by comparing
an incompatible element-poor and an incompatible element-rich specimen. The samples are from lavas presumed to be cogenetic. Enrichment
factors for the Bishop Tuff (Hildreth 1979) are shown with thin lines and
are similar in magnitude and direction (except for Sc and Sm) to those
for these rhyolite lavas.
Topaz Rhyolites
9
TABLE 2. (CONTINUED)
Li
Rb
Cs
Sheep
Creek
Range
15
Jarbidge
16
380
275
96.5
40
42
Silver Tomichi Boston
Dome
Peak
Cliff
22
20
21
107
289
96
304
7.7
16
176
10
1.3
11
23
36
93
Lake City
23
24
167
580
75
378
Grants Ridge
25
26
Black Burro
Range Creek
27
28
185
680
23
348
9.3
160
100
451
7.5
15.6'
1
12*
6
7.5
19
6
5
126
63
600
7
5
29
2.7
0.42
2
55
Y
Zr
Zn
110
154
Nb
Mo
Sn
43
43
1.7
286
176
125
<10
13
Hg
Hf
Ta
46
1. 98
1. 90
13
75
34
101
83
53
5
1
6
105
192
169
25
35
96
36*
121
14
39
16
25
10
<0.5
4.4
5.1
35
50
U
Cl
177
493
Nathrop
19
11
0.1
Cu
Ga
Sc
F
12.8
<12
<22
Ba
Cr
Co
Pb
Th
Little
Belt
Mtns.
18
14.6
B
Be
Sr
China
Cap
17
2800
60
50
49
24
4450
377
15.3
3375
43
34.0
16.2
1700
7.8
3.3
33
5.4
5.9
1350
1399
15. Devitrified rhyolite (Christiansen
et al. 1980; and unpublished data).
16. Average rhyolite (Christiansen unpublished
data).
17. Average rhyolite (Dayvault et al. 1984).
18. Rhyolite (Rupp 1980).
19. Obsidian (Zielinski et al. 1977; and
Christiansen unpublished data). Christiansen
et a1. (1980) report U (16 ppm) and Th (33 ppm).
20. Average of 3 hydrated rhyolite glasses
Mutschler et al. in press) •
21. Average of 3 rhyolites (Ernst 1980).
The differentiates would, nonetheless, have widely varying trace
element characteristics. Crystal settling seems to be an unlikely
mechanism of crystal fractionation; a more plausible method
would be the fractionation model described by McBirney (1980)
and Huppert and Sparks (1984) that involves wall crystallization,
the generation of a buoyant evolved liquid, and its consequent
upward escape to produce a vertically stratified magma chamber.
In contrast to the nearby Spor Mountain rhyolites, no economic mineralization has been found associated with the younger
Topaz Mountain Rhyolite. Bikun (1980) attributes this lack of
mineralization to lower magmatic concentrations of lithophile
elements and to their retention in the spherulitically-devitrified
rhyolite lavas ofthe Thomas Range.
13
2954
47*
59
40
100
8.2
1000
5200
8.1
3800
21
32.4
12.9
3800
1845
397
23. Average of devitrified rhyolites except U and Th
concentrations from 3 marginal vitrophyres
(Steven et al.1977).
Zielinski (1978) reports
U (40, 26, 43, and 41 ppm). Analyses with *
are semi-quantitative.
25. Vitrophyre. (Christiansen unpublished data).
Zielinski (1978) reported the U concentrations.
26. Devitrified rhyolite (Christiansen unpublished
data) •
27. Vitrophyre (Correa 1980; Christiansen unpublished
data) •
28. Average rhyolite vitrophyre (Moyer 1982; Burt
et al. 1981).
The Thomas Range lies in the central portion of the Deep
Creek-Tintic mineral belt (Shawe and Stewart 1976; Stewart et
al. 1977b), an east-trending zone of basement highs, Cenozoic
volcanic centers and associated mineralization (Figure 4). Like
the Pioche mineral belt to the south, it is expressed as a series of
aeromagnetic highs. Cenozoic magmatism along the belt (Lindsey
et al. 1975; Lindsey 1982) began about 42 Ma with the eruption
of a calc-alkaline sequence of intermediate-composition lavas, ash
flows, and small intrusions. Oligocene (38 to 32 Ma) volcanism
in the Thomas Range region was more silicic and is represented
by several ash-flow tuffs that emanated from collapse calderas.
An 11 m.y.lull in magmatic activity preceded the eruption of the
Spor Mountain Formation, which also contains topaz (see
10
TABLE 3.
Sample No.
Rb
(ppm)
Sr AND Pb ISOTOPIC COMPOSITION OF TOPAZ RHYOLITES
FROM THE WESTERN UNITED STATES
Sr
(ppm)
Age
Ma
Analy stl
SM-6lc
Sl1-62b
SM-29-206
184.6
371.8
433
79.34
58.60
2.5
Thomas Range, Utah
6.740
0.70938
18.38
0.71338
495
0.75141
0.70879
0.71174
0.7071
6.3
6.3
6.3
EHC
EHC
LJ
MR-l
188.4
37.47
Mineral Range, Utah
14.56
0.70616
0.70606
0.5
EHC
WW-6b
WW-9
STC-4
381. 4
626
596
18.19
20.2
10.9
Wah Wah Range, Utah
60.78
0.71635
89.7
0.72752
159
0.75485
0.70599
0.7122
0.7092
12.0
12.0
20.2
EHC
LJ
LJ
DRS-155-62
DRS-149-62
631. 7
627.2
38.38
19.44
Cortez, Nevada
47.71
0.71810
93.65
0.72843
0.70800
0.70862
14.9
14.9
EHC
EHC
IZ-l
358
28
0.7085
13.8
LJ
RT-MC
TjR-l
TjR-2
294
178
352
35
70
16.1
0.7142
0.7106
0.7101
15.4
15.4
15.4
LJ
LJ
LJ
NAT-2
318.8
0.7141
0.7080
29.3
30.8
EHC
72L-47K
281
0.7054
18.5
PWL
HC-8
350
3.7
0.7158
0.7108
27.7
29.0
LJ
15a
15b
529
523
Little Belt Mountains, Mont~na
8.51
179.9
0.8341
0.7094
8.91
169.8
0.8269
0.7092
48.8
48.8
ZP
ZP
3.24
112
Sheep Creek Range, Nevada
37
0.71577
Jarbidge,
24.3
7.36
63.5
Nevada
0.71949
0.71217
0.72397
Nathrop, Colorado
288
0.83433
Lake City, Colorado
7.23
0.7073
Black Range, New Mexico
270
0.82203
Note 1:
EHC-Eric H. Christiansen, analyst at USGS, Denver.
Rb, Sr, and isotope
ratios by mass spectrometry and isotope dilution.
LJ -Lois Jones, analyst at Conoco, Ponca City.
Rb by XRF; Sr and isotope
ratios by mass spectrometry and isotope dilution.
ZP -Zell Peterman, analyst at USGS, Denver.
Rb, Sr, and isotope ratios by
mass spectrometry and isotope dilution (Marvin et al. 1973).
PWL-from Lipman et al. (1978a).
Decay constant for Rb=1.42 x 10-11/y •
below). Scattered centers of rhyolitic and basaltic lavas were 2. Spor Mountain, west-central Utah
formed after about 10 Ma including the eruption of the Topaz
Mountain Rhyolite 6-7 Ma. Although the rhyolites of the
The topaz rhyolite exposed around· the margins of Spor
Thomas Range were not emplaced in a strictly bimodal volcanic Mountain in west-central Utah is related to the largest commerfield with contemporaneous mafic and silicic lavas, they are part cial source of beryllium known in North America. The mineraliof this regional sequence of basalt or basaltic andesite (Figure 5) zation occurs in an altered pyroclastic deposit cogenetic with a 21
and high-silica rhyolite. Mafic lavas with ages of about 6 and 1 Ma rhyolite flow (Lindsey 1982). The most recent studies of the
Ma are exposed at Fumarole Butte 23 kIn to the west (Peterson rhyolite and the mineral deposits include those of Lindsey (1977,
and Nash 1980; Best et al. 1980).
1982), Bikun (1980), and Christiansen et al. (1984).
Topaz Rhyolites
NV
40°
11
UT
CD
c
0
N
CD
:J
39°
0
()
CD
c.
Ol
Cil
E
0
38°'
"-
CD
« ...,,-
Aeromagnetic High
km
,
o
80
160
Mineral Belts
OU
DT
P
01
-Oquirrh - U i nt a
-Deep Creek Tintic
-Pioche
-Delamar-Iron Springs
Figure 4. Index map of eastern Nevada and western Utah showing the location of east-trending
structural, mineral, igneous, and aeromagnetic lineaments (modified from Rowley et al. 1978b). Note
the corresopndence of the locations of topaz rhyolites in Utah (filled circles) with the location of the
major lineaments-the Deep Creek-Tintic (DT), the Pioche-Marysvale or Pioche (P), and the DelamarIron Springs (D!) mineral belts of Shawe and Stewart (1976).
Spor Mountain consists of a block of tilted and intricately
faulted lower and middle Paleozoic sedimentary rocks that are
chiefly carbonates (Figure 2). Numerous, relatively small rhyolite
plugs, dikes, and breccia pipes have intruded the sequence. The
pre-volcanic surface was disrupted by northeast-trending ridges
and valleys, perhaps formed by faulting (Williams 1963). Posteruption basin-and-range faulting has further complicated the
structure making it difficult to estimate the number of vents involved. Lindsey (1979) identified at least three major vents. The
eruptions commenced with the emplacement of a series of ignimbrites, pyroclastic air-fall sheets, and pyroclastic-surge units, and
were terminated by the extrusion of lavas over the tuff. The tuff
contains lithic inclusions of dolomite (altered to fluorite near the
top of the tuff) and older volcanic rocks that were entrained from
the country rock as the pyroclastic material moved through the
vent. Locally, the tuff is absent and the lava rests on Paleozoic
sedimentary rocks, but where present the tuff reaches a thickness
of almost 100 m (Williams 1963; drill core information). A
12
variation diagrams) and by sanidine rims on sieve-textured calcic
plagioclase cores (see,for example, Hibbard 1981).
The Spor Mountain rhyolite is generally phenocryst-rich (20
10-8
to
40%).
Major phases include sanidine (Orso to Or60), smoky
~ (J
6-7 )}. ~a
quartz, plagioclase (AnlO to An13), and aluminous Fe- and F-rich
~ ~ (f)X5.3,6.0
4.~
biotite (Figures 31 and 32). Magmatic accessories include Fe-Ti
21 ~
:,,, 1.0
<
~::..
oxides, zircon, fluorite, and allanite. The groundmass of felsitic
c
samples
is granophyric, probably as a result of the thickness and
3.4
6.1
~
z
slow
cooling
of the flow. The groundmass consists of alkali feld110.3 -Delta
spars,
silica
minerals,
fluorite, topaz, and biotite or hematite.
0.4
miles
40
o
Topaz also occurs in miarolitic cavities that surround the mafic
~
60
o
km
inclusions. Two-feldspar geothermometry indicates the pheno.tj{~:
,
::/12:';, 0.1,0.2
390
crysts in the rhyolite equilibrated at 680°C (Table 4) and the
0.9 "',!
- Filmore
composition of the biotites suggest equilibration near the QFM
2.3 ··0~0.4
oxygen buffer (log f0 2 = -18 to -19; Figure 30; Christiansen et
al.
1980). The mafic inclusions contain plagioclase (AnSO-30
:
_. 0.5
'~"
Il Ab4S-600r4-11), augite (Ca37Mg3sFezs), titaniferous magnetite,
0.3 .~.!:" Cove Fort
24.
and ilmenite in a quench-textured matrix of needle-like pyroxene.
Sanidine rims (Or60Ab34An6) on plagioclase suggest tempera0.5- .8[1' 1.0 u&~8~SV181e
tures
of 910°C, while co-existing microphenocrysts of Fe-Ti ox20-22~'f::
23
7.9
1.1
lQ
12.
•
o .'!i 13
~
__
ides
yield
equilibration temperatures of 1100 to 1200°C and f0 2
7.6 b.
8
:
eo
'
1
0:
9-11
c;;,;: 23
21.1
5 :':~".,
near the QFM buffer.
<I
~
50' :';'~
;.:: 7 2' 6.4
. 0 ·f:-·
The average major-element composition of the Spor Moun~:~"
•
5.4 °
I-I~~'d_---::=----~t_-------_;__ 38
tain rhyolite is presented in Table 1. Although generally similar in
its characteristics to other topaz rhyolites, the Spor Mountain
lavas contain lower SiOz (72 to 74%) and higher AlZ03 (13 to
Figure 5. Distribution of Miocene and younger rhyolite or mafic lava
associations in western Utah. Rhyolites are unpatterned; ages (in Ma) are 14%) and FeZ03 (1.2 to 1.5%) than others. In spite of slightly
in bold-face numbers. Mafic flows are stippled; ages are in smaller lower SiOz contents, the low concentrations of P, Ti, and Mg,
numbers (Best et aI. 1980; Best et aI. 1985; Evans and Steven 1982); and high concentrations of F and Na suggest an "evolved" combasalt (0), andesite-trachyandesite (x), potassic mafic lavas (open position, an observation corroborated by the extreme enrichtriangles).
ments of incompatible trace elements. Relative to the rhyolites
from the Thomas Range, samples of the Spor Mountain rhyolite
plot nearer the Ab apex of a normative Q-Ab-Or ternary diagram
central welded zone is developed in the thicker sections of tuff (Figure 36). This is consistent with their F-rich compositions and
(Williams 1963). Bikun (1980) has suggested that in places it Manning's (1981) experiments. His work shows that increased F
may be a result of the accumulation of vitrophyric bombs instead displaces residual melts in the water-saturated granite system toof compactional welding. Although the tuff has been called ward the albite corner. Although devitrified samples are generally
"water-laid" because of its stratified appearance, no textures in- corundum normative, the glasses have normative diopside sugdicative of an epiclastic origin have been recognized (Bikun gesting that alkalies were lost during devitrification (Christiansen
1980). However, in the valley that separates Spor Mountain from et al. 1984). The trace element composition (Table 2) of the
the Thomas Range,a uranium-mineralized lens of tuffaceous vitrophyres bears out the distinctive nature of the Spor Mountain
sandstone and limestone conglomerate occurs beneath the tuff at rhyolite. It contains very high concentrations of U, Th, Rb, Ta,
the Yellow Chief uranium mine (Lindsey 1978). A breccia zone Nb, Be, Li,Y, Ga, Pb, and Sn. A typical REE pattern is shown in
is exposed at the base of the rhyolite lava in several of the open- Figure 40b. It shows a deep Eu-anomaly, indicative of feldspar
pit beryllium mines. This breccia is interpreted as an over-ridden fractionation and the relatively flat REE pattern typical of many
apron of talus that accumulated at the front of the moving flow topaz rhyolites. Based on vitrophyre-felsite comparisons, sub(Bikun 1980). The rhyolite lava has a maximum known thickness aerial crystallization resulted in the loss of 25 to 50% of the F and
of 300 m (Williams 1963). Dark mafic inclusions with globular U in the glasses. Be losses may have been on the order of 75% durand contorted lensoidal shapes are found in the lava. Their tex- ing devitrification (Bikun 1980). Many other elements (including
tures and chemistry led Christiansen et al. (1981) to suggest that a Rb) do not appear to have been released by devitrification.
more mafic magma was injected into the magma chamber of the
The tuff beneath the Spor Mountain rhyolite is the host for
Spor Mountain rhyolite shortly before its eruption. The mafic major Be and minor U, F, Li, Mn, Zn, Nb, and Sn mineralization
(trachyandesite) inclusions themselves show signs of pre-eruption (Lindsey 1977; Bikun 1980). The Be-mineralization is strongest
mixing with the rhyolite in their bulk chemistry (linear trends on in· the upper few meters of the tuff where BeO concentrations
.1140'
113°
112°
I-!I+--------+-----~~-r__
40°
1.0~2.5e
~
Topaz Rhyolites
13
TABLE 4. GEOTHERMOMETRY OF TOPAZ RHYOLITES FROM THE WESTERN UNITED STATES
Location
Method
Temperature
Range (OC)
Turley and Nash 1980
Christiansen et al. 1980
5
E.H. Christiansen
unpublished data
2
685
3
1
620 - 770
650 - 780
3
3
Evans and Nash 1978
Evans· and Nash 1978
2 Feldspar
650
1
Fe-Ti oxides
830
1
Apatite-bio
590
1
E.H. Christiansen
unpublished data
E.H. Christiansen
unpub l i shed data
2 Feldspar
650
1
2 Feldspar
2 Feldspar
Fe-Ti oxides
690
630
725
Spor Mountain, UT
2 Feldspar
680
Honeycomb Hills, UT
2 Feldspar
605
Smelter Knolls, UT
2 Feldspar
Fe-Ti oxides
630
665
Mineral Mountains, UT
2 Feldspar
Fe-Ti oxides
Wah Wah Mountains, UT
Nathrop, CO
Reference
4
1
1
Thomas Range, UT
Chalk Mountain, CO
- 790
n
-
-
690
E.H. Chr istiansen
unpublished data
Notes:
2 Feldspar geothermometry using equations of Stormer (1975)
recalculated 100 bars.
Fe-Ti oxides calculated using recalculation method of
Stormer (1983) and solution model of Spencer and Lindsley (1981).
Apatite-biotite
geothermometer after Ludington (1978).
reach 1% (Griffitts and Rader 1963; Lindsey 1977, 1982; Bikun region is summarized by Lindsey et al. (1975), Shawe (1972),
1980). Mineralization is associated with dolomite clasts altered to and Lindsey (1982) and is reviewed in the section describing the
fluorite and opal in association with feldspathic alteration of the Thomas Range. The episode of rhyolitic volcanism at Spor
matrix (Williams 1963; Lindsey 1977). A thick zone containing Mountain maybe related to the initial development of the Basin
Li-bearing clays underlies the ore zone. Bikun (1980) and Burt . and Range province in a back- or intra-arc setting oflithospheric
and Sheridan (1981) suggest that the ore deposits were formed as extension (see, for example, Zoback et al. 1981; Eaton 1984a).
Be, V, P, and other elements were released from the lava by The evidence for magma mixing suggests that magmatism may
granophyric crystallization and then concentrated in the upper have been bimodal. The episodic recurrence of silicic magmatism
part of the tuff by meteoric fluids at fairly low temperatures. along the Deep Creek-Tintic belt suggests that it may be a proBeryllium ore (bertrandite) coprecipitated with fluorite formed found flaw in the continental lithosphere-perhaps an intraconby reaction of this fluid with the carbonate lithic inclusions. Some tinental transform of the type proposed by Eaton (1979);
support for this model is provided by V-Pb ages of opal nodules Regardless of its origin, the Deep Creek-Tintic belt is very similar
in the beryllium tuff. One of these zoned nodules has ages that to the more southerly Pioche belt in its characteristics and history.
decrease outward from 20.8 ± 1.0 Ma to 8.2 Ma (Ludwig et al. Mid-Cenozoic igneous activity commenced about 10 Ma earlier
1980). These results suggest that the mineralized nodules formed along the northern belt (Stewart et al. 1977b), but the later develat the same time as the rhyolite erupted, but continued to grow opment of both regions was very similar and characterized by the
(perhaps episodically) by the deposition of opal from ground eruption of bimodal suites of mafic lavas and topaz rhyolites after
water. In contrast, Lindsey (1977) and Williams (1963) sug- about 22 Ma.
gested that hydrothermal fluids rose along fractures from an unexposed pluton, intersected the porous tuff and deposited Be and 3. Honeycomb Hills, west-central Utah
P below the lava. On Spor Mountain, uraniferous fluorite without associated beryllium minerals occurs in breccia pipes formed
Honeycomb Hills, the western-most topaz rhyolite occurby partial venting of the rhyolitic magma. As mentioned above, a rence along the Deep Creek-Tintic trend, lies 30 km west of the
sedimentary V -deposit (Yellow Chief mine) occurs in an epiclas- Thomas Range. Studies of the geology of the Honeycomb Hills
tic conglomerate that locally underlies the tuff.
and t~eir rhyolites include those of Turley and Nash (1980),
The volcanic history of the Spor Mountain/Thomas Range Hogg (1972), McAnulty and Levinson (1964), Shawe (1966),
14
and Erickson (1963). The Honeycomb Hills have attracted attention primarily as a result of low-grade beryllium and rare-alkali
mineralization in tuffs associated with a small topaz-rhyolite
dome complex emplaced 4.7 Ma (Lindsey 1977; Turley and
Nash 1980).
Two rhyolite domes with a total volume of less than 0.5
km 3 form the Honeycomb Hills. Eruption of the domes was
preceded by the deposition of a lithic-rich tuff. The tuff is slightly
altered and contains macroscopic fluorite. The rhyolite has a
basal vitrophyre that is exposed locally but it generally consists of
felsitic rhyolite with highly contorted flow foliation. The rhyolite
apparently intrudes unrelated older (Miocene or early Pliocene)
trachyandesitic (shoshonitic) and dacitic lavas and tuffs (Hogg
1972). Paleozoic sedimentary rocks were encountered at about
50 m in a hole drilled between the two hills (McAnulty and
Levinson 1964).
The rhyolites from the Honeycomb Hills contain up to 40%
phenocrysts of sanidine (Orso to Or60), smoky quartz, plagioclase
(AnlO), and F- and Fe-rich biotite (Turley and Nash 1980; Figures 31 and 32). Two-feldspar geothermometry yields a temperature of 605°C (Turley and Nash 1980). Magmatic accessories
include fluorite and traces of Fe-Ti oxides. Topaz, fluorite, and
biotite occur as devitrification products within the groundmass of
the rhyolites. The lava also contains globular inclusions of topazand fluorite-rich material. The textures in these inclusions range
from pegmatitic to aplitic. They may represent fragments of a
cogenetic granite pluton at depth. Inclusions of the metamorphic
basement (quartzite, metaconglomerate, and meta-arkose) and
inclusions of the surrounding mafic lavaS also occur in the
rhyolite.
The compositions of several samples from Honeycomb Hills
are given in Table 1. Fluorine contents in vitrophyres range from
0.95 to 1.2% (Christiansen et al. 1980). The rocks are slightly
richer in Al and Na, and have less Si than many topaz rhyolites
but are similar to those from Spor Mountain, Utah. These compositional characteristics probably resulted from the effect of high
F in the magma. As noted above, Manning (1981) has shown
that the effect of F is to push the water-saturated minimum in the
granite system toward the albite apex as the quartz field expands,
producing more sodie and aluminous residual melts. The single
major-element analysis reported by Turley and Nash (1980) is
very unusual (Table 1). It contains 11.1% CaO, 8.0% F, and only
63.6% SiOz-reportedly due to large amounts of fluorite in the
groundmass. Perhaps this analysis represents one of the globular
inclusions described earlier. Some trace-element analyses are also
available for samples from the Honeycomb Hills (Table 2). Consistent with their higher fluorine contents, the Honeycomb Hills
rhyolites appear to be more "evolved" than those of the Thomas
Range. Analysis 5 is the low-silica sample analyzed by Turley
and Nash (1980) and analysis 6 is for a "normal" rhyolite for
which no major element data have been published. The other
trace-element analyses (7 and 8) represent a pegmatitic inclusion
and a rhyolite respectively (Christiansen et aI. 1980). These rhyolites contain high Rb, Li, and U but are depleted in Th (about 30
ppm) compared to other topai rhyolites, especially if compared
to those with approximately 1% F such as the Spor Mountain
rhyolite. The silica-poor sample contains 151 ppm U and high
Mo, while the pegmatitic inclusion contains anomalous Li (590
ppm) and Sn (65 ppm). REE patterns for the rhyolite are similar
to those from Spor Mountain (Figure 40c).
Low grade Be, Li, Cs, and Rb mineralization occurs in the
pyroclastic deposit beneath the western dome. The mineralization
occurs in aim thick zone about 1 m below the base of the lava
and is reported to have been developed when ascending magmatic fluids were blocked by the impermeable rhyolite (McAnulty and Levinson 1964).
4. Smelter Knolls, west-central Utah
The youngest (3.4 Ma by K-Ar method; Armstrong 1970;
Turley and Nash 1980) of the topaz rhyolites identified on the
Deep Creek-Tintic trend consists of a rhyolite dome and flow
complex at Smelter Knolls located about 25 km west-northwest
of Delta. The geology of Smelter Knolls is described by Turley
and Nash (1980).
The rhyolite complex atSmelter Knolls is 5 km in diameter
and contains about 2.2 km 3 oflava. Basal vitrophyres and flow
breccias are exposed locally. No pyroclastic deposits related to
the lava are exposed. The lavas were apparently erupted through
thick alluvial deposits related to the development of the basin and
range topography of the area.
The rhyolite from Smelter Knolls generally contains 15 to
20% phenocrysts set in a devitrified flow-banded matrix. Quartz,
sanidine (Orso to Or6S), plagioclase (AnlO to Anzo), Fe- and
F-rich biotite (Figures 31 and 32), and Fe-Ti oxides (ilmenite is
rare) occur in all samples. Accessory phases include zircon, fluorite, and allanite. Topaz occurs in the devitrified groundmass and
along with hematite in miarolitic cavities. Fe-Ti oxide and twofeldspar geothermometry indicate that the phenocrysts crystallized at low temperature (660 to 685°C; Table 4) and at f0 2 near
the QFM oxygen buffer (Figure 30; log f0 2 = -19.9 bars).
The average composition of four samples analyzed by Turley and Nash (1980) is presented in Table 1. The rocks are very
similar to other topaz rhyolites. Fluorine ranges from 0.65 to
0.78% in vitrophyres; a single felsite shows the common fluorine
depletion that accompanies devitrification (F = 0.5%). Devitrification also produced depletions of U, CS,and Sb. As expected, the
rhyolites are enriched in incompatible lithophile elements (Table
2). The REE patterns (Figure 40d) are also similar to those from
other topaz rhyolites with low La/LuN (1.60), La/CeN (1.06)
and Eu/Eu* (0.02).
The two major rhyolite hills at Smelter Knolls are separated
by a normal fault that stretches northward for about 8 km. Faults
of this orientation are common in the area. Mafic volcanism in
the immediate vicinity has K-Ar ages of 0.31 Ma (tholeiitic basalt, 48% SiOz: Turley and Nash 1980) and 6.1 Ma (basaltic
andesite, 57% SiOz: Turley and Nash, 1980) thus bracketing the
eruption of the rhyolite. A low-shield volcano composed princi-
Topaz Rhyolites
pally of basaltic andesite developed 1 Ma (Peterson and Nash
1980). It lies 12 km to the north and is centered on the fault trend
noted earlier. It appears that the tectonic-magmatic setting is
similar to other Great Basin rhyolites that were emplaced during
late Cenozoic east-west extension and during episodes of mafic
volcanism.
5. Keg Mountain, west-central Utah
Rhyolite flows and domes in the Keg Mountain area
(termed McDowell Mountains on some maps) have been
mapped as part of the Topaz Mountain Rhyolite (Erickson
1963). A brief description of them is included here, as Shawe
(1972: 72) implies they bear topaz in lithophysae. However,
Erikson (1963) and'Lindsey et al. (1975) report no topaz from
the Topaz Mountain Rhyolite at Keg Mountain. In our visits to
the Keg Mountains we have not yet identified any topaz-bearing
rhyolites.
Erickson (1963) identified atleast two vents and estimated
that about 15 km 3 of rhyolitic volcanic rocks are preserved in the
Keg Mountain area. The emplacement of the rhyolites was similar to other domes and flows with an upward sequence of lithicrich tuff, basal vitrophyre and/or flow breccia, and flow-banded
rhyolite. The rhyolite lavas attain maximum thicknesses of about
60 m near their vents. The underlying tuffs accumulated to similar thicknesses and thin toward the margins of the flows. The
lavas are 8 Ma (Lindsey et a!. 1975).
Most of the lava is felsitic and flow-banded with less than
10% phenocrysts of quartz, sanidine, plagioclase, and sparse biotite (Erickson 1963). Hematite and possibly topaz occur in lithophysae within the rhyolite.
The geochemistry and mineralogy of the tuff associated with
the rhyolites has been the subject of several investigations to
determine the role of alteration on the mobility of Be and V
(Lindsey 1975; Zielinski et a!' 1980), but the lavas remain poorly
studied. The tuffs contain abundant clasts derived from older
volcanic rocks (some as large as 5 m in diameter) that were
probably included in the ash as it explosively vented to the surface. Broken phenocrysts of quartz, sanidine, plagioclase, and
biotite constitute 5 to 15% of the tuff. Magnetite, hematite, titanite, hornblende, augite, zircon, fluorite(?), and apatite occur in
smaller amounts, but may be xenocrystic.·The average trace element composition of the tuff is given in Table 2. The relatively
high concentrations of Sr and Ba and low concentrations of Ga,
Nb, Y, Th, and V suggest that if the rhyolites are topaz-bearing
they are relatively poor in incompatible lithophile elements compared to others from the western Vnited States.
The so-called "alkali"rhyolites of the Keg Mountain area
are deposited on an older sequence of (apparently topaz-free)
rhyolitic lavas (10 Ma), ash-flow tuffs (32 to 30 Ma), and
intermediate-composition lavas and breccias (38 to 39 Ma)
(Shawe 1972; Lindsey et a!. 1975). The temporal magmatic and
tectonic development appears to be similar to that in the nearby
Thomas Range.
15
6. Mineral Mountains, western Utah
In the Mineral Mountains of western Vtah topaz-bearing
rhyolite domes occur discontinuously along the crest of the range
(Evans and Nash 1978; Lipman et a!. 1978b). The Mineral
Range is a typical basin and range horst that consists predominantly of a multiple-phase granitic pluton of Tertiary age (V-Pb
zircon age 25 Ma: Aleinikoff et a!. 1985). The Roosevelt Hot
Springs Known Ge~thermal Resource Area (KGRA) is located
along one of the western range-front faults.
Three sequences of rhyolite lavas exist in the Mineral Range
(Lipman et a!. I978b). The general distribution of the lavas is
shown in Figure 6. The oldest episode of volcanism (7.9 Ma; all
dates are by K-Ar method) is represented by a single dome (Trd)
on the western flank of the range. Quaternary' obsidian flows
(0.77 Ma) several kilometers long and up to 80 m thick were
erupted shortly before the most viscous topaz-bearing domes
(0.58 to 0.53 Ma). At least 11 domes were erupted. The domes
usually have a basal vitrophyre (5 to 10 m thick) that in places
forms part of a flow breccia. The vitrophyre grades upward into a
devitrified, flow-banded rhyolite. A frothy mantle of perlite forms
a carapace around parts of some domes. Lithophysae and other
gas cavities occur in flow interiors. In contrast to the obsidian
flows, several of the domes overlie pumice-rich pyroclastic deposits that represent the vent-opening eruptions. The domes· range
from 0.3 to I km in diameter and are up to 250 m high.
The rhyolite domes are crystal-poor with 2 to 10% phenocrysts of sanidine, quartz, oligoclase, and sparse biotite. Allanite,
titanite, zircon, Fe- Ti oxides, and apatite are accessories. Topaz
occurs in vugs along with pseudobrookite and specular hematite
as a result of crystallization from a vapor phase (Evans and Nash
1978).
Chemical analyses of rhyolites from the domes show that
they are typical of topaz rhyolites elsewhere (Tables I and 2). The
fluorine content of the vitrophyres from the domes ranges from
0.1 to 0.4%. Relative to the slightly older obsidian flows, the
domes are enriched in f, Na, Mn, Rb, Th, Nb, HREE, and Zn
and are depleted in K; Ca, Fe, Ti, Ba, Sr, Zr, light REE (La
through Tb), and Eu. Temperatures, derived from Fe-Ti oxides
(Evans and Nash 1978) and O-isotope geothermetry (Bowman et
a!. 1982) are consistently lower in the F-rich domes than in the
older obsidian flows (650 versus 760°C). The dome lavas also
crystallized at lower oxygen fugacities (log f0 2 ranges from -16.8
to -18.2). These and other observations led Evans and Nash
(1978) to suggest that the magma in the domes was a differentiate
of the obsidian-flow magmas. The REE pattern of the rhyolites
from the Mineral Mountains is different from most other topaz
rhyolites (Figure 40e). The principal difference lies in the relatively low concentrations of the middle REE (Nd through Tm);
in this regard they are similar to the patterns of the Lake City,
Colorado, topaz rhyolites. This depletion is probably the result of
titanite fractionation. A sample from one of the older lavas (the
flow of Bailey Ridge; Figure 6) was analyzed for its Sr-isotope
composition (Table 3). Its'ratio (0.7062 ± 0.0003) indicates that
16
'.
Qac
Tg
(I)
38°
30'
Z
<
I-
Z
:)
o
Tg
:E
I
I
I
I
-J
«
Qac
0::
w
z
:E
T~
38°
corral
*
\
o
2
I
I
Probable vent
3
I
KM
22'L-----......,..----'~----------------'---'-----......,..~
30"
Figure 6, Generalized geologic map ohhe central Mineral Mountains, Utah (after Lipman eta!' 1978a).
Topaz-b!laring rhyolite domes (Qrd) are underlain by rhyolitic tephra (Qrp). Topaz-free obsidian flows
(Qrl) generally lie directly on the granite of the Mineral Mountains (Tg). Surficial deposits (Qac) and
hot-spring deposits (Qhl) occur along a range-front fault (bold line). A Tertiary rhyolite (Trd) lies in the
extreme ,southwest comer of the map. Stars indicated vents for the rhyolite domes.
Topaz Rhyolites
17
contemporaneous fluorine-poor rhyolites by their low crystal
content (less than 5%) and the presence of smoky quartz. A few of
the older topaz rhyolites are phenocryst-rich (e.g., the rhyolite of
the Staats mine). The major phenocrysts in both age groups are
the same and include sanidine, quartz, sodic plagioclase, and
sparse Fe-rich biotite. A few samples contain Fe-rich hornblende.
Magmatic accessories include zircon, allanite, apatite, fluorite,
and Fe-Ti-Mn oxides. Two-feldspar geothermometry indicates
that some of the topaz rhyolites crystallized at temperatures
around 650°C (Table 4). Topaz, fluorite, alkali feldspars, hematite, bixbyite, and silica minerals line gas cavities and occur in the
devitrified matrix of some lavas. Red beryl (up to 1 cm in diameter) occurs in the eastern Wah Wah Mountains (Ream 1979).
Vapor-phase garnet occurs in some topaz rhyolites but it crystallized as a magmatic phase in the tuff of Pine Grove erupted from
a developing (molybdenum-mineralized) intrusive system in the
southern Wah Wah Mountains (Keith 1980, 1982). This tuff is
distinct in age (24 Ma) and chemistry from the younger topazbearing lavas (Keith 1980, 1982). Although the composition of
the rhyolitic portion of the tuff of Pine Grove lies within the range
of all topaz rhyolite compositions, it is enriched in Ba, Sr, Sc, and
Al and depleted in F, Zr, Rb, Nb, Ta, Th, U, Yb, Y, and Mo
relative to the topaz rhyolite lavas of SW Utah. The late "Lou
Rhyolite," which cuts the Pine Grove stock, is nonetheless geochemically identical to other topaz rhyolites from SW Utah
7. Wah Wah Mountains and vicinity, southwestern Utah
(Keith 1982). Its age is not known.
and southeastern Nevada
The chemistry of samples from southwestern Utah demonNumerous fluorine-rich rhyolitic lavas have been identified strates their similarity to other topaz rhyolites (Table 1). No
in the southern parts of three ranges- the Wah Wah Mountains, systematic differences in the compositions of the two age groups
the Needle Range, and the White Rock Mountains of southwest- of topaz rhyolites have yet been demonstrated. Fluorine concenern Utah and adjacent Nevada (Christiansen 1980; Best et al. trations in vitrophyres range from 0.2 to 05%. The trace element
1985). Interest in these lavas has been spurred by the discovery of chemistry is summarized in Table 2, and shows the typical ena porphyry molybdenum deposit in the southern Wah Wah richments of U, Th, Rb, Ta, Cs, and Li. REE patterns for topaz
Mountains (Tafuri and Abbott 1981; Keith 1980, 1982) and by rhyolites from this region are typical of others with low La/LuN,
the occurrence of numerous small deposits of uranium in and La/CeN, and Eu/Eu* (Figure 40t). Initial Sr-isotope ratios range
near the rhyolites (Christiansen 1980).
from 0.706 to 0.712 (Table 3).
The distribution of the rhyolites is shown in Figure 7. Two
Extensive alteration of rocks at the surface is associated with
episodes of rhyolitic magmatism have been delineated-one at 22 the fluorine-rich rhyolites in the Wah Wah Mountains. Large
to 18 Ma and a second at about 12 Ma (Rowley et al. 1978b; areas of jasperoid (Lindsey and Osmonson 1978) and alunite
Best et al. 1985). The rhyolites occur as isolated plugs without (K-Ar age 15 Ma; Best et al. 1985) occur in this region. Fluorite
significant pyroclastic deposits (e.g., Observation Knoll), as iso- and uranium were deposited in tuffs and in the intrusive margin
lated domes or flows with underlying pyroclastic breccias and of the topaz rhyolite at the Staats mine (Christiansen 1980; Lindtuffs (e.g., the rhyolites at the Staats Mine and at the Tetons) and sey and Osmonson 1978; Whelan 1965). Other uranium prosas groups of coalesced domes and flows with interlayered tephra pects associated with the rhyolites are shown in Figure 7. Gold
deposits (e.g., the Broken Ridge and Steamboat Mountain areas). and silver have been recovered from quartz-carbonate-fluorite
The tuffs generally consist of thin pyroclastic flow, surge and veins in the Stateline area of the southern White Rock Mounminor air-fall units. Explosion breccias near some vents contain tains. These deposits are temporally related to the older episode of
abundant lithic inclusions of the local country rock. Evidence of rhyolitic magmatism (Keith 1980). The previously mentioned
mixing between rhyolitic and mafic magmas before eruption is porphyry molydenum deposit at Pine Grove in the Wah Wah
found in tuffs from the southern Needle Range (Christiansen Mountains, which bears topaz in the alteration zones (Tafuri and
1980). Vitrophyres a few meters thick are present at the bases of Abbott 1981), is slightly older (24 versus 20 Ma) than the first
some flows. Felsitic, flow-banded lavas with abundant vapor- episode of topaz rhyolite magmatism (Keith 1980).
phase cavities in their upper portions are typical.
The tectonic and magmatic history of the region is outlined
The topaz rhyolites can usually be distinguished from nearly by Rowley et al. (1978b, 1979) and Best et al. (1980, 1985). The
it was derived from a crustal source with a relatively low Rb/Sr
ratio like those typical of granulite facies metamorphic rocks.
This interpretation is consistent with the O-isotopic composition
of the lavas and domes (8 18 0 = 6.3 to 6.9 %0; Bowman et al.
1982), which also suggests a crustal source of high metamorphic
grade. Both lines of isotopic evidence probably preclude the derivation of the rhyolites from metasedimentary rocks in the Proterozoic age basement. No metalliferous ore deposits are known to
be associated with the rhyolites of the Mineral Mountains.
The Mineral Mountains are part of a broad region of the
western United States that experienced late Cenozoic extension
and bimodal volcanism. Extensive fields of contemporaneous basalt and rhyolite (Figure 5) are located north and east of the range
(Haugh 1978; Clark 1977; Crecraft et al. 1981; Hoover 1974;
Evans and Steven 1982). Most ofthe fields appear to be less than
2.5 Ma. The near contemporaneity of basalt and rhyolite volcanism in the Mineral Range area and the inclusion of vesiculated
basalt "xenoliths" in the obsidian flows implies a close genetic
connection between the magmas (Evans and Nash 1978). The
rhyolites of the Mineral Mountains lie along the .PiocheMarysvale mineral belt of Shawe and Stewart (1976), which is
described more fully below. The distribution of volcanism may
be controlled by fractures that parallel this trend.
18
o.<8)
Ql
Cl
c
as
a:
Ql
+
'C
Ql
Ql
Z
Spar
Horse
Q~~Cottonwood
~/ .
Resevolr
Creek
•
~nto\springS
o
Thermo Hot
Springs
Uranium prospects
Rhyolite flows & intrusions
(Miocene & Pliocene)
~
L:::..:..:..:.J
Granite & quartz monzonite
(Miocene)
Granodiorite
(Oligocene)
o•
i
kilometers
25
Figure 7. Regional distribution of Miocene and Pliocene rhyolites (many of which contain topaz) and
Oligocene and Miocene intrusive rocks in southwestern Utah and eastern Nevada (after Christiansen et
al. 1980; and Best et al. 1985). The locations of uranium prospects and of the topaz-bearing molybdenite
deposit at Pine Grove are also shown. Crystal-poor rhyolite lavas in the western Needle Range, at Dead
Horse Reservoir, and at Thermo Hot Springs contain no topaz and are chemically distinguished from the
topaz rhyolites by their low concentrations of Rb, Y, and Nb, and by high concentrations of Sr. Basin
and range fault-block mountains are outlined with solid lines.
volcanic development of the region was dominated by events
along the east-trending Pioche mineral belt of which the Blue
Ribbon lineament is one element (Rowley et al. 1978b). This
feature coincides with a broad aeromagnetic high and·is marked
by numerous volcanic and intrusive centers of Oligocene to Holocene age. Calc-alkaline volcanism (andesite-dacite-rhyolite)
began in the early Oligocene (about 32 Ma) and continued until
the Miocene (until at least 19 Ma in the Black Mountains: Rowley 1978). This volcanism produced widely scattered, partly clustered, composite volcanoes with andesitic to dacitic lavas along
the belt. The intervening lowlands were covered by widespread
dacitic to rhyolitic ash flows erupted from collapse calderas. The
older F-rich rhyolite domes (20 to 18 Ma) are nearly contemporaneous with the latter portion of this episode but were accompanied by the eruption of trachyandesite (62 to 54% SiOz) lavas
in the Wah Wah Mountains, forming a bimodal suite there (Fig-
ure 42; Keith 1980; Best et al. 1985; Christiansen and Wilson
1982). High-K andesites erupted along the belt from 25 to 21 Ma
(Best et al. 1980, 1985). Fluorine-poor, crystal-rich rhyolite lavas
were also erupted during this time beginning about 22 Ma. A
genetic relationship between the F-rich and F-poor rhyolites has
not been established. A Miocene "lull" in volcanic activity in the
eastern Great Basin was followed by renewed bimodal magmatism that began about 13 Ma ago (Best et al. 1980). Mafic lavas
(trachybasalts to trachyandesites) were again accompanied by the
eruption of the younger topaZ-bearing rhyolites in the southern
Wah Wah-Needle Range area and by topaz-free rhyolite elsewhere (e.g., Dead Horse Reservoir, Thermo Hot Springs). In the
Wah Wah Mountains, mafic lavas and rhyolite were not erupted
from the same centers but their ages correspond closely (Figure
5). The two episodes of topaz rhyolite volcanism are reminiscent
of those to be described in New Mexico and Colorado. The
Topaz Rhyolites
19
Pioche belt is paralleled by NE-trending faults (some of which the northern fringe of the aeromagnetic high which marks the
may have strike-slip movement) and a conjugate NW-trending Pioche mineral belt (Fig. 4). 2) The mafic part of the bimodal
set. Best et al. (1985) suggest that the tectonism along the trend suite consists of more mafic lavas-trachybasalts with up to 2.4%
may be the result of relaxation of the Cretaceous compression K20 at 50% silica. Farther east, lavas as mafic as basalt do not
that produced regional thrust sheets. Alternatively the tectonic appear until much later (8 to 9 Ma; Fig. 5). 3) The topaz rhyolites
and magmatic events may be localized by a lithospheric fracture in the Wilson Creek Range are about 3 Ma older than the oldest
produced or reactivated by differing rates of extension north and topaz-bearing lavas in the Wah Wah Mountains. 4) The rhyolite
south of the Pioche belt (Rowley et aI. 1978b).
lavas may be cogenetic with earlier voluminous ash flows.
8. Wilson Creek Range, southeastern Nevada
9. Kane Springs Wash, southeastern Nevada
Barrott (1984) reports the presence oftopaz in rhyolite lavas
that are exposed in the northern part of the Wilson Creek Range
near Rosencrans Knolls. The lavas are part of a strongly bimodal
sequence of rhyolitic ash flows and basaltic lavas that were emplaced over a short period of time about 23 Ma. The topazbearing lavas (K-Ar age 22.6 ± 0.9 Ma; all ages are from Barrott,
1984) cap knolls underlain by thick sequences of unwelded lithic
tuffs. Interstratified with this lower tuff section is a prominent
rhyolitic welded tuff (K-Ar age 23.4 ± 0.9 Ma) and a series of
trachybasalt lavas (K-Ar age 23.9 ± 1.0 Ma).
The rhyolite lavas are geochemically similar to other topaz~
bearing lavas in their uniformly high Si02 contents (75 to 76
wt%), and their low concentrations of Ti02, Fe203, MgO, and
Cao (Table 1). No fluorine analyses have been reported A
slightly older (23.4 ± 0.9 Ma K-Ar) welded tuff also consists of
high-silica rhyolite that is only slightly enriched in Ti, Fe, Mg, and
Ca relative to the rhyolite lavas. The trace element differences
between the two rhyolites is more striking. The topaz-bearing
lavas are significantly enriched in Rb (1 to 3 times), Y (2 times),
and Nb (2 to 3 times) relative to the zoned(?) tuff. Likewise, the
lavas are depleted in the feldspar-compatible elements Ba (¥.!) and
Sr (lh). The trace element composition for five analyses (Barrott
1984) is summarized in Table 2. The genetic association of the
two rhyolite units has not been adequately examined but their
close association in space and time suggests that they may be
co-genetic.
The rhyolitic ash flows and early topaz-bearing lavas of the
northern Wilson Creek Range are part of a Miocene series of
potassic mafic lavas and rhyolite with no intermediate compositions known to be contemporaneous. This early Miocene volcanic field developed within the bounds of a huge Oligocene
caldera complex-the largest of these calderas was the Indian
Peak caldera which formed upon the eruption of several thousand cubic kilometers of dacitic magma (Best et aI. 1985). The
younger high silica tuff may have been erupted from a trap-door
caldera and has a much smaller volume (less than 20 km 3; Barrott 1984). The Miocene volcanic rocks were erupted about the
same time as the bimodal sequence of potassic mafic lavas and
rhyolites described from the southern end of the Wah Wah and
Needle Ranges. However, they are distinct from these lavas in
several important ways. 1) The rhyolites of the northern Wilson
Creek Range are not a part of the more southerly volcanic belt,
which includes topaz rhyolites of southwestern Utah. They lie on
An unusual association of topaz rhyolite lavas with the development of a mildly peralkaline caldera system has been reported by Novak (1984). The Kane Springs Wash caldera (Noble
1968) is located in southeastern Nevada and lies at the extreme
southwestern part of the ENE-trending aeromagnetic ridge of the
Delamar-Iron Springs mineral belt (Figure 4).
The following description of the evolution of the Kane
Springs Wash volcanic center is taken from Novak (1984). The
eruption oftrachytic lavas 14.2 Ma (all ages are by K-Ar method)
was the first recorded event at the volcanic center. A 19 by 13 km
caldera (Figure 8) collapsed slightly later (14.1 ± 0.2 Ma) as a
result of the eruption of several compositionally zoned (comenditic or fayalite rhyolite to trachyte) ash flows that had a total
volume of over 130 km 3. Immediately following collapse, the
eruption of trachyte lavas formed a cumulodome on the floor of
the caldera. High-silica rhyolites (with ferroedenite) and basaltic
to trachyandesitic lavas were then erupted, marking a fundamental change in the magma chemistry to non-peralkaline compositions. These units have K-Ar ages that are indistinguishable from
those of the other post-caldera volcanic rocks. The caldera experienced no structural resurgence. More trachyandesite lavas and
several biotite rhyolite domes are the last eruptive products of the
Kane Springs Wash center. Two of the biotite rhyolite domes lie
in the southwestern part of the caldera's moat and contain vaporphase topaz. Novak (1984) reports a K-Ar (sanidine) age of 13.2
± 0.2 Ma for one of these domes, about 900,000 years younger
than the comenditic ash-flow eruptions. The largest dome is
about 2 km across and is underlain by an initially erupted blanket
of tephra. Novak interprets the topaz rhyolites to be cogenetic
with the early trachytes and comendites and suggests that they
were derived from a residual pocket of magma within the solidifying sub-caldera chamber. Another topaz rhyolite dome occurs
10 km to the SE, where it plugs the vent of a post-caldera trachyte. Small patches of variably potassic olivine basalt (OJ to
2.3% K20) cap sections both inside and outside the caldera and
have K-Ar ages of 12.7 to 11.4 Ma.
Important differences in the mineralogy exist between the
early rhyolites and the younger biotite and topaz rhyolites. Novak
(1984) reports that the biotite rhyolites contain both sanidine
(Or6o) and plagioclase (Ab76) while the ferroedenite and earlier
comenditic rhyolites contain only sodic sanidine or anorthoclase.
In addition, the topaz rhyolites contain quartz, Fe-rich biotite,
ferroedenite, and fayalite, which is rarely reported in other topaz-
20
i
N
Kane Springs Wash
Caldera, Nevada
o
!
km
(
5
Figure 8. Generalized geologic map of the Kane Springs Wash caldera, southeastern Nevada (after
Novak 1984). Map symbols: Tb = basaltic lava flows; Trt = topaz rhyolite lava domes; Tr = post-caldera
rhyolite lavas and associated tephra; Ttr =interstratified post-caldera trachyandesite lavas and rhyolitic
ash-flows; Tts =post-caldera extrusive trachyte and syenite dome complex; Tkw =intracaldera exposure
of zoned (trachyte to rhyolite or comendite) Kane Wash tuff which is associated with collapse of the
caldera; extra-caldera rocks are stippled; * = vents fur rhyolite domes.
bearing rhyolites (Table 8). Accessory phases include magnetite, than Cl in silicate (especially hydrous) minerals (cf. Kovalenko et
ilmenite, and zircon. Earlier rhyolites either contained no hydrous al. 1984). Nonetheless, the eruption of comendite and attendant
phases (some of the ash flows) or ferroedenite without biotite.
loss of a Cl-rich vapor from the residual melt could leave a
To date, little geochemical information has been published non-erupted residue with a higher FICI ratio. Subsequent fracfor the volcanic rocks of Kane Springs Wash. Novak (1984) tionation of this magma might produce a metaluminous·rhyolite
reported one major-element analysis of the topaz rhyolite show- enriched in incompatible elements. Such a scenario was outlined
ing that it is indeed a typical metaluminous high-silica rhyolite by Christiansen et al. (1983a) for the generation of some meta(Table 1). Compared to the comenditic ash flows, the topaz luminous rhyolite magmas.
rhyolite is enriched in F, AI, Ca, and K, but it is depleted in CI,
The Kane Springs Wash center is one of a number of peralFe, and Ti. A vitrophyre from the topaz rhyolite dome contains kaline volcanic centers in the Great Basin of Nevada and adjacent
0.49% F compared to 0.34%F in a glassy comendite. More impor- parts of California and Oregon (Noble and Parker 1974). Howtantly, the F/CI ratio of the comendite (2) is much lower than ever, the volcanic field lies on an east-west belt of predominantly
that in the younger topaz rhyolite (10), even though both units metaluminous volcanic rocks ·of late Cenozoic age (Figure 4).
are enriched in fluorine and are presumed to be co-magmatic. A The segment of the belt in eastern Nevada is paralleled by the
less F-rich ferroedenite rhyolite has 0.17% F and an FICI ratio of Pahranagat shear system (Tschantz and Pampeyan 1970), which
slightly less than 3 as well. It is doubtful that the closed-system shows late Cenozoic strike-slip movement. Basin and range faultfractionation ofa parental comendite could produce an aluminous ing developed several million years after the Kane Springs Wash
melt or elevate F/CI ratios. Fluorine should be more compatible volcanism.
Topaz Rhyolites
21
Eastern Great Basin Chronologie Summary
Age
m.y.
o
Basalt
K-mafic
Rhyolite
B. Andesite
Calc-Alkaline
T
x
;;
x
.
x
;(
~. (
10
E~W.
><
n
\
i'"
Comments
t
Extension
1
-
NE-SW Extension
Miocene
20
(
Pt.
)1\
u
~Iull"
Colorado Plateau
uplift
Rhyolite
of Pine Grove
,!
30
.r"\
,......,
..
.c
:5
40
I,
I~ )
~
Figure 9. Schematic representation of magmatic and tectonic activity in the eastern Great Basin of
Nevada and Utah. There appear to be several separate eruption episodes for topaz rhyolites (x) in this
region: 1) contemporaneous with potassic mafic lavas and waning calc-alkaline magmatism in western
Utah and southeastern Nevada which peaked about 20 Ma. (This volcanism was accompanied by
NE--:SW faulting along the Pioche-Marysvale mineral belt and the formation of a topaz-bearing Climaxtype Mo deposit at Pine Grove.); and 2) a younger series of topaz rhyolites that are contemporaneous
with basalts and basaltic andesites in western Utah and southern Idaho. (Tectonism accompanying this
last episode was produced by E-W extension and normal faulting.)
Topaz rhyolites in the eastern Great Basin: A summary
Two groups of topaz rhyolites can be discerned in the eastern Great Basin of Nevada and Utah (Figure 9). The division is
based on the age and nature of the associated magmatism. An
early Miocene group, associated with K-rich mafic lavas or calcalkaline intermediate composition (dacites and rhyodacites) lavas
and tuffs, was erupted 23 to 18 Ma ago. This group is represented
by the older topaz rhyolites of the Wah Wah Mountains vicinity;
the lavas of the northern Wilson Creek Range, the rhyolite at
Spor Mountain, and the Be-rich Sheeprock granite (Christiansen
et al. 1983b). A late Miocene to Pleistocene group, associated
with potassic basaltic magmatism, was erupted after about 13
Ma. Intermediate composition lavas and tuffs are very small in
volume, but include the andesites of the area around the Mineral
Mountains, Utah, and the trachytes of Kane Springs Wash, Nevada. These magmatic episodes are separated by an apparent
Miocene "lull" in volcanic activity in the extreme eastern Great
Basin. Both groups of rhyolitic rocks are compositionally similar
and were erupted during apparent crustal extension. Most of the
rhyolites occur in EW-belts nearly parallel to the presumed direction of extension. The change in the dominant magma composi-
tions from intermediate/silicic to potassic basaltic lavas is similar
to that observed in Colorado but postdates it by about 6 to 10
Ma.
10. Cortez Mountains, north-central Nevada
Several plugs and flows of Miocene rhyolite occur in the
southern end of the Cortez Mountains in north-central Nevada
(Gilluly and Masursky 1965). Wells et al. (1971) were the first to
report that the rhyolites contained topaz in "vesicles." The rhyolites have been studied in several investigations concerned with
the age and origin of Au-Ag mineralization at the nearby Buckhorn and Cortez mines (Wells et al. 1971; Rye et al. 1974).
The rhyolites crop out over an area of about 3 km 2 and
occur in slightly dissected domes with pronounced flow banding
(Figure 10). Topaz rhyolites from the vicinity of Horse Canyon
have K-Ar ages of 15.3 Ma (Wells et al. 1971) and 14.5 Ma
(Armstrong 1970). They intrude or overlie slightly older flows of
basaltic andesite (16.3 Ma) and the Ordovician Vinini Formation
(siltstone, shale, and sandstone). Similar bimodal associations of
Miocene rhyolite and basaltic andesite are common in northcentral Nevada (e.g., Shoshone Range and Sheep Creek Range).
22
Cortez - Buckhorn, Nevada
i
N
0
, 1
, 2,
km
Tba
BUCKHORN
MINE
+
Tv
Figure 10. Generalized geologic map of the Buckhorn area in the southern Cortez Mountains, Nevada
(after Rye et aI. 1974). Miocene topaz-rhyolite lavas and domes (Tr) were ernpted in close proximity to
slightly older flows of basaltic andesite (Tba). Map symbols: Qts = undivided Quaternary and Tertiary
sedimentary rocks; Tr = Miocene rhyolites; Tv = mid-Tertiary volcanic rocks; Ii = Jurassic intrusive
rocks; Pu = undivided Paleozoic sedimentary rocks.
The rhyolites are generally phenocryst-poor (less than 3%) are similar in their Pb-isotopic composition to geochemically
felsites with a few lithic fragments. The phenocrysts are smoky distinct Oligocene volcanic rocks in the Cortez Mountains. Thus,
quartz, plagioclase and sanidine accompanied by sparse biotite. the radiogenic Pb isotope ratios may have arisen by contaminaMiarolitic cavities are common and are lined with topaz, fluorite, tion of both the rhyolitic magmas and the older Oligocene magmas by the same crustal reservoir. The Pb-isotope ratios may not
and silica minerals (Wells et aI. 1971).
No published information exists for the major element geo- be inherited from their sources.
The Miocene rhyolites of the Cortez Mountains appear to be
chemistry of the Miocene rhyolites. The average concentrations
of some trace elements are given in Table 2. Their Rb-rich nature spatially and temporally associated with Au, Ag, and Hg mineralindicates that they are similar to other topaz rhyolites. In addition, ization at Horse Canyon and at the Buckhorn mine, which sugWells et aI. (1971) state that they also contain high concentra- gested to Wells et aI. (1971) that they were the source of the
tions of Sn (20 ppm) and Be (10 ppm). Initial strontium-isotope ore-forming fluids. In contrast, on the basis of low D/H ratios in
ratios determined on two felsites with relatively high Sr concen- the alteration minerals, Rye et aI. (1974) demonstrated that metrations are similar to those from the Thomas Range (Table 3; teoric water was the principal ore-fluid. Meteoric waters may
0.7086 ± 0.0002; 0.7080± 0.0002 calculated at 14.9 Ma) and have been heated by the rhyolitic magmatism creating a small
are consistent with the involvement of a crustal component in geothermal system. The Au and Ag mineralization occurs in veins
their generation. Lead isotope ratios for two samples of the rhyo- controlled by NNW-trending faults.
The Miocene volcanism in this area is apparently related to
lite (Rye et aI. 1974) fall within the lower portion of the Area II
field delimited by Zartman (1974). However, they have lower the development of the Cortez (or Nevada) rift (Mabey et aI.
87Sr/86Sr and 207Pb/ 204Pb than many other felsic Area II rocks 1978; Stewart et aI. 1975). The rift coincides with a prominent
that show evidence in terms of high initial 87Sr/86Sr, aeromagnetic high that trends NNW across north-central Ne207Pb/ 204Pb, and 208Pb/ 204Pb ratios and low epsilon Nd (D. E. vada, and is marked by voluminous basaltic andesite and rhyolite
Lee, unpublished data; Stacey and Zartman 1978; Farmer and with minor basalt (Figure 11). Rifting postdates Oligocene erupDePaolo 1983) for a significant sedimentary or metasedimentary tions of calc-alkaline rhyodacite to rhyolitic ash flows. The rift
component in their sources. Nonetheless, these Nevada rhyolites merges to the north with the Snake River Plain-Columbia River
have distinctly higher uranogenic Pb isotope ratios than the topaz volcanic fields (Figure 12). Mabey et aI. (1978) suggest that the
rhyolites from Lake City, Colorado, which suggests that the Cortez rift was produced by the initial Cenozoic extension in this
sources of the Cortez rhyolites are older or had higher U/Pb part of the Basin and Range province. It is parallel to anp conratios. However, it is important to note that the topaz rhyolites temporaneous with several other northwest-trending extensional
Topaz Rhyolites
23
-----,--,---I
I
l
I
I
'.
I
[
,
\
'" ,
,-I
I
~
Daa
~
Tr
Imm~~~
Tba
lI
u
o
I
KM
I
o
I
50
km
Figure 11. Generalized geologic map of north-central Nevada showing
the distribution of 13.8 to 16.3 Ma lava flows and rhyolite domes (after
Stewart et al. 1975). The volcanic rocks are concentrated along the
Cortez rift that opened 15 to 16 Ma. The locations of topaz rhyolites in
the Sheep Creek and Cortez Mountains are shown. Also shown are the
traces of the Golconda (GT) and Roberts Mountain (RMT) thrusts
(Stewart 1980) that may mark the approximate site of the rifted margin
of the Precambrian continent. Map symbols: Qa - Quaternary alluvium;
Tr - Miocene rhyolite domes and flows; Tba - Miocene basaltic andesite,
locally including andesite, basalt, and dacite of uncertain age; U - undivided Tertiary, Mesozoic and Paleozoic rocks. Bold lines are normal
faults.
features in the northwestern United States, including the western
Snake River Plain graben and the vents for the Columbia River
basalts (Figure 12). The continuity of the rift is interrupted by
younger northeast-trending faults produced after reorientation of
the regional stress-field about 10 Ma (Zoback and Thompson
1978). The Cortez rift appears to have formed near the presumed
eastern boundary of the Precambrian crystalline basement (Figure 1), which also appears to limit the distribution of topaz rhyolites (Christiansen et al. 1983a).
11. Sheep Creek Range, north-central Nevada
Farther north along the Cortez rift, another topaz rhyolite
Figure 12. Map showing the distribution of Miocene (20 to 10 Ma)
volcanic rocks and tectonic features of the northwestern United States
(after Davis 1980). The relationship of the Cortez rift (C-R) to the
western limit of the Precambrian continental crust (heavy dashed line) is
shown along with other NW-trending tectonic features: DS = feeder
dikes for the CRE = Columbia River Basalt; WSR = western· Snake
River Plain graben; M = Monument dike swarm; OV = Orevada rift
(Rytuba and McKee 1984). The subduction-related volcanic arc (CR =
Cascade Range) active at the time was located near the continental
margin. The topaz rhyolites of northern Nevada were erupted during this
period of NE-SW extension and basaltic volcanism and are contemporaneous with peralkaline rhyolites in NW Nevada, NE California, and
SE Oregon.
has been identified in the southern Sheep Creek Range near the
old Izenhood Ranch, which lies north of Battle Mountain (Fries
1942). The rhyolites are the hosts for cassiterite/wood tin mineralization similar to that found in the Taylor Creek Rhyolite in
New Mexico.
Miocene rhyolites (ca. 14 Ma; Stewart et al. 1977a) extend
nearly continuously from the Nevada-Oregon border southsoutheast for 150 km in a zone about 40 km wide along the
Cortez rift (Figure 11). The rhyolites were apparently emplaced
as coalesced domes and lava flows like those in the Thomas
Range, Utah. Individual domes and flows cover from 3 to 100
km 2 (Stewart et al. 1977a). In the Sheep Creek Range, the rhyolite intrudes a thick section of 14.8 Ma basaltic andesite (59%
Si02) and is overlain by younger (10 Ma) lavas of olivine basalt
(48% Si02) (Stewart et al. 1977a). The rhyolites are part of a
broadly contemporaneous (14 to 16 Ma) series of basalts and
basaltic andesites that occur along the Cortez rift.
The topaz-bearing phase of the rhyolite, near Izenhood
Ranch, is a crystal-rich (25 to 35%), flow-banded felsite. It is not
known if topaz is restricted to a single dome or a group of domes.
24
Northern Great Basin Chronologie Summary
Age
m.y.
o
Basaltic or
B. Andesite
Calc-Alkaline
Rhyolite
x 0
Comments
T
Development of
Snake River Plain
10
-
I I
~~
x
~g
0
\J
20
30
\
I
E- W extension
1
Cortez Rift opens
(NE-SW extension)
Calc-alkaline
Volcanism
N. Nevada/UT.
40
Compression ends
Figure 13. Schematic representation of magmatic and tectonic activity in the northern Great Basin. x =
topaz rhyolite age; 0 = peralkaline rhyolite age (Noble and Parker 1974; Rytuba and Conrad 1984). The
eruption of the rhyolites followed the decline of calc-alkaline intermediate to silicic volcanism of the
Oligocene. The topaz rhyolites appear to be closely associated with the opening of the Cortez rift, the
development of the Snake River Plain, and the eruption of basalt and basaltic andesite. Peralkaline
rhyolites were erupted contemporaneously, but they are chemically and for the most part spatially
distinct. Compiled from sources cited in the text and from Stewart and Carlson (1976).
The principal phenocrysts are quartz, oligoclase, and sanidine incrustations in fissures produced by the cooling of the lava.
with biotite, zircon, titanite, apatite, and Fe-Ti oxides as accesso- Meteoric waters mobilized by the hot rhyolites were important
ries. The vapor-phase mineralogy as described by Fries (1942) ore fluids, but magmatic fluids released during devitrification
consists of topaz, pseudobrookite, sanidine, silica minerals, fluor- probably produced the initial mobilization of Sn. Boiling would
ite, garnet, (described as andradite but it is spessartine- be expected at such a shallow level.
almandine), and possibly cassiterite. No pyroclastic deposits are
The magmatic and tectonic setting of the rhyolitic volcanism
exposed near the Izenhood Ranch locality and they are rare in the is identical to that in the Cortez Mountains. Lithospheric extensouthern Sheep Creek Range (Stewart et al. 1977a).
sion (NE-SW) along the Cortez rift was accompanied by the
The felsitic phase of the lava at Izenhood Ranch is composi- eruption of rhyolite and basaltic andesite. Along the rift zone
tionally similar to other topaz rhyolites (Table 1 and 2). How- basalt is minor in Nevada (but more voluminous in Oregon).
ever, it is higher in Fe, Zr, Y, Ga, and Sr than most and toward Judging from the age of the volcanic rocks, the rift formed 15 to
the low end ofJhe observed range for Rb (compare Figures 35 16 Ma in a back-arc environment (cf. Snyder et aI1976). Peraland 38). Our preliminary studies of the chemical composition of kaline rhyolites (Noble and Parker 1974), generally erupted from
other rhyolite lavas in the area suggest that they have similar calderas, are also typical of this period (less than 20 Ma) in the
compositions. The devitrified lava is fairly rich in uranium (12 Great Basin of Nevada (Figure 13). For example, 16 Ma comenppm; Christiansen et al. 1980); presumably the magma contained ditic ash flows were erupted from the McDermitt caldera comeven more, because uranium is generally lost during devitrifica- plex (Rytuba and McKee 1984), which developed on the western
tion of fluorine-rich lavas (Bikun 1980; Christiansen 1980). The flank of the northern Cortez rift.
initial 87Sr/86Sr ratio of one sample is 0.7085 (Table 3).
12. Jarbidge, northern Nevada
Fries (1942) describes wood-tin and cassiterite in veinlets
from the lava. The similarity of the mineral assemblage associated
The Jarbidge Rhyolite is a widespread volcanic unit in
with the tin mineralization to that developed in miarolitic cavities northern Nevada along the southern margin of the Snake River
suggests that the ores originated as fumarolic or pneumatolytic Plain (Coats 1964). It correlates with similar rhyolitic lavas as far
Topaz Rhyolites
,';
east as the Utah line and as far south as the East Humboldt Range
(Coats et al. 1977). Topaz has been identified in a single thin
section of the rhyolite (Coats 1964).
The rhyolite was emplaced as a series of volcanic domes and
flows that are underlain locally by flow-breccias and pyroclastic
units. The country rocks through which it was erupted are slightly
older volcanic rocks. Coats et al. (1977) report K-Ar ages for two
samples of Jarbidge Rhyolite as 16.8 and 15.4 Ma. The Jarbidge
Rhyolite is overlain by the voluminous Cougar Point Tuff (12.2
Ma; Coats and Stephens 1968), which is related to the development of the Snake River Plain.
The lavas contain phenocrysts of quartz, sanidine, and
oligoclase-andesine. The presence of small amounts of pigeonitic
clinopyroxene and hornblende as well as the absence of biotite
distinguish the Jarbidge Rhyolite from most other topaz rhyolites
in the western United States. Accessory phases include garnet
(apparently magmatic), zircon, apatite, and Fe-Ti oxides (Coats
1964).
The average of three analyses of the Jarbidge Rhyolite are
presented in Tables 1 and 2. These rhyolites are similar to those
from the Sheep Creek Range in their high concentrations of Fe,
K, and Zr relative to other topaz rhyolites. Nonetheless, these
features are typical of bimodal rhyolites of the Snake River Plain
and northern Great Basin (Wilson et al. 1983). The Jarbidge
Rhyolite is also lower in Rb (275 ppm) than most topaz rhyolites.
Analyses presented by Coats et al. (1977) suggest that F (less than
500 ppm) and Be (less than 5 ppm) are low in these lavas. The
scarcity of topaz is indicative of its lack of enrichment in incompatible lithophile elements. Initial 87Sr/86Sr ratios for three samples of the rhyolite collected by R. T. Wilson range from 0.7101
to 0.7142 (Table 3); ratIos that are typical of Cenozoic rhyolites
from north central Nevada and the Snake River Plain (Wilson et
al. 1983; Leeman 1982a).
The Jarbidge Rhyolite is the host for Au-Ag mineralization
that yielded about $10 million in these commodities before 1942
(Granger et al. 1957). The mineralization is found in epithermal
quartz-adularia veins.
The Jarbidge Rhyolite is part of a bimodal assemblage of
silicic (dacitic to rhyolitic) and basaltic volcanic rocks (Coats
1964) that was erupted during the formation of the western
Snake River Plain graben (Armstrong et al. 1975; Leeman
1982a). The lavas are approximately the same age as the basalts
and rhyolites in southwestern Idaho and the Columbia River
Basalt ofIdaho and Oregon (Figure 12). The Cortez rift, 100 km
to the west, was developing at about the same time. These rocks
may be magmatic products of the initial NE-SW extension of the
northern Basin and Range province (cf. Zoback et al. 1981).
Later faults that cut the Jarbidge Rhyolite (and other correlative
volcanic units) are north-trending and block out the present
mountain ranges of northern Nevada (Figure 13).
13. Blackfoot Lava Field, southeastern Idaho
A group of five small rhyolite domes occurs near the south
end of the Blackfoot Reservoir, about 15 km north of Soda
25
Springs, Idaho. D. R. Shawe (oral communication, 1982) and
Dayvault et al. (1984) report that topaz occurs as a devitrification
product in at least one of the domes, China Cap (called Middle
Cone on some maps). The rhyolites erupted during the development of the predominantly basaltic Blackfoot lava field. Armstrong et al. (1975) report K-Ar ages of 0.04 ± 0.02 Ma (whole
rock), 0.08 ± 0.04 Ma (sanidine), and -0.1 ± 0.1 Ma (whole
rock) for three specimens collected from these domes. Leeman
and Gettings (1977) report concordant K-Ar, hydration rind, and
thermoluminescence ages of 50,000 years, and G. B. Dalrymple
(cited in Pierce et al. 1982) reports a 61,000 ± 6,000 year age for
sanidine from the rhyolite at China Hat. These ages make this
rhyolite the youngest known topaz bearing rhyolite. Three other
small rhyolite domes occur about 25 km north of China Hat in
the similar Willow Creek lava field. S. H. Evans (written communication, 1980) reports K-Ar (sanidine) ages of 1.56 ± 0.06
and 1.28 ± 0.15 Ma for two specimens from these rhyolites.
The Blackfoot lava field lies in a NW-SE-trending Tertiary
graben flanked by mountain ranges that expose Paleozoic and
Mesozoic sedimentary rocks. Mansfield (1927) presents the most
complete geologic description of the region. Three rhyolite domes
(China Hat, China Cap, and North Cone) lie on a NE-trending
line transverse to the graben structure. Vents for basaltic lavas are
also aligned along this trend. China Hat is the largest of the domes
and has a maximum dimension of 2 km parallel to the lineament;
it rises to a height of almost 300 m above the basaltic lavas. Two
smaller bodies of rhyolite occur as islands within the reservoir,
one of which has a K-Ar age of 1.4 Ma. (S. H. Evans, cited in
Feisinger et al. 1982). The domes are older than the basalts
exposed at the surface, but inclusions of still older basalt and
andesite are found in the rhyolite lava (Mansfield 1927) and in
tephra (air fall and base surge deposits) beneath China Hat.
Therefore, the rhyolites formed in the midst of the mafic magmatism. Mabey and Oriel (1970) identified small positive aeromagnetic anomalies associated with the rhyolite domes and a
compound gravity low over the portion of the lava field where
the rhyolites occur. Based on an interpretation of the magnetic
high associated with the basalts, they suggest that the basalts are
about 1000 m thick. From the gravity expression of the field,
Mabey and Oriel (1970) proposed that a caldera centered on the
rhyolite domes collapsed after extrusion of the basaltic lavas, or,
alternatively, the anomalies were caused by a buried granitic
intrusion. An interpretation more consistent with the tectonic
setting (typified by late Cenozoic basin and range faulting) and
volcanological style (rhyolite domes with many small low shield
volcanoes and cinder cones in a plains-style basalt field) is that
the lava field formed over a small sediment-filled graben. The
low-density basin fill could produce the negative gravity anomalies without invoking caldera collapse, which is rarely associated
with the eruptions of small, isolated rhyolite domes. Mabey and
Oriel (1970) interpret the structures of nearby Gem Valley injust
such a manner. There, a slightly elongate gravity low is centered
on the principal basaltic shield volcano of the Gem Valley lava
field.
26
The rhyolites of the Blackfoot lava field are petrographically
similar to one another. The domes consist of phenocryst-poor,
generally devitrified, flow-banded lava. Groundmass textures
range from glassy to granophyric. Small phenocrysts of resorbed
quartz, oligoclase, biotite, and Fe-Ti oxides are ubiquitous. Apatite and zircon are common accessory minerals. Dayvault et al.
(1983) identified euhedra ofthorite in glassy specimens as well as
allanite, epidote, oxyhornblende, and unidentified grains of a CeTh phosphate (monazite?) and a Nb-Y-Ti-Fe oxide. Tridymite,
quartz, hematite, and topaz occur in lithophysae from China Cap.
The average major and trace element composition of six
specimens analyzed by Dayvault et al. (1983) are presented in
Tables 1 and 2. The major element composition ofthese rhyolites
is indistinguishable from other topaz rhyolites with high Si and
alkalies and low Ti, Fe, Mg, and Ca. Although none of the
specimens analyzed were obsidians, most contain glass in their
grounclmass. As a result, fluorine concentrations are quite high
(3500 to 5800 ppm) and may be close to magmatic concentrations. Chlorine concentrations are uniformly low (130 to 580
ppm) yielding characteristically high F ICI ratios (>9). The halogen ratios may have been disturbed by devitrification with the
preferential loss of Cl. Trace element concentrations in the rhyolites of the Blackfoot lava field are also typical of other topaz
rhyolites with elevated concentrations of Be (> 10 ppm), Li (>80
ppm), Sn (> 10 ppm), Y (> 150 ppm), U (> 15 ppm), and Th (27
to 60 ppm). Feldspar compatible elements such as Ba and Sr are
strongly depleted.
No known mineralization is related to the topaz rhyolites of
the Blackfoot lava field.
The Blackfoot lava field is centered about 60 km SE of the
margin of the eastern Snake River plain in the northern Basin and
Range province. Armstrong et al. (1975) outline the temporal
development of late Cenozoic volcanism in southern Idaho. They
point out that a time-transgressive series of rhyolitic ash-flow tuffs
were emplaced across southern Idaho. Volcanism began in
southwestern Idaho approximately 15 Ma and migrated to the
northeast to its present culmination in the Yellowstone area in
Wyoming. Basaltic volcanism was initiated in the wake of the
caldera-forming eruptions and has continued intermittently, creating the Snake River plain. Some of these plains-style (Greeley
1982) basaltic lavas were erupted from long fissures, but others
form low-shield volcanoes capped by small craters less than 1 km
across. The Blackfoot lava field is similar to the Snake River plain
in several important ways, including its age and the style of
basaltic volcanism, but the development of rhyolite domes late in
the history of a basaltic field is atypical of the Snake River Plain,
although several such domes formed in the central part of the
eastern plain. Likewise, no early ash-flow volcanism can be tied
to the Blackfoot field. Perhaps most importantly, the rhyolites of
the Snake River plain proper are generally pyroxene andlor
fayalite rhyolites, with moderately high Fe and Zr contents and
high KINa ratios that show high equilibration temperatures
(Leeman 1982a; Hildreth 1981; Hildreth et al. 1984; Wilson et al.
1983); none of these characteristics are found in the Blackfoot
lava field rhyolites nor in topaz rhyolites in general (with the
possible exception of the Jarbidge Rhyolite, see above). Smith
and Christiansen (1980) relate the formation of the Snake River
plain to the passage of the North American plate over a mantle
plume, while Christiansen and McKee (1978) relate its development to transform-style accommodation of E-W extensional
faulting proceeding at different rates to the north and south of the
plain. Because of the location of the Blackfoot lava field on the
floor of a fault-bounded graben, an extension of the Basin and
Range province to the south, we prefer to place the development
of the Blackfoot field into a context of basin-and-range extension,
bllSalt intrusion, and bimodal basalt-rhyolite volcanism.
14. Elkhorn Mountains, western Montana
A group of Oligocene topaz rhyolites has been identified in
and near the northern Elkhorn Mountains of western Montana.
Chadwick (1978) has included them in the Helena volcanic field
(Figure 14). The Elkhorn Mountains are bounded on all sides by
major faults and were probably uplifted during the mid-Cenozoic
(Smedes 1966).
The rhyolites in the Elkhorn Mountains were emplaced as
isolated intrusive plugs, dikes, and small domes and lava flows.
Some of the extrusive rhyolites overlie cogenetic pyroclastic deposits. Inclusions of the Butte Quartz Monzonite, the dominant
phase of the Cretaceous Boulder batholith, occur in the pyroclastic deposits erupted from centers within the mountains (Smedes
1966). Most of the rhyolite masses are less than 1 km across; four
were mapped by Smedes (1966). Chadwick (1978) obtained a
K-Ar age of 35.8 Ma on a sample from the topaz-bearing rhyolite
at Lava Mountain. Two other similar rhyolite flows from the area
have been dated and have K-Ar ages of 37.3 and 36.9 Ma
(Chadwick 1978; Figure 14).
Most of the rhyolite is felsitic, flow-banded, and phenocrystpoor. Obsidian is pres~rved along the margins of some dikes.
Smoky quartz, sanidine, and sodic plagioclase{An 10 to An12) are
the principal phenocrysts. The grounclmass is usually spherulitically devitrified or altered by vapor-phase crystallization. Topaz
(up to 30 mm long), along with quartz and fluorite, line miarolitic
cavities and lithophysae (Smedes 1966).
Smedes (1966) reports whole-rock chemical analyses of
several phases of the rhyolite from the Helena volcanic field that
demonstrate their chemical similarity to other topaz rhyolites in
the western United States (Table 1). Two samples of "soda rhyolite" from the Elkhorn Mountains, which may be from the topazbearing units, contain an average of 8.9 ppm Uand 37.6 ppm Th
(Tilling and Gottfried 1969). Greenwood et al. (1978) report that
the topaz rhyolite from Lava Mountain contains high amounts of
F (0.2 to 0.5%), Sn (10 to 50 ppm), Be (1 to 20 ppm), Nb (30 to
100 ppm) and moderate amounts ofMo (5 to 10 ppm).
Silver-bearing galena, sphalerite, fluorite, and quartz occur
in veinlets cementing a brecciated phase of the rhyolites at Lava
Mountain (Smedes 1966). However, most of the Ag, Au, and Pb
mineralization in the rocks of the Boulder batholith appears to be
Topaz Rhyolites
~CRATER MTN.
Late , Cenozoic Volcanism
In S W Montana
~
MONT
0
7.8U
-
29
ANA
0
37 .3
cf36.9 Helena
~~
_
1'\ '"
~"ON l.:.,:)
r\€.\..\)
FIELD
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3'-4~~"
~~~~
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(9
,
37
I77.l Basalt
VOLCANO BUTTE
BASALT ~
_
29.1
~,:f(; X:vQ
-:
27
ICLa
0
LiNEAA:1VtENT
_
Lava
Mountain
_
_
Rhyollite
Sample dated in m.a.
o
Rhyollite
Basalt
J;;l Trachyandesite
•
X Butte
o
,
,
!
5,0
KM
X Bozeman
32.7 34.4
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30.3 ",,\~G\,/
•
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38.9,>: Dillon
Q
(j>
URNS
.-,
HEPBSA
8.4 :~ALTS
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1.9
t::l ~_ _ If:l.~o,:..:.?~MT
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UPPER ON I
Wy
22.9. '0,"
tA~~iC\)[£)1
::' ',::,
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:. ',,", : : : . '
,I. . '
Y ELL OW S TON E
I
_/ SRP
I
.
Figure 14. Distribution of post 40 Ma volcanic rocks in southwestern Montana (after Chadwick 1978,
1981). The topaz rhyolite lavas and domes of Lava Mountain in the Elkhorn Mountains and other
rhyolites in the Helena and Beaverhead volcanic fields are about 10 Ma older than dated basalts from
the rest of the region but may be part of a bimodal (basalt-rhyolite) suite associated with the inception of
block faulting in the region (Chadwick 1981), Younger volcanic fields include those at Yellowstone and
the SW-trending Snake River Plain (SRP).
older than the rhyolitic magmatism (Smedes 1966). The Bald
Butte molybdenum prospect (Rostad 1978) is located about 25
km north of the topaz rhyolite locality, but it is older (about 48
Ma) and is more likely related to magmatism of the type described below from the Little Belt Mountains.
The rhyolites of the Helena volcanic field are part of a
broadly bimodal assemblage of basaltic and silicic lavas that were
erupted in southwestern Montana after about 40 Ma (Chadwick
1978, 1981). Individual fields do not generally contain both
magma types, but basalt flows of uncertain age (Pliocene to Oligocene) do occur in the Helena field (Greenwood et al. 1978).
This volcanic episode postdates dominantly calc-alkaline intermediate to silicic magmatism 45 to 55 Ma (Armstrong 1978) and
may mark the initiation of extensional tectonics and basin subsidence in western Montana (Pardee 1950; Chadwick 1978).
Chadwick (1978) suggests that the Helena and other volcanic
fields lie along the Montana lineament (or Lewis and Clark line)
that marks the northern boundary of a Precambrian basement
province (Weidman 1965), This lineament coincides with the
northernmost boundary of basin and range faulting and may have
acted as a transform allowing lithospheric extension and blockfaulting to proceed south of it (Reynolds 1977). The tectonic and
magmatic history of the region is summarized in Figure 15.
15. Little Belt Mountains, central Montana
The oldest Cenozoic topaz rhyolites we know of (about 50
Ma) are exposed in the Little Belt Mountains of central Montana
(Weed 1900; Pirsson 1900; Witkind 1973). The mountainous
terrain exposes several major alkaline and calc-alkaline plutons
that may be genetically related to the volcanic cover (Witkind
1973).
The topaz rhyolites and their less fluorine-enriched counterparts occur as sills and cylindrical plugs (Figure 16). Topaz has
been descnbed from a sheeted-sill complex that intrudes Cambrian sedimentary rocks near Yogo Peak (Pirsson 1900) and
within the rhyolitic "bysmalith" (a forcefully emplaced plug that
has pushed the overlying strata up along one or more circumfer-
28
Montana Chronologie Summary
Age
m.y.
Basalt
Calc-Alkaline
Rhyolite
20
Comments
Beginning of
Miocene "lull"
A
f\
30
1I
I I
I I
x
40
50
Block-faulting
begins (?)
(
\
60
\]
Mo-mineralized
plutons
Last Regional thrust
sheets
Figure 15. Schematic representation of Cenozoic magmatic and tectonic activity in western Montana.
Topaz rhyolites (x) were erupted during basalt-rhyolite volcanism and during older high-K calc-alkaline
and alkaline magmatism of central Montana. Compiled from sources cited in the text.
ential faults) at Granite Mountain (Pirsson 1900; Witkind 1973).
The rhyolite at Granite Mountain has a K-Ar age of 48.8 ± 2 Ma.
The rhyolite at Granite Mountain is described as being dense
and fine-grained with a groundmass of quartz, alkali feldspars,
biotite, and Fe-Ti oxides. Witkind (1973) classified topaz, along
with albite, as phenocrystic. However, in view of the apparently
devitrified nature of the rocks and its typical development elsewhere, the topaz is probably the result of growth from a vapor
phase. The rhyolite sills at Yogo Peak are also fine-grained but
contain granules of tourmaline in addition to topaz in the
groundmass (Pirsson 1900)~both are probably the result of the
exsolution of vapor.
Pirsson (1900) and Witkind (1973) report chemical analyses of the topaz rhyolites (Table 1). The sill analyzed by Pirsson
may have been slightly altered; it contains slightly more AI, Mg,
Ca, and P than most topaz rhyolites. The rhyolite of Granite
Mountain is typical of its class but has slightly lower Ti and
higher Mg than other topaz rhyolites. Rubidium is enriched
(about 525 ppm) and Sr depleted (about 9 ppm) in this phase as
well (Witkind 1973), which is consistent with other occurrences.
Rupp (1980) reports that a specimen from Granite Mountain
contains 3400 ppm F and is likewise enriched in Rb, U, and Nb
but not Sn (Table 2). As for most other topaz rhyolites Ba and Sr
are extremely depleted. Initial Sr-isotope ratios of two samples
from Granite Mountain average 0.7093 (t = 48.8 Ma) (Marvin et
al. 1973).
No mineralization is directly associated with the rhyolites,
but an intermediate-composition stock of about the same age is
related to Ag-Pb mineralization. Minor quantities of molybdenite
and scheelite occur in fissure veins within the stock (Witkind
1973). The Climax-type(?) Big Ben molybdenite deposit (Olmore
1979) is located less than 20 km to the west and has an age of
49.5 Ma (Marvin et al. 1973) making it approximately the same
age as the topaz-bearing Granite Mountain rhyolite, as well as the
Bald Butte Mo-prospect in the Elkhorn Mountains.
The topaz rhyolites in the Little Belt Mountains are part of a
complex assemblage of plutonic and volcanic rocks that were ail
emplaced between 54 to 48 Ma (Marvin et aI. 1973). The plutons
and laccoliths are composed of felsic (70 to 72% SiOz) to intermediate (61 to 67% SiOz) composition rocks. Syenite, shonkinite,
and lamprophyre (in places mixed with rhyolite in composite
dikes) are also exposed in small bodies (Witkind 1973). In Montana and Wyoming, Eocene magmatism was common 54 to 45
Ma (Armstrong 1978; Figure 15). Calc-alkaline and alkaline intermediate to silicic volcanic rocks and subjacent plutons are the
principal expressions of this activity. Lipman (1981) and Snyder
et al. (1976) relate the magmatism to subduction of oceanic
lithosphere near the continental margin. However, Armstrong
(1978) suggests that intra-arc rifting and basin sedimentation may
have been contemporaneous with this episode, which is correlative with the Challis volcanism of Idaho and arc-type volcanism
and graben formation in northern Washington (Davis 1980).
Monger and Price (1979) and Ewing (1980) also suggest that the
subduction-related volcano-plutonic arc formed nearer the con-
29
Topaz Rhyolites
o
I
I
I
I
I
5I
km
MPu
Little Belt Mtn., MT
EXPLANATION
TERTIARY INTRUSIVE ROCKS
U:::';r;}
-' -
~)-;1
/.
Rhyolitic/granitic
Intrusions
Q lame and
syenite intrusions
PRE-CENOZOIC ROCKS
Mesozoic-Paleozoic
sedimentary rocks
pC sedimentary rocks
pC igneous and
metamorphic rocks
Figure 16. Generalized geologic map of the Little Belt Mountains, Montana (after Marvin et al. 1973),
showing the locations of topaz rhyolites (x) and granites relative to alkaline intrusive rocks of the same
age. The location of the Big Ben molybdenite prospect is also shown.
tinental margin on a strike-slip faulted terrane. Thus, the Eocene
volcanism in Montana may have been "back-arc" in nature, in
keeping with the extensional faulting and the apparent tectonic
setting of other topaz rhyolites.
16. Specimen Mountain, north-central Colorado
I
I
"
Wahlstrom (1941) first reported topaz from gas cavities in
rhyolite lava flows from Specimen Mountain on the crest of the
Front Range in Rocky Mountain National Park. From Wahlstrom's (1941, 1944) descriptions it appears to be similar in its
mineralogy, chemistry, and emplacement history to other topaz
rhyolites.
The rhyolite lavas of Specimen Mountain (Figure 17)· are
underlain by a thick (lOOs of meters) sequence of pyroclastic
deposits (apparently composed of fall, flow, and breccia units).
These in tum overlie a basal complex of trachyandesite (quartz
latite) lavas and pyroclastic deposits. Wahlstrom (1944) describes
the core of Specimen Mountain as an intrusive rhyolite plug that
formed shortly after the emplacement of the upper rhyolite lavas.
All of the volcanic rocks overlie or intrude Precambrian gneiss,
granite, and pegmatite. Fragments of these rock types also occur
as inclusions in the vent agglomerates. A series of arcuate faults
cut the plug and flows-apparently formed as the rhyolite plug
collapsed back down its conduit shortly after emplacement (see,
for example, Corbett 1966, 1968). These relationships suggest
that the rhyolite flows and plug are part of a small dome complex.
The entire complex contains about 1.5 km 3 of material.
The topaz-bearing lavas generally show conspicuous flow
banding and are locally rich in lithophysae. Quartz and sanidine
are the only reported phenocrysts (Wahlstrom 1944). The rhyolite plug contains biotite, oligoclase, magnetite, and rare resorbed
hornblende in addition to quartz and sanidine. No topaz is reported from this phase of the complex. Wahlstrom (1944) includes the rhyolites and trachyandesites, with phenocrystic
plagioclase, amphibole, biotite, and augite, as part of the same
30
105· 50'
Specimen Mountain. CO
predominant volcanic rocks in the region are poorly studied rhyolitic ash-flow tuffs. Their temporal and chemical relationships to
the topaz rhyolites at Specimen Mountain are unknown. Likewise, the age of an older series of basalts(?) and trachyandesites is
not established (Eocene to Oligocene) but they represent only a
small volume of the Cenozoic effusive rocks (Corbett 1968). The
Never Summer stock, a zoned granodiorite to quartz monzonite
intrusion that outcrops about 5 km west of Specimen Mountain,
is also the same age (± 1 Ma) as the silicic volcanic rocks (Corbett
1968).
17. Chalk Mountain, central Colorado
o
,
I
1,
Km
40.25' L-._......L
--l
Figure 17. Generalized geologic map of Specimen Mountain, Colorado
(after Wahlstrom 1944). Map symbols: Trp = rhyolite plug; TrZ = topazbearing rhyolite lava; Trt = rhyolitic tuff and lava; Tl = trachyandesite
(quartz latite) lavas; pCg Precambrian gneiss and granite.
volcanic sequence and suggests that the rhyolites were derived
from them by crystal fractionation. It is difficult to test this hypothesis in the absence of more chemical information or precise
ages for the trachyandesites and the younger rhyolites. If the topaz
rhyolites were indeed derived from the trachyandesites, this complex may be unique among U.S. occurrences. However, a prominent Si02 gap (64.5% to 75.8% Si02 in nonhydrated samples) is
obvious on variation diagrams. Wahlstrom (1944) also notes
disequilibrium features in the trachyandesites-ealcic plagioclase
in intermediate composition lava and resorbed hornblende and
biotite-possibly suggesting that an even more mafic magma
mixed with a rhyolitic magma before eruption to produce the
trachyandesite. These trachyandesites are chemically similar to
the magmatic inclusions in the Spor Mountain rhyolite in Utah,
which show similar evidence of magma mixing.
Chemically, the topaz rhyolites are virtually indistinguishable from others in the western United States and have high Si, K,
and Na and low Ti, Mg, and Ca (Table 1, compare with Figure
35). Analyses of the topaz-free rhyolite plug are similar but more
variable. For example, in the plug, Si02 ranges froni 68.9% to
77.8% and Al203 ranges from 11.7% to 15.7%, perhaps as a result
of mixing or alteration. The dacites have potassic intermediate
compositions (64% Si02).
The rhyolites of Specimen Mountain are contemporaneous
with a variety of dacites and rhyolites (Corbett 1968) that were
emplaced between 27 and 28 Ma in north-central Colorado.
Although the topaz rhyolite of Specimen Mountain is a lava, the
Chalk Mountain is located on the west side of a narrow
valley that separates it from the Climax molybdenite deposit in
Lake County, Colorado (Figure 18). Cross (1886) first described
topaz and garnet from cavities within this rhyolitic plug. Based on
magnetic and gravity anomalies, Tweto and Case (1972) suggested that the Chalk Mountain stock and the stock at the Climax
mine are both apophyses of a batholith that underlies this portion
of the Colorado mineral belt. Alternatively, Chalk Mountain may
be the downfaulted upper portion of the mineralized intrusive
system (R. P. Smith, oral communication 1982). It lies on the
west side of the Mosquito fault, which separates it from the
Climax deposits. In spite of this provocative suggestion, little
information has been published regarding the rhyolite of Chalk
Mountain. Although most references to Chalk Mountain term the
rhyolitic mass a stock, Cross (1886) describes it as extrusive in
origin. The rhyolite outcrops over an area of about 4 km 2.
The rhyolite contains large phenocrysts of sanidine and
smoky quartz. Andesine (oligoclase in some) and biotite are set in
an aphanitic groundmass that contains topaz, magnetite, rare ilmenite, and apatite. Topaz also occurs in drusy quartz-lined cavities along with garnet, sanidine, biotite, and Fe oxides (Cross and
Hillibrand 1885; Cross 1886; Pearl 1939). New analyses (by
electron microprobe) of biotite from the Chalk Mountain rhyolite
show that it contains intermediate Fe/(Fe + Mg) ratios (ca. 0.5;
Figure 31). Fluorine concentrations in the biotite are not exceptionally high (averaging about 0.5 wt%) and moderate FICI ratios are typical (log FICI clusters around 1; Figure 32). In both of
these respects the biotites are more similar to those from the
Mo-mineralized Pine Grove stock in southwestern Utah (Keith
1982) than to biotites from the hydrothermally altered rocks of
the Mo deposit at nearby Henderson, Colorado (Gunbw et al.
1980). The compositions of abundant magnetite phenocrysts are
uniform eXusp .ca. 0.9) and, when combined with analyses of
sparse ilmenite, indicate that high oxygen fugacities prevailed during crystallization of the Chalk Mountain rhyolite (log f0 2 about
-10.3 at 830°C). This is about 3.5 log units above the QFM
oxygen buffer (Figure 30). Apatite-biotite geothermometry, as
formulated by Ludington (1978), gives sub-magmatic temperatures (-590°C) for apatite inclusions in biotite.
An analysis of the Chalk Mountain rhyolite published by
Cross (1886) shows that the rhyolite is similar to other topaz
Topaz Rhyolites
Volcanic Center
0
39
Deer Peak
Tomichi Dome
Bonanza volcanic field
Mount Aetna caldera (MA)
San Juan volcanic field
39 Mile volcanic field
Nathrop Volcanics
Silver Cliff-Rosita
@
Chalk Mountain
Buffalo Peaks
Cripple Creek
Hillside
31
Approximate Age
(Ma)
38-32
38
38-33
36
35-26
34
18
29
32-26
27
29
<29
Magma Type
andesite-trachyte (Iatite)
topaz rhyolite
andesite-dacite-rhyolite
rhyolite
andesite-dacite-rhyolite
andesite
Waugh Mtn basalt
topaz rhyolite
andesite-trachyte-rhyolite,
topaz rhyolite
topaz rhyolite
andesite
phonolite
andesite-trachyte (Iatite)
~ ~~SILVER
CLIFF' •.
e
"
.
San Juan
Volcanic Field
'.
San Luis Valley
•
ioo~~~
ROSITA
'.\~~
.
~Deer
~)eak
0,).
38
~'.:
Figure 18. Locations oftopaz rhyolites (solid dots) in central Colorado in relation to other middle to late
Cenozoic volcanic fields (from Epis and Chapin, 1975). The volcanic fields are listed in their approximate order of development. Age and magma composition references are given in the text.
rhyolites (Table 1). The principal differences lie in slightly higher
A1203, MnO, and MgO contents than in most topaz rhyolites
(compare Figure 35). The validity of the analysis is difficult to
assess as no modern analyses of samples from Chalk Mountain
have been published.
Chalk Mountain contains small molybdenite occurrences
and small amounts of silver ore have been removed from the
rhyolite contacts (Pearl 1939). The Chalk Mountain rhyolite is
27 ± 1.9 Ma (Tweto and Case 1972) and is one of numerous
rhyolitic plugs, stocks, and dikes adjacent to the approximately
30 Ma Climax stock (White et al. 1981), and is probably 00magmatic with this fluorine-rich rhyolite porphyry. An older
series (Late Cretaceous) of diorite to granodiorite plutons appear
to be the only other igneous rocks in the area. The mid-Tertiary
rhyolites are cut by the Mosquito high-angle normal fault, which
is part of the northward extension of the Rio Grande Rift. The
initial tectonism associated with the development of the rift in
Colorado occurred 26 to 30 Ma (Tweto 1979) and was contemporaneous with the magmatism at Chalk Mountain.
18. Nathrop, central Colorado
The Nathrop Volcanics lie on the west side of the Mosquito
Range about 10 km north of Salida, Colorado. These rhyolites
were first reported to contain topaz and garnet by Smith (1883)
and Cross (1886). Van Alstine (1969) and Schooler (1982)
further described their geology.
In general, the volcanic rocks rest directly on Precambrian
gneissic quartz monzonite and form low isolated hills along the
front of the Mosquito Range (Figure 19). At Ruby Mountain,
where topaz and spessartine garnet are found in lithophysae,
pumiceous tuff, and breccia (ca. 30 m thick), with fragments of
rocks from the Precambrian basement, are overlain by perlite (ca.
35 m thick) that apparently forms the basal vitrophyre to a capping rhyolite lava (ca. 100 m thick). The lava is flow banded with
lithophysal and spherulitic textures. Sugarloaf Mountain is also
underlain by a basal tephra and breccia unit. Scott (1975) suggests that all the rhyolite vented from Bald Mountain (Figure 19).
However, the field study of Schooler (1982) demonstrates that
32
Qal
Arkansas
Valley
Graben
Xgd
EXPLANATION
I Qal!
~
M'\""~"
i:;<]
Quaternary alluvium
Nathrop volcanics (28-29 m.y.)
rhyolite
Wall Mountain Tuff (36 m.y.)
rhyodacite
IXgdl Precambrian granodiorite
o
2km
I
I
-+-+-+-
Dike
-
Fault
and quartz diorite
Figure 19. Simplified geologic map of the Nathrop Volcanics from central Colorado (after Scott 1975; Van Alstine 1969; and Schooler 1982).
The rhyolites near Nathrop include lavas (cross-hatched) and tuffs (solid
color). Three vents probably exist near Nathrop and a fourth at Bald
Mountain.
vents also existed at Sugarloaf Mountain, Nathrop Butte, and
Ruby Mountain.
The rhyolite lavas contain sparse phenocrysts of sanidine
(Or6oAb38), oligoclase (Or8Ab81; yielding a two-feldspar temperature of 630°C), smoky quartz, and traces of biotite (Van
Alstine 1969). The groundmass contains chlorite, topaz, magnetite, and fluorite. Lithophysal cavities in the lava contain topaz,
spessartine garnet, sanidine, silica minderals, magnetite, hematite,
opal, and calcite.
The major and trace element chemistry ofthe Nathrop rhyolite has been reported by several investigators; analyses are presented in Table 1. The major element chemistry is similar to all
topaz rhyolites with high Si, K, Na, Fe/Mg, and F and low
concentrations of Ca, Ti, and Mg. Zielinski et al. (1977) report
three trace element analyses of separate phases of the lava (Table
2) and show that the vitrophyre is enriched in Mo and the lithophile elements U, Th, Be, Li, and Nb relative to most rhyolites;
these values are typical of topaz-bearing varieties (Christiansen et
al. 1980). The REE concentrations are shown diagramatically in
Figure 40g where the Nathrop rhyolite is compared to a calcalkaline rhyolite from Summer Coon volcano (Oligocene age
from the San Juan volcanic field; Zielinski and Lipman 1976).
Although the pattern is reminiscent of those from other topaz
rhyolites, it has higher La/LuN and Eu/Eu* than most. A single
new Sr-isotope analysis of a whole-rock sample suggests that its
initial 87Sr/86Sr ratio is relatively high (0.714 ± 0.0060, Table
3). The large uncertainty is a result of the high Rb/Sr ratio and
uncertainty of its age. (29.3 ± 1.5 Ma was used-the oldest of
three ages reported by Van Alstine 1969. Other K-Ar ages are
29.1 ± 0.9 and 28.0 ± 0.8 Ma.) Additionally, the effects of even
small amounts of upper crustal contamination are readily apparent in rocks with low Sr content (3.2 ppm) like the Nathrop
rhyolites. For example, 1%assimilation of a component with 300
ppm Sr and an 87Sr/86Sr ratio ofO.nO (values not unrealistic for
the Precambrian country rocks at 30 Ma) would double the
present Sr content and elevate the initial ratio from 0.706 to
0.713. Any conclusions based on this isotopic analysis must be
tempered by these facts.
The Nathrop Volcanics are spatially associated with fluorite
deposits that formed in a near-surface hot-spring environment at
temperatures of 119° to 168°C (Van Alstine 1969). Van Alstine
assigns the deposits a post-Miocene age and does not relate their
formation to the fluorine-rich rhyolite volcanism.
The topaz rhyolites from Nathrop are located near the western margin of the Thirtynine Mile volcanic field (Figure 18; Epis
and Chapin 1968). Rhyodacitic to rhyolitic ash flows are the
oldest volcanic units in the field. They were probably erupted
from centers west of the Thirtynine Mile volcanic field (e.g., the
36 Ma Wall Mountain Tuff; Chapin and Lowell 1979). The
Thirtynine Mile volcanic field consists predominantly of andesitic
lavas and breccias with minor amounts of basaltic lavas, diorite
plugs, and rhyolitic dikes (Epis and Chapin 1968; Epis et al.
1979). Presumably, the volcanism was related to the development of a composite volcano about 34 Ma. A rhyodacitic ash
flow was erupted at about the same time as the Nathrop Volcanics (the 29 Ma old Gribbles Park Tuff; Steven 1975) and overlies
some of the intermediate composition lavas. The topaz rhyolites
of Nathrop (K-Ar age 28-29 Ma; Van Alstine 1969) thus appear
to be part of an andesite-rhyodacite-rhyolite series of calc-alkaline
nature. Nonetheless, strongly alkaline volcanism at Cripple Creek
(75 km east) is contemporaneous (Steven 1975). A marked
change in the nature of the magmatism occurred about 18 Ma
when the "andesite" of Waugh Mountain was erupted in the
southern part of the Thirtynine Mile volcanic field (Wobus et al.
1979). These lavas are similar in age and composition to "silicic
alkalic basalts" of a post-Oligocene episode of bimodal basaltrhyolite volcanism (Lipman and Mehnert 1975; 133). Tweto
(1979) estimates that the Arkansas Valley graben, adjacent to the
Nathrop rhyolites, began to subside approximately 28 Ma. The
apparent superposition of the rhyolite dome complex on some of
the range-front faults (Figure 19) suggests that the tectonism and
volcanism may have been nearlyconcutrent. The tectonic and
magmatic activity of central Colorado are summarized in Figure
20.
19. Silver Cliff-Rosita, central Colorado
The Silver Cliff and nearby Rosita volcanic fields are 10-
Topaz Rhyolites
33
Colorado Chronologie Summary
Age
m.y.
o
Basalt
Alkaline
Calc-Alkaline
A-D-R
Rhyolite
Renewed uplift
ranges
II
, I
r--
10
'I
,
~
x
Miocene ·'u/l"
IMOPIulons
"---
20
Spanish Peaks
intrusions
l)
30
Comments
(iIMO PIulons
{
( )
y
+
Rio Grande Rift
~
opens
40
Figure 20. Schematic representation of magmatic and tectonic activity in southern Colorado. The ages
of topaz rhyolites (x) show that they were erupted in at least two episodes-one contemporaneous with
high-K calc-alkaline magmatism (andesite-dacite-Iow silica rhyolite) and the other group contemporaneous with basalt (as used by Lipman and Menhert 1975) and high silica rhyolite. The ages of
molybdenum-mineralized stocks (Mo plutons) of the Climax-type are also shown. The term calcalkaline is used here to denote igneous rock series whose members contain greater than about 60% Si02
and generally display continuous silica variation diagrams. Basalt or basaltic andesite is used where no
intermediate composition rocks are observed.
cated along the northeastern side of the Wet Mountain Valley
graben (Figure 18). Cross (1896) first reported topaz and garnet
from rhyolitic lavas from these volcanic fields. Subsequent investigations by Siems (1968), Kleinkopf et aI. (1979), and W. N.
Sharp (1978) have demonstrated that the volcanism in the two
areas occurred contemporaneously 32 to 26 Ma. The rhyolites in
both fields have been related to poorly documented collapse
calderas-an unusual mode of occurrence for topaz rhyolites.
Both complexes are small; neither covers much more than
30km 2.
Acording to Siems (1968) the initial eruptions at Silver Cliff
were pyroclastic eruptions that emplaced a thick sequence of
non-welded rhyolitic tuff and breccia on the Precambrian basement. Apparently these deposits accumulated in a subsiding basin
or were preserved by caldera collapse (Figure 21; Siems 1968).
Tephra accumulations exceed 600 m as exposed in mine shafts;
low residual gravity anomalies suggest they are contained in a
trough-shaped graben (W. N. Sharp, cited in Scott and Taylor
1975). Extrusion of rhyolite domes and flows with basal vitrophyres closed the volcanic cycle at Silver Cliff. The lavas are 40
to 50 m thick and have ages of 27 to 26 Ma.
The volcanic center at Rosita, centered about 8 km south-
west of Silver Cliff (Figure 21), produced a more diverse group of
rocks and spanned a much longer period of time. Early eruptions
(32 to 29 Ma) of smaIl volumes of andesite and rhyolitic tephra
from the Deer Peak volcanic center (Figures 18 and 21) were
followed by extrusions of (rhyo-)dacitic lavas. The rhyolitic products of this early phase of volcanism are grouped together as Tro
and the non-rhyolitic rocks as Tmo in Figure 21. After a 1 to 2
Ma lull a younger sequence of trachyte and trachyandesite (latite)
lavas (Tmy) was accompanied by the eruption of topaz rhyolite
lavas (27 to 26 Ma). (Most of the mafic lavas in both sequences
have potassic affinities.) Siems (1968) described this center as an
incompletely developed resurgent caldera, but documents no
large-scale collapse or structural resurgence. W. N. Sharp (1978)
cites the formation of dikes and fissures as evidence for resurgence. The duration of the activity and the interlude between the
older and younger volcanic events suggests that two perhaps
unrelated volcanic cycles are represented at Rosita. (The dates are
from Scott and Taylor 1975; and W. N. Sharp 1978.)
Modern analyses (phair and Jenkins 1975) show that the
rhyolites from Silver Cliff-Rosita are very similar to. all other
topaz rhyolites in their major constituents (Table 1). Two rhyolite
specimens have an average U content of 20 ppm and Th content
34
Silver Cliff/Rosita, Colorado
108
0
105 20'
sc
\~
105
,\0
)
~
+
40
f
Specimen Mtn
127 28
• )
IT
r~: ~::o::~
0
Boston Peak R~und
Tomichi Dome _
OHenderso~ ~
~Chmax"
~
MIn
(38)
I
5
""
of 31 ppm. Trace element analyses of vitrophyres reported by
Mutschler et al. (1985) show the rhyolites to contain 850 to 1200
ppm fluorine and relatively low concentrations of incompatible
elements like Rb (<250 ppm) and Li (<25 ppm). The rhyolites
are nonetheless enriched in Nb and depleted in Zr (-100 ppm),
Sr «20 ppm), and Ba «60 ppm) in common with other topaz
rhyolites. Many analyses (Cross 1896; Mutshler et al. 1985) of
felsitic specimens have high K20/Na20 ratios that suggest
alteration.
Deposits of hypogene and supergene silver, gold, lead, zinc,
and copper are associated with the Tertiary volcanic rocks. In the
Silver Cliff district the deposits are cavity fillings or replacements
in the lavas and tuffs. The most productive mines are breccia
pipes that formed within the Precambrian gneisses; their relationship to the volcanic cycle is unclear but their similarity to vent
facies rocks suggest that they are also related to the younger
rhyolites. In the Rosita Hills district, mineral deposits occur in
fIssure veins and along faults in all of the volcanic units.
The Wet Mountain Valley is a tectonic basin that began to
form in late Cretaceous or early Eocene times. However, major
uplift of the flanking ranges occurred in early Miocene to late
\
\
39
"\
\~28-29)
Sitv", Cliff (26)
La~~8~lty ~ ~\ \y \\
38
\
km
Figure 21. Generalized geologic map of Silver Cliff-Rosita area, Colorado (after W. N. Sharp 1978). Topaz-bearing rhyolitic lava flows and
dome (Trf) are underlain by tephra and sedimentary rocks (Trt) that
accumulated in a small NW-trending graben at Silver Cliff (8C). Rhyolites surround a central core of older (Tmo) and approximately contemporaneous (Tmy) mafic volcanic rocks (andesite, trachyandesite, and
trachyte) at Rosita (R). Vents (*) and mineralized breccia pipes (+) are
also shown. Both volcanic centers are located near a fault (dashed) that
separates a sediment-(QTs)-filled graben from Precambrian gneisses and
granitic rocks (peg) of the Wet Mountains. A portion of the Deer Peak
volcanic center is also shown in the SE comer of the map.
LNathrop
r
~
•
38
o
40
~ ~
/Chalk Mtn
(27-28)\( -9
38
104
108
107
I~; /\ I
106
104
Figure 22. Faults with major Neogene movement in the area of the Rio
Grande rift in Colorado (after Tweto 1979), compared to the location of
topaz rhyolites (filled circles) described in the text. The ages of the
rhyolites are shown in parentheses. The Rio Grande rift began to form 30
to 27 Ma (Eaton 1979; Tweto 1979), about the same time as the topazrhyolite magmatism was initiated in the region. The locations of topazbearing molybdenite deposits (0) younger than 30 Ma (White et aI.
1981) are shown for comparison.
Pliocene time (Scott and Taylor 1975). Apparently this tectonic
episode occurred shortly after the development of several volcanic centers on the eastern side of the graben (Scott and Taylor
1975)-Silver Cliff-Rosita (andesite to rhyolite, 32 to 26 Ma),
Deer Peak (andesite to trachyandesite, 38 to 32 Ma), and Hillside
(andesite to trachyandesite, less than 29 Ma). The locations of
these fields are shown in Figure 18. The tectonism and magmatism of the Wet Mountains appear to be part of the development
of the Rio Grande rift system in Colorado (Tweto 1979; Eaton
1979) and are contemporaneous with the eruption of other topaz
rhyolites in Colorado (compare Figure 22).
20. Tomichi Dome, central Colorado
Tomichi Dome is located about 35 km east of Gunnison,
Colorado. Stark (1934) reported topaz from the rhyolite as part
of a study of heavy minerals from Tertiary "intrusions" in central
Colorado. The geology of the rhyolite (Figure 23) is described by
Stark and Behre (1936) and briefly by Ernst (1980).
Topaz Rhyolites
35
Tomichi Dome, Colorado
o
!
1
2
I
!
krn
Tertiary rocks
t~;:H?ll
Massive rhyolite
t,,~.:j
Spherulitic rhyolite
•
Summit
- - - - Fault
Figure 23. Generalized geologic map of Tomichi Dome, Colorado (after Stark and Behre 1936).
As its name suggests, the rhyolite caps a domical mountain others of this group they contain relatively low concentrations of
that rises about 600 m above a generally flat region. The rhyolite U (4 to 7 ppm) and Rb «300ppm). Felsites contain about 0.17
was emplaced through Cretaceous sedimentary rocks (sandstone, wt%F.
shales, and limestones). The initial eruptions resulted in the emF. E. Mutschler (written communication, 1983) reports a
placement of a poorly-exposed basal explosion breccia and tuff whole-rock K-Ar age of38 Ma for Tomichi Dome. This is 9 Ma'
over 200 m thick. Ernst (1980) identified a breccia pipe (ca. 500 older than any other topaz rhyolite in Colorado and makes it
m across) on the northeastern margin of the rhyolite. Fragments contemporaneous with the calc-alkaline magmatism of the Boof sedimentary rocks and Precambrian granite are included in the nanza and San Juan volcanic fields to the south (Varga and Smith
breccia pipe and in the pyroclastic deposits. Overlying the tuff is a 1984; Lipman et aI. 1976).
phenocryst-poor rhyolite lava flow or dome. The lower 300 m of
lavas have spherulitic textures and the upper part is a denser 21. Boston Peak, central Colorado
equigranular rock with a granophyric groundmass and prominent
flow banding. The exposed dome is 2 to 3 km in diameter and the
Three vent complexes of topaz rhyolite are exposed at Bosvolume of rhyolitic rock exposed is approximately 3 to 4 km 3. ton Peak (Ernst 1980), which lies about 40 km northwest of
A sill, composed of a rock similar to that in the main mass, Tomichi Dome in Gunnison County, Colorado (Figure 22). All
intrudes shales at the base of the dome. The sill is 6 to 10m thick of the rhyolite plugs are small; the largest is only about 800 m
and a thermal contact aureole extends away from it for several across. No bedded pyroclastic deposits are associated with the
meters into country rock.
emplacement of the Boston Peak plugs but a breccia pipe is
The principal phenocrysts are biotite, sanidine, oligoclase, exposed adjacent to one of the rhyolite vents. Angular to suband smoky quartz. The matrix consists of these same minerals rounded fragments of underlying sedimentary units and Precamand some glass. Magnetite, zircon, and apatite are magmatic ac- brian granite are included in the breccia pipe.
cessories. As devitrification products, topaz and garnet occur in
The rhyolites contain phenocrysts of resorbed quartz, saniirregular clots, and biotite crystals occur in radiating groups in the dine, albite (An 3 to 6), biotite, zircon, and Fe-Ti oxides. A
upper part of the lava. Garnet, of unspecified association, is also vitrophyre is preserved at the margin of one of the vents, otherpresent in the basal tuffaceous sequence. Small quantities of wise the phenocrysts are contained in fine-grained, flow-banded,
hornblende, ilmenite, and titanite were found in heavy mineral and felsitic matrix. Lithophysae with concentric shells of quartz
separates but were not observed in thin section (Stark 1934).
and topaz are common; zeolites flil some of the cavities. Fluorite
Ernst (1980) analyzed eight specimens from Tomichi Dome crystals occur in fractures in the phenocrysts and, along with
for major and trace element concentrations. An average of two topaz, garnet, and tourmaline, are post-magmatic. Muscovite,
analyses of the upper topaz-bearing part of the dome is given in probably as an alteration product, occurs as small grains in the
Table 1 and 2. In many ways the analyses are typical of other groundmass of two of the rhyolite plugs.
topaz rhyolites from the western United States, but compared to
The average of six analyses (Ernst 1980) of rhyolites from
36
the three vents at Boston Peak are presented in Tables 1 and 2.
The major element compositions of these rhyolites are in all ways
like their counterparts elsewhere; they show none of the compositional "anomalies" of the Tomichi Dome rhyolites. A vitrophyric
specimen contains 0.51% F; felsites contain 0.12 to 0.49% F. In
accord with these relatively high F concentrations, the rhyolites
have relatively high concentrations of incompatible trace elements, e.g., Li (90 to 270 ppm), Rb (390 to 820 ppm), U (7 to 24
ppm; highest in the vitrophyre), and Nb (35 to 160 ppm). The
ppm),
feldspar-compatible elements, Ba (<75 ppm) and Sr
are strongly depleted.
The age of the rhyolites is not known, but they lie on a
NW-trend that includes the rhyolites at Mt. Emmons and Redwell Basin (17 Ma-a molybdenum mineralized rhyolite
complex that bears topaz; Thomas and Galey 1982; J. E. Sharp
1978), Treasure Mountain dome (12.5 Ma; Obradovich et aI.,
1969), Round Mountain (14 Ma; Cunningham et aI. 1977), and
Tomichi Dome (38 Ma Mutschler, written communication
1983). With the probable exception of Tomichi Dome, these
plugs appear to be part of a post 20 Ma bimodal suite of basalt
and high-silica rhyolite typical of southwestern Colorado (Figure
20).
Lake City, Colorado
\
\
,...
\
o
Te
5
L...L..L...L.J.
km
«10
Te
22. Lake aty, southwestern Colorado
An ENE-trending line of topaz-bearing rhyolite plugs, sills
and laccoliths is exposed in the northwestern portion of the San
Juan volcanic field (Steven et aI. 1977). The trend starts about 5
km north of Lake City, Colorado. The belt of intrusions extends
for 20 km and consists of about 10 separate bodies that are all
described as being mineralogically and chemically similar to one
another. The geologic setting, mineralization potential, and trace
element chemistry of these rhyolites are reported by Steven et aI.
(1977), and isotopic analyses (Pb, Sr) of one plug are presented
by Lipman et aI. (1978a). The petrology of these rhyolites and the
rocks of the Lake City caldera are presently being studied by R.
A. Zielinski and K. Hon of the U.S. Geological Survey.
The rhyolites (Figure 24) were emplaced as discordant intrusions, generally less than 1 km across, within the volcanic and
sedimentary fill of the Uncompahgre caldera, which formed 28
Ma. All of the intrusions are small, with maximum dimensions of
1 to 2 km; Although none of the rhyolites are reported to be
extrusive, the preservation of marginal vitrophyres and the formation of lithophysae suggests that they were emplaced at very
shallow levels. K-Ar dating indicates they are 18.5 Ma (Lipman
et aI. 1978a); at least 8 Ma younger than other dated topaz
rhyolites in Colorado.
The rhyolites are generally devitrified and light gray in color.
Reported phenocrysts include abundant quartz and sodic sanidine
along with sparse biotite and oligoclase. Titanite is a prominent
microphenocryst (R. A. Zielinski, oral communication, 1982)
along with apatite, zircon, and sparse Fe-Ti oxides (Ernst 1981).
Small crystals of topaz and fluorite occur within some cavities
and along fractures.
~
Post-caldera mafic laval
I@ I
Granitic Intrusive Rocks
~ Rhyolite plugs (18-19 m.y.l
r:fs;l
Sunshine Peak Tuff
~
(22m.y.)
EXPLANATION
.'. ._Caldera
:....... ....
wall
Older ash flows and
lavas (30-26 m.y;)
Early intermediate composition
lavas arid tuffs (35-30)
Figure 24. Generalized geologic map of the Lake City area, southwestern
Colorado (after Steven et al. 1977; and Lipman et al. 1978). Plugs of
topaz rhyolite (cross-hatched) are nestled within the rim of the Uncom,
pahgre caldera (formed 28 Ma), but are much younger (18 to 19 Ma).
The rhyolites are also younger than the Silverton/San Juan (28 Ma) and
the Lake City (22.5 Ma) calderas and probably represent a distinct
volcanic episode in this area.
Ernst (1981) reports chemical an,alyses of the rhyolite plugs
(Table 1) that show them to be typical high-silica rhyolites. Many
of the samples have high Kz0/NazO, indicating post-magmatic
redistribution of alkalies; no major-element analyses of vitrophyres have been published. A variety of trace element analyses
reported by various investigators is given in Table 2. Ernst
(1981) reports that the rhyolites have relatively high Sr (ave. 126
ppm) and Ba (600 ppm) for topaz rhyolites, but his analyses also
demonstrate an enrichment in Rb and Li. Semi-quantitative anal-
Topaz Rhyolites
yses (Table 2) also show that the rhyolites are enriched in other
incompatible trace elements: Be (7 to 20 ppm), Nb (20 to 80
ppm), Pb (30 to 70 ppm), Sn (up to 10 ppm), Y (up to 20 ppm),
and Mo (up to 15 ppm). Delayed neutron activation analyses
show that the rhyolites are also enriched in U (10 to 42 ppm) and
Th (26 to 64 ppm). Fluorine concentrations, even in "vitrophyres," range erratically from 500 to 1300 ppm, values that are
much lower than for most other topaz rhyolites. Vitrophyres have
higher values of U, Be, and Mo than their devitrified equivalents
and Steven et al. (1977) suggest that these elements were released
from the lavas during cooling and crystallization in the form of
halogen complexes. The REE patterns for the topaz rhyolites of
Lake City display prominent middle REE depletion (R. A. Zielinski and K. Hon, written communication 1982) that may have
resulted from the fractionation of titanite. These patterns are similar to those of the titanite and topaz-bearing rhyolites from the
Mineral Range, Utah. The Sr- (0.7054) and Pb- isotopic composition of these rhyolites are relatively unradiogenic (Table 3) and
imply that they may have arisen from a source with relatively low
-Rb/Sr, Th/Pb, U/Pb and Th/Pb ratios. Lipman et al. (l978a)
suggest that the source for these rhyolites (and other similar
Miocene-Pliocene rhyolites) is in amphibolite-facies lower crust
because of the inferred low Th/U ratio of the source. Low Th/U
ratios are atypical of some, but not all, exposed granulite-facies
metamorphic terranes.
All of the young Lake City rhyolites are anomalously radioactive (scintillometer readings over them are about two times
that found over their hosts-Steven et al. 1977) and they have
been extensively prospected for uranium. In 1960, a small quantity of uranium ore was removed from a supergene deposit near
one of these plugs (Steven et al. 1977). Much of the mineralization in the San Juan Mountains is associated with the emplacement of young (less than 22 Ma) silicic magmas that are grossly
similar to the Lake City rhyolites but generally topaz-free (Lipman et al. 1976).
. The initial eruptions (35 to 30 Ma) in the San Juan volcanic
field formed clusters of central volcanoes composed of andesitic to
rhyodacitic lavas and breccias. Large ash-flow eruptions of more
silicic (about 72% Si02) magmas occurred 30 to 26 Ma, at the
same time as the Rio Grande rift was developing farther east (Figures 20 and 22). Eruptions of minor volumes of andesitic to dacitic
lavas accompanied resurgence of the calderas. About 20 to 25 Ma
(early Miocene) volcanism in the San Juan Mountains had
changed to a bimodal assemblage of basalt (or basaltic andesite)
and "high-silica alkali rhyolite" (Lipman 1981). The topaz rhyolites at Lake City are included in this latter group. The only caldera
in the San Juan volcanic field to develop during this episode was
the Lake City caldera (22.5 Ma) with which the topaz rhyolites
(18.5 Ma) are spatially associated. It is important to note that the
topaz rhyolites of the San Juan Mountains are younger than other
dated topaz rhyolites from Colorado (26 to 38 Ma versus 18.5 Ma)
and they do not occur within the Rio Grande rift system (Figure
22). In addition, contemporaneous mafic lavas are alkali basalts,
not andesites or their derivatives (Lipman and Menhert 1975).
37
Topaz rhyolites in Colorado: A summary
Although geochemically similar, there appear to be at least
two separate groups of topaz rhyolites in Colorado (Figure 20)an older Oligocene group emplaced before about 25 Ma (represented by Specimen Mountain, Chalk Mountain, Nathrop, Silver
Cliff-Rosita, and possibly Tomichi Dome, which is 38 Ma) and a
younger Miocene group emplaced after about 22 Ma (represented by the Lake City group and perhaps the plug at Boston
Peak). The Oligocene group is contemporaneous with the waning
stages of generally calc-alkaline volcanism (characterized by extended Si02 variation diagrams) in the San Juan Mountains and
other centers on the flanks of the Rio Grande rift. Their distribution coincides with the locus of Neogene tectonism as they preceded or were contemporaneous with the development of the rift.
The topaz-bearing stocks at the Climax and Henderson molybdenum mines are probably part of this group. The younger group of
topaz rhyolites were erupted as part of a clearly bimodal group of
basaltic (or basaltic andesite) lavas and rhyolitic lavas and tuffs.
The locations of the younger rhyolites are not limited to the rift
proper, instead they occur on the western side of the rift. In
addition, they are not clearly associated with faulting related to
the still active Rio Grande rift (Tweto 1979). The rhyolitic stocks
at Mount Emmons and Redwell Basin (topaz-bearing Climaxtype molydenite deposits; J. E. Sharp 1978; Thomas and Galey
1982) are included in this younger group.
23. East Grants Ridge, west-central New Mexico
Grants Ridge is a discontinuous basalt-capped mesa that
slopes southeast towards Grants, New Mexico. The eastern part
of Grants Ridge is underlain by a rhyolite dome complex that
contains topaz and garnet in lithophysae (Kerr and Wilcox
1963). The volcanism is related to the development of the andesitic composite volcano at Mount Taylor, which is centered about
12 km to the northeast (Figure 25).
The initial volcanic activity at Grants Ridge is represented
by a rhyolitic tuff. The tuff includes black obsidian bombs and
large xenoliths of Precambrian(?) granitic rocks. Precambrian
rocks are not exposed at the surface and must lie several 1OOOs of
meters deep in the vicinity of Grants Ridge (Thaden et al. 1967).
The tephra probably formed a low tuff cone and are comprised of
pyroclastic flow, surge, and fall deposits. Separate perlite and
felsite domes rose through the tuff; both contain topaz and garnet.
The perlite dome is intimately flow banded and Kerr and Wilcox
(1963) suggest that hydration occurred under magmatic conditions and is not the result oflow temperature hydration of obsidian at the surface. The lava dome consists of a felsitic rhyolite
core surrounded by a collar or rind of obsidian and perlite. Flow
banding is well-developed in all phases of the body and suggests
that the dome has a concentric internal structure. Lithophysae,
some 10 cm in diameter, are common in the outer part of the
dome. Rhyolite from the dome has K-Ar dates of 3.2 Ma (Bassett
et al. 1963) and 3.3 Ma (Lipman and Mehnert 1980). The rhyo-
38
pTr
pTr
Qb
r---.---I'l-~R
a
Cubero
35~----~""':';";'~-_J-_-------_"""L.-I""'--------""
o
10
I
I
20
I
K II o,m e t e r s
Figure 25. Generalized geologic map of volcanic rocks in the region surrounding Grants Ridge,
New Mexico (from Lipman and Mehnert 1980). The inset shows the relationship of Grants Ridge
(GR) and the Mount Taylor volcanic field to the Raton (R)-8pringerville (8) or Jemez lineament
of New Mexico and Arizona; late Cenozoic volcanic fields in the region are outlined. Map symbols:
Qb - Holocene and Pleistocene tholeiitic basalt and the alkalic Zuni Canyon basalt flow;
Qbt - Pleistocene alkalic basalts of the high mesas surrounding Mt. Taylor; Qa - Pleistocene andesites,
basalts, and pyroclastic flows of Mt. Taylor; QTr - Pleistocene and Pliocene rhyolite domes and
tuffs including the topaz-bearing rhyolite at East Grants Ridge; Tb - Pliocene basalt; pTr - pre-Tertiary
sedimentary rocks.
lites are partially covered by slightly younger olivine basalt flows
that contain inclusions of the rhyolite. Several scoria cones represent the final activity at the basalt vents. Although the basalt
extrusions have not been dated, similar mafic lavas that form part
of the Mount Taylor volcanic field have ages that range from 2.9
to 1.6 Ma (Lipman and Mehnert 1980).
The rhyolites at East Grants Ridge are all phenocryst-poor
but contain small phenocrysts of quartz, sanidine, and sodic plagioclase. Devitrification of glass has produced a light-grey felsitic
groundmass in most of the dome. Vapor-phase crystallization of
garnet, topaz, and tridymite occurred in lithophysae that are present in obsidian, felsite, and some phases of the perlite dome.
Baker and Ridley (1970) report an average of two rhyolite
analyses from the Mt. Taylor volcanic field, but do not give the
locations of the samples. The average is similar to other topaz
rhyolite analyses reported in Tables 1 and 2 and shows the characteristic depletion ofTi, Mg, and Ca. Our trace element analyses
also show that the Grants Ridge rhyolites are similar to other
topaz rhyolites with enrichments ofF (0.5% in a vitrophyre), Li,
Rb, and Sn. Baker and Ridley (1970) report two rhyolite analyses
for Rb (602,537 ppm) and Sr (nd., 2 ppm) that are in accord
with these analyses and typical of other topaz rhyolites. Zielinski
(1978) reports that the uranium contents of the lavas at Grants
Ridge range from 7 to 8 ppm, relatively low concentrations for
topaz-rhyolite lavas. Baker and Ridley (1970; citing P. Pushkar)
report that a dacite from Mt. Taylor has an initial 87Sr/86Sr ratio
of 0.7193 and believe that it was derived by mixing of mafic
magma with a rhyolitic magma (represented by the domes and
tuff) derived by partial melting of the crust.
No known mineralization is associated with the rhyolites.
The Cenozoic volcanism is unrelated to the large uranium deposits of the Grants region that are Late Jurassic to mid-Cretaceous
in age (Brookins 1980).
Mount Taylor consists of an andesite composite cone with
early high-K basaltic andesites (trachytes) or dacites (4.44 Ma;
Lipman and Mehnert 1980) and rhyolitic tuffs exposed in a central amphitheater. The cone is composed of porphyritic andesite
lavas, and developed 2.4 to 2.9 Ma. The surrounding mesas are
capped by a differentiation series of high-K basalts to andesite
(alkali basalts to trachytes; Crumpler 1980) that developed contemporaneously 4.3 to 1.5 Ma (Lipman and Mehnert 1980; Lipman et al. 1979). The rhyolites and basalts of Grants Ridge are
part of the peripheral volcanism. The development of the Mt.
Taylor volcanic field was concurrent with minor north-northeast
(r
I
!'
I.
Topaz Rhyolites
39
faulting in the area. Lipman and Mehnert (1980) relate the development of the volcanic field to activity on the east-northeasttrending Jemez Lineament or Springerville-Raton zone (Figure
25). Volcanism of similar age occurred elsewhere along this belt
that stretches from southwestern Arizona to northeastern New
Mexico.
24. Black Range, southwestern New Mexico
The Taylor Creek Rhyolite is a topaz-bearing lava that crops
out in the northern Black Range of southwestern New Mexico
(Fries 1940; Fries et al. 1942; Ericksen et al. 1970; Lufkin 1976,
1977; Correa 1980). The rhyolite has been extensively studied in
part because it contains low-grade concentrations of tin as cassiterite and wood tin. The Black Range lies on the eastern margin
of the Mogollon-Datil volcanic field that was active from about
40 to 20 Ma (Elston and Bornhorst 1979).
The Taylor Creek Rhyolite, which has a volume of 130 km 3
(Rhodes 1976), was emplaced as a series of endogenous domes (5
to 15 separate vents may exist). Pyroclastic material that preceded the lava eruptions is exposed at the present margins of the
lavas and presumably underlies the lavas as well. Locally, an
autoclastic breccia forms the base of the flows. Flow-banding is
well-developed and reveals large folds that show the internal
structure of the domes and lava flows. The rhyolite has K-Ar
(sanidine) ages of 24.6 Ma (Elston 1978) and 27.7 Ma (Ratte et
al. 1984) from the same locality. In any case, the extrusion of the
domes probably spanned a considerable length of time. According to Ratte et al. (1984) the Taylor Creek Rhyolite may represent ring-fracture volcanism of the Bursum caldera (source of the
Bloodgood Canyon and Apache Springs Tuffs), which formed
29-28 Ma. About 30 m of Bloodgood Canyon Tuff overlie the
Taylor Creek Rhyolite (Ratte et al. 1984) discounting the suggestion that the Taylor Creek Rhyolite is the devolatilized posteruption residue of the Bloodgood Canyon magma chamber as
suggested by Rhodes (1976). Nonetheless, it is just as likely to be
unrelated to the development of any of the large calderas (T.
Eggleston, oral communication 1985). The distribution of the
lava and the location of some of the calderas are shown in
Figure 26.
Most of the Taylor Creek Rhyolite is composed of devitrified and vapor-phase altered lava. It contains 20 to 40 percent
phenocrysts of quartz and sanidine with lesser amounts of plagioclase. Biotite and ferrohornblende are present in most samples;
ferroaugite (Mg24Ca3SFe41) occurs in biotite and hornblende-free
flows (Correa 1980). Accessory phases include zircon, titanite in
some specimens, Fe-Ti oxides (mostly titaniferous magnetite),
and fayalite(?) (Lufkin 1976, 1977). Vapor-phase minerals that
occur in miarolitic cavities and veinlets include quartz, alkali
feldspar, hematite, bixbyite «(Mn,Feh03), pseudobrookite
(Fe2TiOs), cassiterite, topaz, monazite, fluorite, (Lufkin 1976),
garnet (Fries et al. 1942), and beryl (Kimbler and Haynes 1980).
Inclusion-ridden topaz crystals in miarolitic cavities at Round
Mountain reach lengths of 2 to 3 cm. Although granophyric
lIT] Calc-alkalic
rhyolite
50
,
KM
Figure 26. Tectonic map of southwestern New Mexico showing the
relationship of the topaz-bearing Taylor Creek Rhyolite to the major
Cenozoic faults and calderas in the region (Elston 1978, 1984; Ratte et
aL 1984). Mid-Tertiary calderas are indicated by dashed lines and normal faults by heavy solid lines. The Mogollon caldera (M) was the
source of the Cooney Tuff and appears to have formed 34 Ma. The Gila
Cliff Dwellings caldera (GCD) formed 30 to 29 Ma and may have been
the source of the Davis Canyon and/or Shelley Peak Tuffs. The large
Bursum caldera (B) formed 29 to 28 Ma and was the source of the
Bloodgood Canyon and Apache Springs Tuffs. The relationship of the
Taylor Creek Rhyolite (24.6 to 27.7 Ma) to these calderas is open to
question.
textures are the most common, spherulitic devitrification textures
are also present. The upper parts of the flows are vapor-phase
altered and contain the most abundant gas cavities.
The Taylor Creek Rhyolite is chemically similar to most
topaz rhyolites (Table 1) with high Si, K, Na, and Fe/Mg and
low Mg, Ca, Ti, and P (Correa 1981). Vitrophyres from the flow
contain up to 0.4% F and are enriched in Sn and Rb (Table 2).
Eggleston and Norman (1984) mistakenly reported that the lavaS
were Cl-rich: The Taylor Creek Rhyolite contains less D, Ta, and
Th than some of topaz rhyolites from western Dtah (Figure 35),
but is still enriched when compared to other high-silica rhyolites
from the region. In addition, it is decidedly rich in Y (> 100 ppm).
The REE pattern of the rhyolite (Figure 40b) has a deep Eu
anomaly (Eu/Eu* = 0.07), low La/LuN (2.4), and low La/CeN
(Correa 1981), typical of topaz rhyolites in general. Eggleston
and Norman (1984) report that the delta 180SMOW value for the
rhyolite is 8 permil, indicating a crustal history for the magmas'
source. A single Sr-isotope analysis of a whole rock sample of the
Taylor Creek Rhyolite suggests that its initial 87Sr/86Sr is very
high (0.71583 ± 0.0028; Table 3). The usefulness of this result is
questionable both because the sample contains only 3.7 ppm Sr
40
New Mexico Chronologie Summary
Age
m.y.
o
Basaltic
Calc-Alkaline
Basaltic Andesite
Rhyolite
'i.
\
X
Comments
Mt Taylor
Intraplate Block
Faulting
Jemez
10
20
A
I~
I
x
Questa
Back-arc
Extension
+
Rio Grande Rift
opens
30
40
?
Figure 27. Schematic representation of magmatic and tectonic activity in New Mexico (after Elston and
Bornhorst 1979). Topaz rhyolites (x) appear to have erupted in two episodes-one contemporaneous
with the basaltic andesite suite of Elston and Bornhorst and the other contemporaneous with a younger
basalt to andesite suite that includes the andesites and alkalic basalt to trachyte suite at Mt. Taylor. The
age of the Questa molybdenite and peralkaline volcanism there are also shown (0).
(thus it could have been easily contaminated at the magmatic
stage) and because the sample was a hydrated vitrophyre. (Hydration may radically change the Sr content and hence Srisotopic composition of ·vitrophyres: e.g., Hargrove 1982). A
detailed Br-isotope investigation of the Taylor Creek rhyolite is
being conducted by D. Norman and T. E. Eggleston (oral communication 1985).
Tin mineralization in the form of cassiterite and wood-tin
veinlets (Lufkin 1976, 1977) deposited as a result of fumarolic
activity is common in the upper parts of the Taylor Creek Rhyolite (Correa 1980; Eggleston and Norman 1984). The mineralization is restricted to the flanks of intensely vapor-phase
recrystallized zones just below the carapace of the domes~ Fluid
inclusion and oxygen isotope studies indicate that quartz and
topaz crystallized from saline magmatic fluids (8180 = +6 to +10
%0) at temperatures over 600°C (Eggelston and Norman 1984)
and miarolitic cassiterite-hematite-quartz was deposited at
temperatures above 500°C (Rye et al. 1984); probably also froni
magmatic fluids derived by degassing of the rhyolite lavas. Late
cassiterite crystallized at 150 to 200°C from boiling fluids with
calculated 8 180 of ~6 to 0 %0 (Eggleston and Norman 1984).
The wood-tin depositing fluids appear to have contained a large
meteoric component. Placer deposits of Sn derived from the lavas
have been mined on a small scale. Wood-tin mineralization is
also associated with topaz rhyolites in Nevada and Mexico (see
below). No known F, Be, or U mineralization is directly aSsociated with the Taylor Creek Rhyolite. However, F, Be, Fe, W,
Sn, and U mineralization occurs in skarns and along faults near
Iron Moutnain, east of the Black Range (Jahns 1944; Hillard
1969). A specimen from the aplitic intrusions associated with the
skarns has a K-Ar age of 29.2 Ma (Chapin et al. 1978) indicating
that these magmas are approximately the same age as the Taylor
Creek Rhyolite.
The mid-Terti~ry magmatism of the Mogollon-Datil province has been interpreted to be the product of three overlapping
magma suites (Elston and Bornhorst 1979; Figure 27). A calcalkaline andesite to rhyolite suite (43 to 29-28 Ma) produced
composite volcanoes and slightly younger compositionally zoned
silicic ash-flow tuffs. This volcanism·appears to have been contemporaneous with subduction of oceanic lithosphere at the
continental margin. Back-arc (or intra-arc) extension led to the
development of a bimodal basaltic andesite and high-silica rhyolite suite that extended from 30 to 19 or 18 Ma. The third stage is
represented by flows of tholeiitic and alkalic basalt and high-silica
rhyolite with ages ranging from 19 Ma to less than 1 Ma. This
suite was erupted during a period of block faulting. A somewhat
different picture was presented by Ratte et al. (1984).They group
the high-silica rhyolitic ash flows and lavas of Elston's and Born-
Topaz Rhyolites
horst's intermediate stage with the oldest suite, interpreting them
as being the first erupted parts ofzoned (dacite-rhyolite-high silica
rhyolite) magma chambers. Ratte et al. (1984) suggests that the
"ring-fracture" rhyolites were also erupted from these zoned
chambers. These authors place voluminous post-caldera andesites
in a "fundamentally basaltic" suite (26 to 23 Ma) but also recognize a bimodal suite of basalt and high-silica rhyolite that appeared about 19 Ma, long after the eruption of the Taylor Creek
Rhyolite. Rhodes (1976) and Ratte et al. (1984) suggested that a
composite granitic batholith underlies the volcanic field. The
topaz rhyolites may have evolved by fractional crystallization and
dehydration of one of these magma bodies following collapse of
one of the large calderas (Rhodes 1976, suggests the Gila Cliff
Dwelling caldera and Ratte et al. 1984, suggest the Bursum caldera). However, Correa (1980) speculates that the Taylor Creek
Rhyolite may have arisen from a separate magma body not directly related to these calderas.
25. Saddle Mountain, eastern Arizona
Information about the topaz rhyolite near Saddle Mountain
in the Galiuro Mountains of southeastern Arizona is sketchy.
Anthony et al. (1977) state that topaz, together with pseudobrookite, spessartine garnet, and bixbyite, occurs in a rhyolite in
the Winkelman area. The topaz rhyolite was not located during a
brief visit· to the area. The rhyolite nearest Saddle Mountain
(which is formed by resistant andesitic lavas) is described by
Krieger (1968) as an intrusive plug. Biotite from the rhyolite has
a K-Ar age of 61 Ma. Younger (ca. 24 Ma) rhyolitic lavas occur
3 to 5 km northeast of Saddle Mountain.
26. Burro Creek, western Arizona
Rhyolitic lavas with lithophysae containing garnet and/or
topaz occur in eastern Mohave County, Arizona (Burt et al. 1981;
Moyer 1981). Topaz and garnet occur in a mesa identified as
"Negro Ed" on the 7lf2 minute quadrangle of the same name
(V.S. Geological Survey 1980). The hill is formed by the remnants of a flat-topped rhyolite lava (Figure 28). A basal vitrophyre (l to 3 m thick) is locally disrupted and forms part of a
flow-produced breccia. Locally, a thin (less than about 50 m
thick in places thinning to about 1 m) pyroclastic deposit and/or
explosion breccia with fragments of the Precambrian country
rock is exposed at the base of the complex. The dome is more
than 200 m high and 2 km across. Several other small domes in
the area contain vapor-phase garnet in lithophysae or along more
coarsely crystalline flow bands. The ages of the rhyolites are not
known, but they are probably late Miocene or Pliocene. Mesacapping basaltic lavas in the area have K-Ar ages of 8 to 9 Ma
(Shafiqullah et al. 1980).
The rhyolites are all phenocryst-poor (3 to 5%) with sparse
phenocrysts of quartz, sanidine, oligoclase, biotite, and Fe-Ti oxides. Spherulitic devitrification is the most common groundmass
texture in the felsites. Moyer (1981) reports that the vapor-phase
41
'"'",,'" :
!f1
'0
'"
'
","<
'",
'"
0",
I
I
,I
D',
N
Negro Ed
,,
,
,,
Tvsp
p€
~TVSP
3.4 37' 30"
Old
() ()
LU::;;:
l..-I
o
Tvs
1 kilometer
>10-<
<
::c,>
0'<
:::;1>-
Figure 28. Geologic sketch-map of the volcanic geology near Burro
Creek, Arizona (after Burt et al. 1981b). The topaz-bearing rhyolite
dome at Negro Ed (Tvst) is underlain by a pyroclastic unit (Tvsp). Other
small rhyolite domes (Tvs) in the area contain vapor-phase garnet and
are underlain by cogenetic tephra (Tvsp). The rhyolites lie on a Precambrian basement of gneisses and granites (pC). Normal fault and sense of
movement shown as dashed line.
garnets are spessartine-almandine solid solutions (Figure 34) with
up to 3000 ppm fluorine (analyzed by ion chromatography and
electron microprobe). They may represent the "almandine" occurrence at the southern end of the Aquarius Range mentioned
by Anthony et al. (1977).
An average chemical composition of vitrophyres from
garnet- and topaz-bearing lavas is given in Table 1 and trace
element analyses in Table 2 (Moyer 1981). The average is typical
of topaz rhyolites. Fluorine concentrations range from 1400 to
2600 ppm; chlorine concentrations are 4 to 5 times lower.
Although a number of lithophile-element deposits (Be, W,
V) occur in the region" most are associated with anorogenic
granites of Precambrian age (Anderson 1983). However, Burt et
al. (1981) have suggested that the fluorine-rich rhyolites or similar rocks may have been the source some of the uranium deposited to the south at the Anderson mine (Sherborne et al. 1979).
The topaz rhyolites of Burro Creek are part of a more or less
bimodal assemblage of late Cenozoic volcanic rocks from western Arizona (cf. Suneson and Lucchitta 1983). In the vicinity of
Burro Creek, Moyer (1981) has delineated a series of low-silica
rhyolite lavas (ca. 70% SiOz) and sub-alkaline basalts to basaltic
andesites (48 to 56% SiOz) that are broadly contemporaneous
with, but slightly younger than, the aphyric (topaz-bearing) lavas.
Tuffs emplaced during dome-forming eruptions are interlayered
with basaltic lavas in the region. A period oflate Cenozoic listric
normal faulting was followed by high-angle normal faulting between 14 and 7 Ma in western Arizona (Suneson and Lucchita
1983). The magma-tectonic setting of this region appears to be
similar to that found elsewhere in the Basin and Range province
where topaz rhyolites occur in association with basaltic lavas and
extensional faulting.
42
OTHER "TOPAZ RHYOLITE" OCCURRENCES
Several other occurrences of topaz in or associated with
volcanic rocks have been reported that: 1) are not sufficiently
documented to warrant separate discussion; 2) are not Cenozoic
in age; or 3) are not from the western United States. Brief discussions of these "other" occurrences are presented here for completeness and for comparative purposes.
Other Cenozoic occurrences, western United States
Trimble and Carr (1976) report "topaz(?)" in heavy mineral
separates from the tuff of Arbon Valley along the southern margin of the Snake River Plain in Idaho. A topaz locality in southcentral Idaho is shown on Shawe's (1976) distribution map of
topaz rhyolites in the United States. At this locality, topaz occurred with fayalite, allanite, chevkinite(?) (D. R. Shawe, oral
communication 1982) and zircon in a stream sediment concentrate. The stream drains Tertiary rhyolites (rhyolites of Magic
Reservoir) on the north side of the Snake River Plain and part
of the Idaho batholith and thus the topaz may not be from a
rhyolite at all. Leeman (1982a, b) reports that the rhyolite dome
at Moonstone Mountain (which is drained by the sampled
stream) is a high-silica rhyolite lava with 0.2% fluorine. Perhaps
this 3 to 6 Ma rhyolite is the source of the topaz. Other F-rich
rhyolite domes occur near Magic Reservoir. However, Bennett
(1980) reports that the 44 Ma Dismal Swamp stock (which
intrudes the southern part of the Idaho Batholith) contains topaz
and beryl and is anomalously radioactive. It is not known if this
Mo-mineralized intrusion has a rhyolitic phase, but such plutons
could also be a source of the topaz. Blixt (1933) reports topaz and
fluorite of uncertain paragenesis associated with gold mineralization in the North Moccasin Mountains of central Montana. Van
Alstine (1969,1974) reports the occurrence oftopaz in the upper
part of a devitrified rhyolitic (68 to 73% Si02) ash flow from near
Poncha Springs in central Colorado. The vitrophyre of the ash
flow is chemically dissimilar to topaz rhyolites and contains only
0.08% fluorine. The ash flow appears to correlate with the 36 Ma
Wall Mountain Tuff mentioned earlier.
Mexican topaz rhyolites
The youngest volcanic sequence in the Sierra Madre Occidental of Mexico contains numerous topaz rhyolites (e.g. Foshag
and Fries 1942; Pan 1974; Huspeni et al. 1984; Ludington et al.
1984; Ruiz et al. 1985). The rhyolites commonly contain tin
mineralization like that described from the Black Range in New
Mexico (e.g. Ypma and Simmons 1969; Huspeni et al. 1984;
Duffield et al. 1984).
Most of the occurrences lie in a southeast-trending belt extending from Durango to near Mexico City (Figure 29) that
parallels the main Tertiary volcanic belt of the region. The rhyolites overlie mid-Tertiary andesitic to rhyolitic lavas and tuffs
(Foshag and Fries 1942; Ympa and Simmons 1969; Cameron et
al. 1980). The topaz rhyolites were erupted from 32-27 Ma
during the climax ofthe mid-Tertiary calc-alkaline magmatism of
the Sierra Madre Occidental (Huspeni et al. 1984; Cameron et
al. 1980). In the states of Durango and Zacatecas, the tin-bearing
varieties occur as subvolcanic plugs and as extrusive domes and
lava flows. They intrude other volcanic units near the margins of
calderas. Some are covered by crystal-rich rhyolitic ignimbrites
that may have erupted shortly after the emplacement of the lavas,
leading some authors to suggest that the Sn rhyolites are directly
associated with the caldera-related rhyolites (Huspeni et al.
1982; Ruiz et al. 1985). Our own field observations and those of
others (e.g. Ludington et al. 1984) suggest that this interpretation
requires further study.
Most of the F-rich lavas are phenocryst-poor with phenocrysts of quartz, plagioclase (AnlO-20), sanidine (OrSO-70), and
traces of ferroaugite and Fe- and F-rich biotite (Pan 1974; Huspeni et al. 1984, Ruiz et al. 1985). Crystallization temperatures
are relatively low (650 to 780°C) as determined by two-feldspar
geothermometry (Pan 1974; Huspeni et al. 1984). One notable
difference from topaz rhyolites in the United States may be the
presence of fayalite (Ympa and Simmons 1969; Pan 1974) in
some tin rhyolites (though not necessarily those that bear topaz).
In rhyolites with moderate fluorine contents (0.1 to 0.2%) the
exchange of Al for Fe in the vapor-phase mineralogy stabilizes
garnet over fayalite, and biotite is generally the stable Fe-bearing
magmatic phase.
Huspeni et al. (1984) and Ruiz et al. (1985) report that the
F-rich lavas are geochemically similar to topaz rhyolites from the
western United States. Like their counterparts to the north, the
Mexican tin rhyolites are high-Si02, metaluminous to slightly
peraluminous lavas, with marked depletions ofTi, Mg, Ca, and P
(Table 5). Many of the analyses reported have anomalously high
K20/Na20 ratios (> 1.5) similar to but higher than those found
in tin rhyolites from the Black Range, New Mexico, and the
Sheep Creek Range, Nevada. The role of post-magmatic processes, such as alkali metasomatism during devitrification, needs
to be examined, but "fresh vitrophyres" show similar ratios. Fluorine concentrations in two vitrophyres were 2000 and 3000 ppm;
no CI analyses were reported. Their trace element concentrations
are indistinguishable from topaz rhyolites in the western United
States (Figure 41), with the typically high concentrations of Rb,
Cs, Ta, Th, U, and variable enrichment of Sn (<20 ppm). REE
patterns show deep Eu anomalies and low La/YbN ratios. In
their high Y·(>100 ppm) and total REE content they are most
similar to the Sn-mineralized varieties in the Black Range, New
Mexico, and the Sheep Creek Range, Nevada. Initial 87Sr/86Sr
ratios of the tin rhyolites range from 0.7054 to 0.7075 (Huspeni
et al. 1984; Ruiz et al. 1985) and are slightly higher than associated calc-alkaline (i.e. F-poor) volcanic rocks.
Precambrian topaz rhyolites
Topaz rhyolite magmatism in not strictly a Cenozoic phenomenon in western North America. Topaz (and fluorite) occurs
Topaz Rhyolites
43
o,'-_ _----'--_
1()0
200
_- - ' -_ _.......J
Miles
Figure 29. Distribution of topaz rhyolites in Mexico (compiled from sources cited in text). 1- AmericaSaporis, Durango; 2 - Cerro de los Renidos, Durango; 3 - Fresnillo, Zacatecas; 4 - Pinos, Zacatecas; 5 Guadalcazar, San Luis Potosi; 9 - Hacienda Sauceda, Guanajuato; 10 - San Felipe, Guanajuato; 11 TIachiquera, Guanajuato; 12 - Leon, Guanajuato; 13 - Tepuxtepec, Guanajuato; and 14 - Apulco,
Hidalgo. Small unnumbered dots correspond to rhyolites with Sn mineralization which may also
contain topaz. For the most part the F-rich lavas occur along the eastern margin of the Tertiary
calc-alkaline volcanic belt of the Sierra Madre Occidental. Also shown are the location of the presently
active volcanoes (stars).
in Precambrian rhyolites of the Keewatin District in the
Northwest Territories of Canada (Le Cheminant et al. 1981).
The anomalously radioactive rhyolites are part of a bimodal
magmatic suite related to an early Proterozoic rift. Topaz has also
been found in heavy mineral concentrates from a Precambrian
meta-rhyolite tuff from the Flying W Ranch area, west of Young,
Arizona (Conway 1976). The origin of the topaz is equivocal. It
may be metasomatic and related to younger beryl mineralization
in nearby quartz veins.
PRINCIPAL CHARACTERISTICS OF
TOPAZ RHYOLITES
From the occurrences reviewed above, all Cenozoic topaz
rhyolites from the western United States appear to be remarkably
similar in terms of their mode of emplacement, mineralogy,
chemistry, associated ore deposits, and volcanic-tectonic setting,
despite their wide distribution and diversity of ages. These similarities are reviewed below. The implications of these characteristics are also considered in companion papers (Christiansen et al.
1983a; Burt et al. 1982).
Distribution and ages
Topaz rhyolites are widespread in western North America
and their occurrence closely coincides with the limit oflate Cenozoic extensional faulting (Christiansen et al. 1983a). In the United
States, their emplacement appears to have spanned most of the
Cenozoic Era (Table 6). Their isotopic ages range from 50 Ma
(LIttle Belt Mountains, Montana) to 0.06 Ma (Blackfoot lava
field, Idaho), although all but 3 are younger than 30 Ma. In
Mexico, isotopic ages of topaz rhyolites cluster at 27 to 32 Ma
(Huspeni et al. 1983; Ruiz et al. 1985).
Most known topaz rhyolites in the western United States lie
within the eastern and southern Basin and Range province and
along the Rio Grande rift and thus appear to surround the Colorado Plateau. As far as is known, no topaz rhyolites occur in the
western Great Basin region of California, Nevada, or Oregon, in
spite of this area's contemporaneous bimodal (basalt-rhyolite)
volcanism and extensional faulting. No topaz rhyolites have been
identified to the west of the initial 87Sr/86Sr = 0.706 line as
determined for Mesozoic plutonic rocks (Figure 1; Kistler 1983;
Armstrong et al. 1977) or Cenozoic silicic volcanic rocks (Wilson
44
TABLE 5.
Si02
Ti02
A1203
AVERAGE COMPOSITION OF TIN RHYOLITES
FROM NORTHERN MEXICO.
1
S.D.
76.9
0.60
0.07
13.0
0.01
0.15
2
77.0
0.06
12.8
S.D.
0.12
0.0
0.12
S.D.
3
75.6
0.71
0.14
0.09
12.8
0.38
Fe203*
1.14
0.08
1.14
0.09
1.12
0.19
MnO
0.02
0.00
0.04
0.02
0.06
0.01
MgO
0.13
0.10
0.09
0.06
0.15
0.07
CaO
0.56
0.38
0.30
0.21
0.133
0.37
Na20
2.76
0.10
3.91
0.08
3.73
0.28
0.35
K20
5.~5
0.36
4.67
0.05
5.04
P205
0.01
0.00
0.00
0.01
0.00
0.01
F
0.31
0.23
0.09
0.21
Rb
548
8
807
18
510
42
Zr
73
21
95
6
114
15
U
16
5
17
2
21
5
0.2
10
0.1
Ta
4.7
5.6
0.8
Note: All analyses in weight percent or ppm and
recalculated H20, C02 and S02 free. S.D. is one
standard deviation. Fluorine analyses from vitrophyres.
*Total Fe as Fe203
1. Average of 3 "host-rhyolites" for Sn mineralization at
Sombrete, Zacatecas (Huspeni et al. 1984).
2. Average of 4 "host-rhyolites" for Sn mineralization at
America-Saporis, Durango (Huspen~ et al. 1984).
3. Average of rhyolites from the Thomas Range, Utah
(Christiansen et al. 1984).
et al. 1983). This line is taken by these investigators to mark the Mode ofemplacement
westernmost extent of the Proterozoic craton in the western
United States. Farmer and DePaolo (1983) suggest from their Nd
Nearly all of the topaz rhyolites described in this report were
and Sr isotope studies of granitoids from the region that the emplaced as endogenous lava domes with or without lava flows
sialic continental margin lies farther inland (where eNd = -7 or (Table 7). Several topaz rhyolites were also emplaced as small
87Sr/86Sr = 0.708; Figure 1), but this datum is not well- intrusive domes or plugs that may not have vented to the surface
constrained in central Nevada. These crustal discontinuities, ex- (e.g. Lake City, Colorado, Chalk Mountain, Colorado, Little Belt
pressed structurally as the Roberts Mountain and Golconda Mountains, Montana, and some occurrences in the Wah Wah
thrusts in central Nevada, mark the eastern limit of a series of Mountains, Utah); their fme grain-size, the presence of miarolitic
allochthonous or "suspect" terranes composed of ocean-floor or cavities and glassy margins suggest that even they were emplaced
island arc crust (e.g. Speed 1979). These terranes may have at very shallow levels. The extrusive rhyolites are generally underformed as oceanic crust at the margin of North America during lain by pyroclastic deposits that appear to be remnants of tuff
the Paleozoic and early Mesozoic Eras (Oldow 1984). Given a rings formed by base-surge eruptions (cf. Sheridan and Updike
crustal origin for the parental magmas of topaz rhyolites and the 1975; Wohletz and Sheridan 1983a). The pyroclastic deposits
absence of topaz rhyolites in this region, the young, mafic crust generally consist of a lower, near-vent ("peel-back") breccia that
does not appear to have a composition appropriate for the gener- represents vent-clearing explosions. Breccia fragments may come
ation of topaz rhyolites.
from as deep as 1 km (e.g. East Grants Ridge, New Mexico). The
45
Topaz Rhyolites
TABLE 6. AGES OF CENOZOIC TOPAZ RHYOLITES IN THE
WESTERN UNITED STATES
Location
Age (Mal
Reference
6-7
21
4.7
Lindsey 1981
Lindsey 1981
4. Smelter Knolls, UT
5. Keg Mountains, UT
6. Mineral Mountains, UT
3.4
8
0.5
Lindsey et al. 1975
Lipman et al. 1978b
7. Wah Wah Mountains, UT
20-18
12
22.6
13.4
Lindsey and Osmon son 1978
Best et al. 1985
Barrot 1984
Novak 1984
15
14
16
Wells et al. 1971
Stevlart et al. 1977
Coats et al. 1977
I. Thomas Range, UT
2. Spor Mountain, UT
3. Honeycomb Hills, UT
8. Wilson Creek R, NV
9. Kane Springs Wash, NV
10. Cortez Mountains, NV
II. Sheep Creek Range, NV
12. Jarbidge, NV
13. Blackfoot Lava Field, ID 0.06
36
14. Elkhorn Mountains, NV
50
15. Little Belt Mtns, MT
16. Specimen Mountain, CO
17. Chalk Mountain, CO
18. Nathrop, CO
28-27
28-27
28-29
19. Silver Cli ff-Ros ita, CO 26
38
20. Tomichi Dome, CO
2I. Boston Peak, CO
22. Lake City, CO
23. Grants Ridge, NM
24. Black Range, NM
Corbett 1968
Tweto and Case 1972
Van Alstine 1969
Sharp 1978
F.E. Mutschler unpub.
Ernst 1980
Lipman et al. 1978a
Bassett et al. 1963
Ratte et al. 1984
18.5
3.3
28
25. Saddle Mountain, AZ
L Cenozoic
26. Burro Creek, AZ
breccia is commonly overlain by stratified pyroclastic-surge units
produced during pulsing unsteady eruptions. Some short and thin
(less than I m) lithic-rich ash-flow tuffs probably resulted from
minor collapse of low eruption columns. Mantling ash-fall unitS
punctuate the record of explosive volcanism. These features suggest that the pyroclastic eruptions were initiated as rising magmas
explosively mixed with groundwater (hydromagmatic eruptions;
Wohletz and Sheridan 1983b). However, the origin of the driving volatiles (magmatic versus phreatic) is difficult to establish
without detailed studies of individual complexes (cf. Taylor et al.
1983). Once the vent was cleared, relatively quiet eruptions of
rhyolite lava proceeded. Lava eruptions may have been caused by
the eruptive degassing of the magma and the evisceration of a
volatile-rich cap of a small magma chamber, or by the restricted
access of ground water to the vent.
The geology of the Mineral Mountains domes in Utah is
instructive in this regard in that the F-rich domes have basal
tephra deposits while earlier F-poor lavas lack them. This observation suggests that some explosive volcanism resulted from
magmatic differentiation and volatile enrichment (Evans and
Nash 1978) of the upper part of a rhyolitic magma chamber. The
viscous domes may have been extruded following devolatilization
of a chamber's cap.
Pierce et al. 1982
Chadwick 1978
Witkind 1973
Anthony et al. 1977
Burt et al. 1981b
Lavas are generally underlain by a flow breccia (about I m
thick) produced as the flow-front crumbled, slumped, and was
overridden by the flow in caterpillar fashion. Rapidly quenched
vitrophyric blocks from the flow front are common in this layer.
The volume of magma in individual domes or flows ranges
from less than 0.5 km 3 to a probable maximum of about 10 km 3.
However, in some cases, fairly large volumes (10 to 100 km 3) of
coalesced domes and flows accumulated over short intervals
(about 1 Ma); for example, the Thomas Range, Utah, the Wah
Wah Mountains, Utah, and the Black Range, New Mexico.
These observations set topaz rhyolite eruptions apart from
the large ash-flow eruptions of high-silica rhyolite that culminate
in caldera collapse. Volumes of ash flows from calderas are
commonly 1 to 2 orders of magnitude larger (Smith 1979) than
those from topaz rhyolites. However, eruptions of large volumes
of magma over geologically short time intervals in the Thomas
Range, Utah, and in the Black Range, New Mexico, suggest that
some topaz rhyolites may emanate from magma chambers with
volumes approaching those of caldera-related plutons. In fact,
some topaz rhyolites may have arisen from magma chambers that
gave rise earlier to large ash-flow sheets (e.g., the Black Range,
New Mexico, and at Kane Springs Wash, Nevada). Some of
these differences are shown schematically in Figure 4K
46
TABLE 7. MODE OF EMPLACEMENT OF TOPAZ RHYOLITES
Location
Area
(km 2 )
Thomas Range, UT
160
Emplacement
Coalesced domes and lava flows with
underlying tuffs.
Lava with underlying tuff and isolated
dome(?).
Lava domes with underlying tuff.
Spor Mountain, UT
5
Honeycomb Hills, UT
1
Smelter Knolls, UT
10
Isolated lava dome with no tuff exposed.
Keg Mountain, UT
30
Partially coalesced domes and lava flows
with underlying tuffs.
~1ultiple isolated domes with underlying
tuffs.
Coalesced domes and lava flows with
underlying tuff1 some isolated plugs and
domes (many <lkm 2 ).
Lava flows with underlying lithic tuffs and
a prominent welded ash-flow sheet.
Isolated intra-caldera dome with underlying
tuff.
Isolated domes or plugs with no tuff
exposed.
Coalesced dome/flow complexes with little
associated tuff.
Coalesced dome/flow complexes (?) with
little tuff.
5 isolated domes with underlying tuffs.
Mineral Range, UT
8
Wah Wah vicinity, UT
75
Wilson Creek R,
20
NV
Kane Springs Wash,
4
NV
Cortez Mountains,
NV
2
Sheep Creek Mtns,
NV
50
Jarbidge,
?
NV
Blackfoot Lava Field, In
4
Elkhorn Mountains, MT
5
Little Belt Mtns, MT
5
Specimen Mountain, CO
8
Chalk Mountain, CO
4
Nathrop, CO
3
Silver Cliff, CO
12
Tomichi Dome, CO
5
Boston Peak, CO
1
Isolated plugs and lava domes, some with
tuff.
Isolated plug (bysmalith) and sills.
Isolated lava flow/plug with underlying
tuff.
Isolated intrusive stock or extrusive plug.
3 isolated domes with lava flow and
underlying tuff.
Coalescing flows overlying thick tuff in
small subsidence structure.
Isolated lava dome with underlying tuff;
sill.
Lake City, CO
Grants Ridge, NM
Black Range, NM
10
6
100
Saddle Mountain, AZ
?
Burro Creek, AZ
4
Isolated flow-banded plugs and breccia
pipe1 no tuff exposed.
Isolated intrusive plugs, dikes and sills
near margin of older unrelated caldera.
Perlite and lava dome intruding tuff, short
lava flows.
Coalesced lava dome complexes with
associated tuff.
?
MUltiple isolated dome/flow complexes each
with an underlying tuff.
Mineralogy
The mineralogy of topaz rhyolites is relatively simple and is
summarized in Tables 8 and 9. Phenocryst-poor (less than 5%)
rocks are the most common, but in some lavas and shallow
intrusions the phenocryst content may be as high as 40%. In order
of abundance, sanidine, quartz, and sodic plagioclase are the
principal phenocrysts. Biotite is common; hornblende, garnet, and
clinopyroxene occur in a few samples. Common magmatic accessories include zircon, apatite, magnetite, ilmenite, allanite, fluorite, and titanite.
Fe-Ti oxides and titanite. Titaniferous magnetite and
ilmenite both occur in topaz rhyolites. These phases may be
altered in devitrified lavas but are commonly unoxidized in vitrophyres. The few Fe-Ti oxide analyses that exist indicate that the
ilmenites are generally Mn-rich. Table 4 contains the results of FeTi oxide geothermometry for topaz rhyolites and shows that most
of these rhyolites crystallized at temperatures between 600 and
850°C; mostly in the lower end of that range. Moreover, these
analyses indicate that f0 2 is commonly low,(Figure 30), near the
QFM oxygen buffer (e.g. the rhyolites of the Thomas Range,
Spor Mountain, and Smelter Knolls, Utah). However, analyses of
Topaz Rhyolites
47
TABLE 8. MAGMATIC MINERALS REPORTED IN TOPAZ RHYOLITES
Location
Sa
Qz
PI
Bt
Mt
Thomas Range., UT
Spor Mountain, UT
Honeycomb Hills, UT
x
x
X
X
X
X
X
X
+
x
x
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
x
Smelter Knolls, UT
Keg Mountain, UT
x
X
X
X
X
X
Wilson Creek Range, NV
Kane Springs Wash, NV
X
X
X
X
X
+
X
X
X
XXXX
Cortez Mountains, ·NV
Sheep Creek Mountains, NV
Jarbidge, NV
X
X
X
X
X
X
X
X
X
Blackfoot lava field, ID
Elkhorn Mountains,MT
Little Belt Mountains, MT
X
X
X
X
X
X
Chalk Mountain, CO
Nathrop, CO
X
X
X
X
Silver Cliff, CO
Tomichi Dome, CO
Boston Peak, CO
Lake City, CO
Grants Ridge, NM
Black Range, NM
Saddle Mountain, AZ
Burro Creek, AZ
Op
x
II
Ho
Px
X
±
±
Fa
Gt
Zr
Al
Ap
Tt
±
X
X
X
X
X
±
Fl
Tz
Th
+
X
X
X
±
X
X
X
X
+
X
±
X
X
X
X
±
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
+
X
X
X
X
X
±
X
X
X
X
±
X
X
X
±
X
X
X
X
±
X
X
X
X
X
X
X
X
X
X
?
X
±
X
X
X
X
X
X
X
xjV
X
±
±
?
X
X
X
X
X
X
X
X
±
X-present in most samples; ±-present in some samples; ?-questionable or uncertain report.
may be vapor phase.
XjV-
Sa-sandine; Qz-quartz; PI-plagioclase; Bt-biotite; Mt-magnetite; Op-unidentified opaque
mineral; II-ilmenite; Hb-hornblende; Px-pyroxene; Fa-fayalite; Gt-garnet; Zr-zircon; AI-allanite;
Ap-apatite; Tt-titanite; Fl-flourite; Th-thorite; and Tz-topaz.
oxides from the Mineral Range, Utah (ca. NNO), and Chalk
Mountain, Colorado (3 log units above NNO), demonstrate that
some topaz rhyolites crystallized under relatively oxidizing
conditions.
In the latter case, these values approach those inferred by
Keith (1982) to have existed in the rhyolitic magma chamber that
gave rise to the tuff of Pine Grove and to a related Climax-type
molybdenite deposit in the Wah Wah Mountains of southwestern
Utah. It is important to note that of all the topaz rhyolites described in this report, the Chalk Mountain rhyolite is most
intimately related to such a mineral deposit-the Climax deposit
itself. Perhaps this bears out the suggestion of Keith (1982) that
high oxygen fugacities are important for the generation of
Climax-type Mo deposits.
In the absence of detailed studies of Fe-Ti oxides, the presence of titanite in several topaz rhyolites may be a mineralogic
indicator of relatively high f0 2 (Haggerty 1976). Titanite occurs
in both the Chalk Mountain, Colorado, and Mineral Range,
Utah, rhyolites, where independent evidence suggests oxidizing
conditions. Titanite also occurs in the Lake City, Colorado, rhyo-
lites that have prominent middle REE depletions (R. A. Zielinski,
written communication, 1982), which probably indicate titanite
fractionation. Titanite is also reported from the Sheep Creek
Mountains and Jarbidge, Nevada, where it is not known if the
rhyolites possess middle REE depletions. It should be noted,
however, that not all topaz rhyolites that bear titanite have middle REE depletions. For example, some samples of the Taylor
Creek Rhyolite, New Mexico, and a few of the lavas from the
Thomas Range, Utah contain sparse titanite but, as far as is
known, none of the lavas possess middle REE depletions. It thus
appears that there are less and more oxidized topaz rhyolites, a
situation perhaps analogous to that described for the Proterozoic
anorogenic granites described by Anderson (1983) to consist of
an ilmenite- and a magnetite-series.
Feldspar. Almost all topaz rhyolites are two-feldspar rhyolites, in contrast to many other bimodal rhyolites-for example,
many of the rhyolites of the Snake River Plain, Idaho (Leeman
1982a; Hildreth 1981) and the peralkaline rhyolites of the western Great Basin (e.g., Rytuba and McKee 1984; Conrad 1984;
Novak 1984; NOble and Parker 1974). In general, one-feldspar
48
TABLE 9. DEVITRIFICATION AND VAPOR-PHASE MINERALOGY REPORTED IN TOPAZ RHYOLITES
Location
Sa
Qz
Thomas Range, UT
Spor Mountain, UT
Honeycomb Hills, UT
X
X
X
X
X
X
Smelter Knolls, UT
Keg Mountain, UT
X
X
X
X
X
X
Wilson creek Range, NV
Cortez Mountains, NV
X
X
X
X
Jarbidge, NV
X
X
Chalk Mountain, CO
Nathrop, CO
PI
Bt
Mt
Gt
FI
Tz
Bx
Ps
Hm
Be
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
?
X
X
X
X
X
Tm
X
X
X
X
X
X
X
X
X
X
X
X
X
Ct
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
Silver Cliff, CO
Tomichi Dome, CO
Boston Peak, CO
X
X
X
X
Lake City, CO
Grants Ridge, NM
Black Range, NM
X
X
X
X
X
Saddle Mountain, AZ
Burro Creek, AZ
X
X
X
X
X
?
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X/M
X
X
X
X
X
X
X
X
X
X
x-present in some samples; ?-uncertain identification; X/M-may be magmatic.
Bx-bixbyite, Ps-pseudobrookite, Hm-hematite, Be-beryl, Ct-cassiterite,
Tm-tourmaline; others as in Table 8.
rhyolites crystallize at higher temperatures than those inferred for
F-rich topaz rhyolites. Sanidine in topaz rhyolites is generally
Or40 toOr60. Plagioclase is generally sodic oligoclase, although.
compositions as sodie as calcic albite are found in evolved topaz
rhyolites like the Spor Mountain rhyolite, Utah. Andesine is
found in less evolved rhyolites like Chalk Mountain, Colorado.
Two-feldspar temperatures (Stormer 1975) are shown in Table 4.
Two.-feldspar (calculated at 100 b to 1 kb) and Fe-Ti oxide
temperatures are generally in good agreement where they have
been analyzed from the same sample. The rhyolites of the Thomas. Range, Utah, show the broadest temperature range (790 to
600°C); temperature ill negatively correlated with F and other
incompatible element concentrations. All equilibration temperatures, as determined by feldspar pairs from other localities, fall
within the lower part of this range.
Mafic silicates. Biotites from topaz rhyolites generally
have high Fe/(Fe+Mg)ratios (Figure 31) reflecting the high
Fe/(Fe+Mg) of the magma (in many cases, molar Mn and Ti
exceed Mg), and perhaps the prevalence of relatively low f0 2 in
these types of magmas. Nonetheless, variable Fe/Mg ratios have
resulted from the variaple oxidation states inferred above produc-
ing more magnesian biotites in some lavas. In general, the Altot is
lessthan 3 moles per 24 (0, OH, F, Cl). In contrast, biotites from
two-mica granites of the Basin and Range province and strongly
peraluminous, S-type, granites the world· over generally contain
more Al than this, a fact indicative ofthe metasedimentary parentage of.S-type granites (Figure 31). The concentrations of F and
Cl have been analyzed in relatively few biotites from topaz rhyolites. Existing analyses demonstrate that the biotites have high
F-contents (up to 5 wt%), Concentrations this high for Fe-rich
biotites suggest crystallization at high fHF and at high fHF/fH20
(10- 1 to 10-3 for the Thomas Range Rhyolites; Turley and Nash
1980). F/CI ratios in the biotites also suggest crys~llization at
high fHF/fHGI. On a molar plot of Mg/(Mg+Fe) vs log F/CI
(Figure 32), biotites from topaz rhyolites fall in the same compositional fields as the ilmenite series granites from Japan (Czamanske et al. 1981) and greisenized Sn-mineralized granites
(Figure 32). Gunow et al. (1980) and Munoz (1984) have shown
that such high F/CI ratios are also characteristic ofthe Henderson
molybdenite deposit. Brimhall et al. (1983) and Munoz (1984)
have generalized this observation to many other Mo, Sn, W, and
Be deposits and contrast these deposits with piotites from Cu-rich
Topaz Rhyolites
-6
-
49
~~".,
-8
-10
,-
,-
,-
....
./
C'iI
./
0
0)
0
:0... :.
,-
••
./
-12
-
./
.'
.
Bishop Tuff
/
-14
./
.. ...
-16
o
-18
Topa2: rhyolites
west-central Utah
-20
-22
600
700
800
900
1000
1100
Temperature ( °C)
Figure 30. Compilation of T-log f02 data for silicic andesites, dacites and rhyolites from western North
America (after Ewart 1979, except as noted). The oxygen buffer curves (HM hematite-magnetite; NNO
nickel-nickel oxide; QFM quartz-fayalite-magnetite) are for one atmosphere pressure. Large dots are
Californian bimodal rhyolites; smalldots are for calc-alkaline intermediate to silicic rocks; and the open
field is for the Bishop Tuff, California (Hlldreth 1977). The topaz rhyolite field includes rhyolites from
the Thomas Range, Smelter Knolls, and Spor Mountain; Utah; large filled circles are for Chalk
Mountain, Colorado, and open circles for the Mineral Range, Utah (all from sources cited in text).
porphyry deposits which have low FICI ratios when corrected
for Fe-F avoidance (Figure 32).
Clinopyroxene (in high T andF~poor rhyolite; e.g. "mafic"
Thomas Range, Utah, and Jarbidge, Nevada, specimens), fayalite
(Kane Springs Wash, Nevada, and some Mexican tin rhyolites),
Fe-rich hornblende (Figure 33), or Fe-Mn garnet are found in a
few specimens from topaz-bearing rhyolites. Fe-enriched mafic
minerals are common in bimodal rhyolite magmas (Ewart 1979)
and in anorogenic or A-type granites (e.g., Anderson 1980,
1983). As for the case of biotite, this association reflects the high
Fe/Fe+Mg ratios of the magmas and, possibly, the prevalence of
relatively low f0 2 in these types of magmas.
The vapor-phase mineralogy of topaz rhyolites (Table 9) is
diverse and includes topaz, fluorite, spessartine gamet, beryl, FeTi-Mn oxides (pseudobrookite, hematite, and bixbyite), silica
minerals, and alkali feldspar. The compositions of these minerals
and their implications for the transport of elements in a vapor
phase remain little studied. However,fluid inclusion and oxygen
isotope studies of topaz and quartz in rhyolites from the Black
Range indicate that vapor-phase crystallization-. occurred
at a temperature in excess of 600°C (Eggleston and Norman
1984). High salinity in the fluid probably reflects fluid boiling. In
addition, the high iron content of bixbyite and the presence of
pseudobrookite instead of rutile indicate initial' temperatures of
over 500°C during crystallization in some cavities (Lufkin 1976).
Burt (1981) has shown that the vapor-phase mineral assemblage
is consistent with moderately high fHF and relatively high fo 2•
Barton's (1982) exposition of the thermodynamic properties of
topaz suggest that the HF/H20fugacity ratio exceeded 10-3 at
temperatures above 600°C, stabilizing topaz and K-feldspar
over muscovite plus quartz. The latter assemblage is more typical
ofgreisens formed at some depth. The general absence of fayalite
50
Annite
3r-------------...,..~----___,
Siderophyllite
0.8 f-
2
H
()
....
u.
o Chalk MIn
• Spor MIn
K Smelter Knolls
H Honeycomb Hills
o Thomas Range
OJ
o
...J
1
2.0
Phlogopite
3.0
4.0
AI atoms/24 (O,OH,F)
Eastonite
Figure 31. Compositions of unoxidized biotite from topaz rhyolites from
the western United States in terms of molar Fe/(Fe + Mg) and total Al
relative to ideal end members. Sources of data for topaz rhyolites are
given in the text. For comparison the compositions of biotites from the
Bishop Tuff, California (RT - Hildreth 1977), from calc-alkaline igneous
rocks from the western United States (Wender and Nash 1977; Hausel
and Nash 1979; Dodge et aI. 1969) and from muscovite-bearing granitoids of the western United States (MG-Best et aI. 1974; Lee et aI.
1981; Kistler et aI. 1981; Dodge et aI. 1969) are also given.
Hastingsite
Pargasite
1.0
u:;
:i
q,
9
o:l'
C'\l
0.8
0.6
.....III
e
....III0
0
;
0.4
... - ....,
/
S.
<E
Wolf
I
River
,Massif
I
0.2
\
'-; "
0.0
0.2
0.4
0.6
Edenite
0.8
H
\
\
\
}
,-
//
1.0
Ferroedenite
Fe/Fe+ Mg
Figure 33. Composition of hornblende in terms of molar Fe/(Fe + Mg)
and octahedra1 Al relativeto ideal end members. The filled circle represents hornblendes from one lava flow in the Thomas Range. Amphiboles
from other anorogenic silicic magmas have similar Fe/(Fe + Mg) ratios
(Wolf River Batholith, Wisconsin: Anderson 1980; and the Pikes Peak
batholith, Colorado (+): Barker et aI. 1975). Hornblendes from calc~
alkaline volcanic rocks of the eastern Great Basin are shown as open
circles (Hausel and Nash 1979; Wender and Nash 1977; Keith 1982).
The solid line encloses the composi~ions of hornblendes from the Sierra
Nevada batholith, California (Dodge et aI. 1969), and the southwest
Japanese batholith (Czamanske et aI. 1981).
o
-1
L-_.l-_.l-_..I-_...l-_...J....._...J....._...J....._...J....._-'-_-'
0.1
0.3
0.5
0.7
0.9
Mg/Mg+Fe
Figure 32. Compositions of unoxidized biotite in topaz rhyolites from the
western United States in terms of molar Mg/(Fe + Mg) and log F/Cl
(molar). Symbols are the same as in Figure 31. Sources of the data for
topaz rhyolites are given in the text. Other fields shown for comparison
are from Keith (1982), Gunow et aI. (1980), Czamanske et aI. (1981),
Jacobs and Parry (1979), and Parry et aI. (1978). Isopleths of equal F
enrichment have a positive slope on this diagram. Topaz rhyolites have
biotites which are consistently F-rich when "corrected" for their usually
high Fe/Mg ratios. In this regard, they are most like biotites from
ilmenite-series granites of Japan (Czamanske et aI. 1981) and from
granites associated with Mo-Sn-W-Be deposits (Munoz 1984).
(as a phenocryst or as a product of vapor-phase crystallization) is
notable; it is presumably unstable with respect to F-bearing biotite. Peralkaline minerals (aegerine, riebeckite, etc.) are likewise
absent. In contrast, the presence of aluminous minerals in miarolitic cavities, especially topaz and spessartine garnet (Fig. 34),
some of which is F-bearing (Moyer, 1982), suggests a link between F and Al in an escaping vapor-phase after eruption. This
comigration may be an extension of their association in aluminosilicate melts proposed by Manning et al. (1980) and Christiansen
et al. (1983a).
Additional detailed studies of magmatic and vapor-phase
mineral compositions could help to further constrain the values of
the intensive parameters that prevailed during cooling and crystallization of topaz rhyolites. Mineral chemistry could also help to
document the role of fluorine in silicate melts and minerals. Studies of phenocryst/melt partitioning are as yet lacking for topaz
rhyolites from the United States. Such studies would be useful for
sorting out the role of volatiles and melt structure in determining
partition coefficients in highly silicic magmas (e.g., Mahood and
Hildreth 1983).
Geochemistry and differentiation trends
The major element composition range of topaz rhyolites is
fairly restricted. All are high-silica rhyolites with high Na, K, F,
Topaz Rhyolites
51
Spessartine
x Burro Creek, Arizono
o
Nathrop, Colorado
• . Thomas Range, Utah
t::.
Pine Grove, Utah
o
Ely, Nevada
A
Canterbury, New Zealand
Kern Plateau, California
---- Pegmatitic garnets
+
\
\
\
\
\
\
\
\
\
\
\
\
\XX
I
.
" \ X XX
\
\
X
\
\
X
"
\
X
\
\
\
\
\
\
Pyrope
L--
---"'--'''''''''=<.....................
Almandine
Figure 34, Compositions of garnet from rhyolitic volcanic rocks in terms of Fe, Mn, and Mg endmembers. Garnets from the rhyolitic tuff of Pine Grove, Utah (Keith 1980), and lavas from Canterbury,
New Zealand (Wood 1974), and the Kern Plateau, California (Bacon and Duffield 1981) are magmatic;
the rest are Mn-rich vapor-phase garnets (Miyashiro 1955; Cross 1886; Christiansen et al. 1980; Moyer
1982; Pabst 1938). The pegmatitic garnet field is from Miyashiro (1955).
and Fe/Mg and low Ti, Mg, Ca, and P (Table 1; Figure 35).
These characteristics are typical of bimodal rhyolites (see below),
Topaz rhyolites are apparently a subclass of this group (a conclusion also justified by their trace element chemistry, Fe-enriched
mineralogy, tectonic setting, and magma associations). The composition of a "typical" topaz rhyolite is shown in Table 1 and
represents modal values from the histograms in Figure 35,
All topaz rhyolites thus far identified from the western United Stateshave Si02 concentrations greater than 72% (all Si02
concentrations have been recalculated on an anhydrous basis)
and a strong mode exists at 76% (Figure 35). Silica contents vary
little with differentiation trends seen in individual dome complexes. For example, silica ranges from 74.2% to 76.7% in the
rhyolites of the Thomas Range, Utah, whereas incompatible elements such as Rb triple in concentration, As a result silica concentrations are poor indicators of the chemical variability of these
rhyolitic magmas. In addition, silica contents are actually lower in
rhyolites that are extremely enriched iIi fluoriIle and incompatible
elements such as the lava at Spor Mountain, Utah. The average
silica concentration of these samples is about 74%. Other highly
evolved topaz rhyolites, such as the one at the Honeycomb Hills,
Utah, and the ongonites discussed below, also show silica contents lower than 76%, Christiansen et al. (1984) interpreted the
low silica content as the result of crystallization near the minimum in the granite system, with elevated fluorine. Manning's
(1981) experiments show that the stability field for quartz expands with increasing fluorine content in a water-saturated haplogranite at 1 kb pressure (Figure 36). It is thus conceivable that
enhanced quartz fractionation could lead to a reversal of normal
Si02-enrichment during fractional crystallization of a fluorinerich rhyolite.
Most topaz rhyolites contain 12% to 14% Al203 (Figure 35).
High aluminum contents are found in the evolved low-silica rhyolites noted previously, reflecting the increased proportion of
feldspar components, especially albite. All topaz rhyolites have
high concentrations of alkalies ranging between 8% and 10%
Na20 plus K20. In general K20/Na20 ratios are greater than
one (typically about 1.2 to 1.4 by weight). Several topaz rhyolite
occurrences seem to be typified by K20/Na20 ratios higher than
1.5-the topaz (and topaz-free) rhyolites of northern Nevada, the
Taylor Creek Rhyolite in the Black Range of New Mexico and
the Sn and F-rich rhyolites of Mexico's Sierra Madre Occidental.
Because only a few vitrophyres have been analyzed from these
areas, it remains to be demonstrated that these high ratios are
magmatic and not the result of alkali-metasomatism during subaerial crystallization, In d!fferentiation sequences, K generally declines with advancing concentrations of F and other decidedly
incompatible elements, while Na increases-a result of the combined fractionation of potassic sanidine and biotite (Figure 36). In
spite of high alkali concentrations, topaz rhyolites are not peralka-
52
70
60
50
51°2
40
30
20
10
72 74 76 78 80
0
0.08 0.16
0.24
10 12 14 16
0.6
1.0
1.4
1.8
0.02 0.06 0.10 0.14
0.18
40
K2 0
MgO
30
CaO
P205
Na 2 0
F
20
10
0.04
0.20
0.36
0.20
0.60
1.0
1.4
3.0
3.8
4.6
4.0
4.6
5.4
0
0.04
0.08
~
0.2
0.6
1.0
n
I
1.4
Figure 35. Histograms of whole-rock chemical analyses of topaz rhyolites from the western United
States (values in wt%). Analyses of 118 samples from 22 locations are represented (sources cited in the
text). Analyses were rejected if H20 was greater than 3 wt% or if K20/Na20 (by weight) exceeded 2.
All analyses were recalculated to 100% on an H20- and C02-free basis. The vertical scale shows the
frequency of each value.
line, and the use of the term "alkali rhyolite," reserved by lUGS
usage for peralkaline rhyolites (Streckeisen 1979), should be
abandoned. Use of this term may lead to confusion with the truly
peralkaline rhyolites with which topaz rhyolites are contemporaneous in the western United States. In the lUGS system topaz
rhyolites are generally rhyolites or alkali feldspar rhyolites. Indeed, many vitrophyres from topaz rhyolites are metaluminous.
The presence ofgarnet and topaz (absent as vapor-phase ttlinerals
in peralkaline volcanic rocks) also reflects the aluttlinous character of topaz rhyolites. Many felsites are slightly peraluminous as
indicated by normative corundum-probably as a result of the
loss of alkalies during crystallization.
We prefer to calculate CIPW norms based on fluorine-free
analyses. The CIPW scheme ties up Ca with F to form fluorite
before Ca is used to form plagioclase, thereby producing a rock
that appears to be more strongly peraluminous, as indicated by
normative corundum, than it would be otherwise. This crystallization order is not that observed in topaz rhyolites. Fluorite is one
of the last phases to crystallize in some extremely F-rich glasses
and more generally fluorite is post-magmatic. In short, topaz
rhyolites are not strongly peraluminous indicating that they are
not the eruptive equivalents of S-type grllnites (e.g., White and
Chappel 1983).
The concentrations of CaO (generally <0.8%), Fez03*
(generally <1%), MgO (generally <0.20%), TiOz (generally
<0.12%), and PzOs «0.02%) in topaz rhyolites are relatively low
and similar to those found in other high-silica rhyolites (Figure
35). In cogenetic suites, all of these elements correlate negatively
with incompatible elements and with fluorine. Their decreasing
concentrations can be shown to relate to the fractionation of
modal proportions of plagioclase, biotite, Fe-Ti oxides and apatite. A few samples from the Honeycomb Hills and Spor Moun-
Topaz Rhyolites
53
Q
.
L6
•
••
LOCATION OF PLOT AREA
D.
0.5 D.
1 D. ~.
D.
2 M.
&. D.D.
3~~:
•
5
*D.
•
1%
D.
D.
D.
D.
*
2%
D.
Ab '--_---'>l._---"----"'-----''------'>l.--....:L.--'''--_--''"--_--''-_ _...l. Or
Figure 36. Normative composition of topaz rhyolites from the western United States in terms of quartz
(Q), albite (Ab), and orthoclase (Or) compared to experimentally determined ternary minima in: 1) the
hydrous granite system at pressures given in kb (PHzO) next to the filled circles (Tuttle and Bowen
1958; Luth et aI. 1964; and Whitney 1975); and 2) the F-bearing granite system at 1 kb (Manning
1981). The numbers next to the stars indicate weight % F in the water-saturated system at 1 kb.
tain, Utah and perhaps elsewhere, show Ca-enrichment as the
result of the accumulation of post-magmatic fluorite. Topaz rhyolites have high Fe/Mg ratios and plot in the "tholeiitic" field on
Si02 versus FeO*/(FeO* + MgO) diagrams such as the one used
by Anderson (1983) to discriminate calc-alkaline from tholeiitic
or reduced anorogenic suites. In addition, Fe/Mg ratios increase
with differentiation, driven principally by biotite fractionation as
FeO*IMgObiotite is substantially less than FeO* IMgOmelt .
Manganese concentrations are also low, less than 0.08%
MnO in almost all samples, but in general Mn appears to behave
as an incompatible element. MnO increases with differentiation
from 0.05% to 0.08% in the rhyolites of the Thomas Range.
Similar increases of MnO with differentiation of rhyolitic magmas have been noted by many authors (see review in Hildreth
1981) and appear to relate to limited biotite fractionation.
Of course, the most discriminating feature of topaz rhyolites
as a group is their high fluorine content. Topaz appears as an
identifiable vapor-phase mineral in lavas whose vitrophyres contain over 0.2% F. Fluorine concentrations of more than 1% are
only known from vitrophyres from Spor Mountain and the
Honeycomb Hills, both in western Utah. Figure 35 shows
concentrations in felsites and vitrophyres, where most of the low
values «0.2%) are from felsites. Comparisons of vitrophyrefelsite pairs almost universally show that F is lost during devitrification. Chlorine is also Ibst by devitrification, as no mineral phase
concentrates CI in contrast to F. Meaningful chlorine concentrations can only be obtained by analysis of vitrophyres or obsidians.
Figure 37 shows the fluorine and chlorine concentrations in
glassy specimens of topaz rhyolites. In specimens thus far analyzed, Cl concentrations remain less than 0.2% and are generally
much lower than this. Important features of topaz rhyolites are
their high F/Cl ratios (greater than about 3) as compared to Fand Cl-rich peralkaline rhyolites, which have lower F/Cl.
Topaz rhyolites are variably enriched in incompatible trace
elements (Li, Rb, Cs, Be, U, Th, Y, Nb, Ta, Ga, Pb, Mo, Sn, W,
and HREE) and depleted in feldspar-compatible trace elements
(Ba, Eu, and Sr). Zirconium and Hf concentrations are generally
low as well, consistent with their compatibility with fractionating
zircon. As for other metaluminous rhyolitic magmas (e.g. Hildreth 1979), Zr/Hf ratios (typically 25 to 20) decline with differentiation as the result of the greater compatibility of Zr in zircon.
The concentrations of many elements that are taken up by mafic
silicates (Ni, Co, Cr, and V) are extremely low, but as they lie
near the detection limits for common analytical metho~ their
concentrations are not well known. Typical concentrations are
listed in Table 2 and illustrated in Figure 38. Using trace element
compositions, Pearce et al. (1984) have attempted to interpret the
tectonic settings of granitic rocks. In this classification (Figure
39), topaz rhyolites from the western United States consistently
straddle the boundary between WPG (within-plate granites, like
54
0.8
eft
~
:;=
0.6
w
Z
.,.
I--l
0::::
a
--l
I
U
0.4
..
. .. ..
0.2
<>
0.6
1. 0
1.2
(wt.%)
Figure 37. Fluorine and chlorine concentrations in glassy topaz rhyolites (open diamonds) and peralkaline rhyolites (closed triangles). A FICl ratio of 1 divides oceanic from continental peralkaline rhyolites
(Bailey 1980). A F/C! ratio of 3 separates peralkaline rhyolites from topaz rhyolites of the western
United States. The composition ofthe Bishop Tuff, California (B: Hildreth 1979) is shown for comparison. Data are from Christiansen et al. (1980), Moyer (1982), Dayvault .et al., (1984), Novak (1984),
Turley and Nash (1980); Conrad (1984); Mahood (1981); Bailey (1980); Macdonald and Bailey
(1973).
.
0.2
0.4
the Nigerian Younger granites) and syn-COLG (syn-collisional
granites, like the leuco-granites of the Himalayas). Although these
diagrams point to a certain uniqueness among topaz rhyolites as
compared to many other granites, it is unlikely that simple discriminant diagrams such as these will provide a unique definition
of the type or tectonic setting of granites because of the wide
variety of crustal and mantle components that are involved in
granite genesis. Note for example the location of the metaluminous Bishop Tuff, California, and the comenditic Tala Tuff,
Mexico.
REE patterns (Figure 40) show some variability (perhaps
inherited from slightly different source rocks and/or differentiation histories) but they generally display low La/CeN, La/YbN (1
to 3 for most but this ratio may be as high as 12 for rhyolites from
Colorado), and Eu/Eu* (0.45 to 0.01 for analyzed specimens).
Light REE concentrations generally do not exceed 200 times
chondrite values and more typically are near or less than 100
times chondrite. Differentiation trends for the REE are seen in
several complexes (Thomas Range, Utah, Mineral Mountains,
Utah,Wah Wah Mountains, Utah, Lake City, Colorado) and
show decreases in LREE and Eu concentrations that are coupled
to increasing HREE and other incompatible element concentrations. Christiansen et al. (1984) have modelled this trend for the
0.8
FLUORINE
Thomas Range rhyolites as resulting from the fractionation of
small amounts of allanite (0.04 wt% of the fractionated mineral
assemblage). The middle REE depletions noted for the Mineral
Mountains and the Lake City rhyolites suggest the fractionation
of titanite.
Devitrification mobilizes a number of trace elements (U, Sb,
F, CI, Li, Be, perhaps Sn, W, and Mo)judging from lower
concentrations in vitrophyre-felsite pairs. However, it has not
been demonstrated that devitrification, without significant vaporphase alteration, significantly changes concentrations of many
other trace elements, including petrogenetically important elements like Rb, Sr, Ba, Y, Zr, Hf, Nb, Ta, Th, REE, Sc, and Ga. In
fact, for some trace~element and isotope studies, dense felsites may
be better samples than variably hydrated vitrophyres, as hydration may involve significant addition of Sr, Ca, and other elements (Hargrove 1982), as well as 0 and H isotope exchange.
Some of the compositional features of topaz rhyolites are
summarized in Figure 41, which compares the chemical composition (Table 2) of a variety of topaz rhyolites on normalized
geochemical diagrams. Such normalized-concentration diagrams
allow the relative concentrations of many elements in a single
sample to be displayed and elemental fractionation is easily portrayed. All concentrations have been normalized to those of the
Topaz Rhyolites
Rb
55
y
Zr
Ba
Sr
.~
0
400
800
0
Li
o
80
20
100
0
Be
160
o
40
200
0
80
Nb
80
o
40
80
160
20
U
120
o
20
100
180
Th
40
20
40
60
Figure 38. Histograms of trace element concentrations in topaz rhyolites from the western United States
(all values in ppm). Analyses of 30 to 90 samples are represented for each element and come from 19
different localities (references cited in Table 2). The vertical scale represents the relative frequency of
each value.
U.S. Geological Survey geochemical reference sample RGM-l (a ofY and Sm relative to geochemically similar elements. Normalrhyolite from Glass Mountain, California) as presented in Govin- ized major element diagrams show the depletions of Fe, Mg, Ti,
daraju (1984). RGM-1 was chosen because the concentrations of and Ca typical of most high-silica biotite rhyolites with two feldmany elements have been accurately determined and because it is spars, including topaz rhyolites.
a sub-alkaline high-silica rhyolite grossly similar to many topaz
Taken together these compositional features suggest that
rhyolites. RGM-1 is an obsidian and the concentrations of ele- topaz rhyolites are highly differentiated or evolved magmas.
ments otherwise mobilized by devitrification and hydration Their common association with less evolved, topaz-free rhyolites
should be nearly magmatic. This is particularly important for F, suggests they may have fractionated from more "mafic" composiCl, and U-elements of interest here. RGM-l is reported to have tions. Such evolutionary relationships are discernible in at least
a Th/U ratio of2.6 (1515.8) and a FICI of 0.6 (3401540); values 15 of the 26 localities described here. Christiansen et al. (1984)
that may be typical of calc-alkaline rhyolitic glasses. The elements have used the extreme depletion of Eu and other compatible
are listed in order of increasing c/r2, so that geochemically similar elements to preclude an important role for variable degrees of
elements are plotted near one another. Chondrite- or mantle- partial melting in the observed chemical variability. This is not to
normalization was deemed to be inappropriate in this case be- suggest that a varying proportion of partial melting was not criticause of the tremendous differences between rhyolites and these cal to the generation of topaz rhyolites, but only that the evidence
other materials. For example, the FICI ratio of chondritic mate- of such a process has been obscured by subsequent fractionation
, processes.
rial is approximately 10.
As indicated throughout this discussion, an examination of
The tremendous enrichments in F, Ta, Nb, Y, Rb, U, Th, .
and HREE and the relative depletion of Zr, Sr, Eu, and Ba are the elemental composition of cogenetic lavas reveals the imporeasily visualized in these diagrams. Chlorine concentrations in tance of crystal fractionation in the evolution of topaz rhyolites.
glasses show up as deep anomalies, indicating the lack of enrich- Those centers examined in any detail document chemical trends
ment relative to F and Rb. Some topaz rhyolites show depletions consistent with crystal fractionation. The correspondence be-
56
1000
syn-COLG
400
-.
E
c.
c.
......
.c
a:
100
81'hOPT~
WPG
ORG
VAG
10
100
10
1.0
Vb +Ta (ppm)
1000
Topaz
Rhyolites
syn-COLG
.....
E
c.
.....c.
.c
100
a:
WPG
VAG
100
10
~
,,(p+Nb
(ppm)
Figure 39. Trace element discriminant diagrams for granitic rocks
(Pearce et al. 1984) showing the compositions of topaz rhyolites from the
western United States, the Bishop Tuff (Hildreth 1977), and the Tala
Tuff (Mahood 1981). VAG = volcanic arc granites; ORG = ocean ridge
granites; WPG = within plate granites; and syn-COLG = syn-collision
granites. Topaz rhyolites clearly overlap the within-plate and collisional
fields.
tween major and trace element models for the evolution of the
Thomas Range rhyolites was noted above, and involved the fractionation of sanidine >quartz >plagioclase »biotite >bxides
>> apatite>zircon>allanite; The extreme depletion offeldsparcompatible elements is thus explained as are enrichments in a
variety·of incompatible trace elements. Crystallization appears to
have occurred close to the mi~imum in the granite system (Q-Ab-
Or-An-F-H20), consistent with the observed small changes in
Si02, A1203, and the alkalies. These variations correlate with
larger changes in trace element concentrations controlled by the
fractionation of major phases and by the observed accessory minerals (i.e., zircon, allanite, apatite, titanite). The correlation of
higher F with lower Si02 and higher Al203 and Na20 in the
rhyolites is also consistent with crystallization differentiation of
more F-rich magmas as predicted by the experimental studies of
Manning (1981). In short, although these rhyolites are among the
most F-rich magmas yet analyzed, we see no compelling evidence
to indicate that convection-driven thermogravitational diffusion
(Le. volatile complexing in the melt or Soret diffusion) controlled
their evolution (cf. Mahood 1981; Hildreth 1981).
The importance of crystal fractionation is strikingly illustrated by the correspondence of the trace element patterns of the
rhyolites with distinctive accessory mineral phases. (Wolf and
Storey (1984) have used this principle in a convincing explanation of the trace element zonation patterns seen in highly alkaline
magma bodies.) In topaz rhyolites and other aluminous rhyolites
the negative Zr (and usually Hf) anomalies and decreasing Zr/Hf
ratios are controlled by fractionation of zircon. Fractionation of
allanite, and perhaps monazite or chevkinite in some lavas, produce the typical LREE depletion patterns shown in Figures 40
and 41. In some rhyolites, titanite appears as a microphenocryst
and its fractionation is probably responsible for the Y, and middle
REE depletions apparent for lavas from the titanite-bearing rhyolites of the Mineral Range, Utah, Lake City, Colorado, and other
areas as noted by the dashed lines in Figure 41. Elements that are
largely incompatible in the major and trace phases become substantially enriched. Such elements include both large-U, Th, Rb,
es, Ta-and small ions-Be, Li. Fluorine also behaves as an
incompatible element as it is only removed by biotite fractionation and biotite occurs in small proportions. Moreover, the biotite/
melt partition coefficient for Fis strongly dependent on the Fe/
Mg ratio of the biotite (Munoz 1984) and hence on the fugacity
of oxygen. On the other hand, the negative Cl anomaly must be a
reflection of the original composition of topaz rhyolite melts,
because Cl should be more strongly incompatible than F.
Individual volcanic centers need to be examined in more
detail to test these conclusions regarding the geochemistry and
differentiation of topaz rhyolites. Of particular importance would
be studies of the halogen· compositions of lavas and/or tuffs that
maybe cogenetic with topaz rhyolites. Changes in F/Cl ratios
could be monitored in this way. Another fruitful avenue of research would be to examine in more detail the relationship of
accessory mineral assemblages and compositions with the trace
element patterns of their host "liquid." Befbre unique geochemical models can be constructed, mineral/liquid partitioning needs
to be studied, but has thus far been completely neglected for the
rhyolites described here. V. I. Kovalenko and others (e.g. 1978
and 1984) have been examining crystal/liquid partitioning in
ongonites. Studies of the geochemical effects of devitrification
have thus far been cursory. The mineralogic controls on this
process need to be determined. Experimental studies are called
Topaz Rhyolites
--
5""-61
57
a
.....
Thom•• Range. UtAh
_
100
SM·2~:z.06
--...-_-- ---
10
-- ...
b
------- ..
------_ ... ---- .._-.
r----='SPc:;O':.:M::..:.n:::••.::;uT'--_- ~ 100
-- ~ ...
....................................
Black R.ngc.NM
................................
Nathrop.CO
...•
10
c
Honeycomb Hill.. Utah
100
10
d
Smelter KnoU.
Utah
100
10
100
e
Minerai Mount.lna. U,...
10
Wah W.h Mountalna
------------- ...
10
-WW·9
___ a
..
WW'41
STC-4
100
g
~---
Nalhrop; lopal
~"':I~~;"-""""''''
---.,
Southw •• tern Colorado
\
I
\
10
,,
\
\
\
"11
,
I
0.1 ~-:'-_-..I._---J
_ _....L-:'--L-L--'
l.
c.
100
Nd
Sm
[u
Gd
Tb
Dy
--JL.....L..:J
Vb L ...
Figure 40. Chondrite-normalized REE patterns for topaz rhyolites and
related rocks from the western United States. Concentrations normalized
to 0.83 times concentrations in Leedey chondrite (Masuda et al. 1973).
a. Thomas Range, Utah (Christiansen et al. 1984a). Sample SM-61
contains relatively low concentrations of incompatible trace elements
(e.g. 5 ppm U) while SM-29-206 contains high concentrations of these
elements (e.g. 19 ppm U). La/LuN and Eu/Eu* are negatively correlated
with U content and apparently decrease as a result of differentiation of
the parental magma. HREE concentrations increase with differentiation.
The lava from which SM-61 was collected appears to be topaz-free, but
Christiansen et al. (1984a) postulate that it is genetically related to the
topaz-bearing rhyolites of the Thomas Range.
b. Specimens from Spor Mountain, Utah (Christiansen et al. 1984),
Black Range, New Mexico (HC-8; Correa 1980), and Nathrop, Colorado(Zielinski 1977) cover the range of REE concentrations seen in
topaz rhyolites from the western United States.
c. Honeycomb Hills, Utah (Turley and Nash, 1981). The low Si, high
Ca and F sample described in the text has a very similar pattern.
d. Smelter Knolls, Utah (average of three analyses; Turley and Nash
1980).
e. Mineral Mountains, Utah, topaz rhyolites (average of samples from
two domes) compared with less differentiated but probably co-genetic
lava (Bailey Ridge flow) (Lipman et al. 1978a). Depletion of light and
middle REE, and especially Eu, with differentiation is apparent. Yb and
Lu show slight enrichment.
f. Wah Wah Mountains, Utah (Christiansen et al. 1980). Samples WW9 and WW-41 are late Miocene rhyolites from the Broken Ridge area;
STC-4 is an early Miocene rhyolite from the plug at the Staats (U-F)
mine.
g. Nathrop, Colorado, topaz rhyolite (Zielinski 1977) compared to calcalkaline rhyolite REE pattern from Summer Coon volcano in the nearby
San Juan volcanic field (Zielinski and Lipman 1976). These rhyolites are
not co-genetic. Note the lower La/Yb ratio and deeper Eu anomaly of
the topaz rhyolite. These features are typical of comparisons between
topaz rhyolites and nearby, often nearly contemporaneous,calc-alkaline
rhyolites.
58
Topaz Rhyolites
Topaz Rhyolites
Mexican Tin
Rhyolites
10
10
5
5
~
I
I
::;;
::;;
(!l
(!l
0:
0:
....
11>
11>
c.
C.
E
'"
0.5
E
0.5
'"
(/)
(/)
0.1
0.1
F
a
CI
Rb Sa Sr Eu
La Sm Yb
Y
Th
U
Zr
Nb Ta
Peralkaline Rhyolites
10
5
11>
C.
E
0.5
'"
(/)
•
PantelJeria
o Sierra La Primaveva
0.1
• McDermitt
b
F
CI
Rb Sa
Sr Eu La Sm Yb
Y
Th
U
Zr Nb Ta
c
F
CI
Rb Sa Sr Eu
La Sm Yb Y
Th
U
Zr
Nb Ta
Figure 41. Concentrations of selected trace elements in rhyolitic rocks
normalized to those in RGM-l, a rhyolite obsidian from Medicine Lake
volcano, California, and U.S. Geological Survey geochemical reference
(Govindaraju 1984). The elements are listed in order of their relative
field strength, increasing to the right. a) Topaz rhyolites from the western
United States typically show negative CI, Ba, Sr, Eu, and Zr anomalies
on such diagrams. Titanite-bearing rhyolites commonly show negative
Y anomalies shown with dashed lines; titanite-free lavas show no negative Y anomalies. b) "Tin" rhyolites from Mexico (vertical bars; Huspeni
et al. 1983), which commonly have vapor-phase topaz as well, have
trace element compositions that largely overlap with those of topaz
rhyolites from the western United States (shaded). Chlorine concentrations from glassy specimens have not yet been reported. c) The trace
element patterns of peralkaline rhyolites (Civetta et al. 1984; Mahood
1981; Conrad 1984; Christiansen et al. 1980) are strikingly different
from those of topaz rhyolites (shaded) and generally have small negative
(or in some cases even positive) CI anomalies. In addition, they generally
have higher REE contents, but marked negative Th and U spikes are
distinctive. The absence of negative Zr anomalies is the result of the
enhanced solubility of zircon in peralkaline melts.
Topaz Rhyolites
for to determine the role of fluorine on melt properties and phase
relationships and to constrain the values of intensive properties of
each system.
Isotopic composition
Initial strontium isotope ratios for 18 samples from 10 localities are reported in Table 3. These analyses indicate that initial
87Sr/86Sr ratios for topaz rhyolites range from 0.7055 to about
0.712. Considerable uncertainties exist in the initial 87Sr/86Sr
because of poorly known ages for some of the samples, coupled
with their extremely high RblSr ratios. The effects of recalculation to different ages are shown for the rhyolites from Nathrop,
Colorado, and the Black Range, New Mexico. These isotope
ratios are suggestive of an important crustal component in the
protolith of topaz rhyolites. Using approximate crustal province
ages (Farmer and DePaolo 1984), these ratios correspond to
crustal sources with relatively low RblSr ratios. For example, an
initial 87Sr/86Sr of 0.709 from the eastern Great Basin corresponds to a RblSr ratio in the source of about 0.07 (assuming
separation of the crust 2.2 Ga ago from a bulk-earth reservoir
with RblSr ratio ofO.029-parameters from DePaolo and Wasserburg 1977); a ratio of 0.706 implies a source RblSr of 0.04.
Initial ratios between 0.709 and 0.7055 from Colorado and New
Mexico (with crustal ages of 1.7 to 1.8 Ga) imply RblSr ratios of
0.08 to 0.04. These elemental ratios are consistent with the hypothesis that topaz rhyolites are derived from high-grade
metamorphic rocks, perhaps of granulite grade (see compilation
in Pettingill et al. 1984). Moreover, these values are lower than
would be expected if the rhyolites were derived from metasedimentary materials of similar Proterozoic ages. For example,
Farmer and DePaolo (1984) estimated that the RblSr ratios for
the sources of strongly peraluminous (S-type) granitoids in the
western Cordillera range from 0.3 to 0.05. These aluminous granitoids are thought to be derived from a middle crustal
metasedimentary source (e.g. Lee and Christiansen 1983; Farmer
and DePaolo 1983). Upper crustal contamination of cogenetic
topaz rhyolites may be indicated by variable initial Sr-isotope
ratios for samples from the Thomas Range, which belong to a
coherent suite as judged from trace and major element geochemistry (Christiansen et al. 1984).
Oxygen isotope ratios for topaz rhyolites are available only
from the Mineral Mountains, Utah, (Bowman et al. 1982) and
Lake City, Colorado, (R. A. Zielinski, written communication,
1982). These bear out the suggestion of a crustal source in their
low to moderate values (8 180 = 6.3 to 6.9 %0 for the Mineral
Range and 7 to 10 %0 for Lake City).
Interpretation of the Pb-isotopic data in such a manner is
equivocal as the· two topaz rhyolites thus far examined (Lake
City, Colorado, and Cortez, Nevada) do not show elevated
208Pb/ 204Pb ratios. Elevated thorogenic Pb isotope ratios are
commonly taken to indicate derivation from granulitic rocks because of the presumed depletion of U relative to Th during metamorphism. Although many granulite terranes do show high
59
ThlU ratios (>5), many, especially felsic granulites, do not (see
compilation in Iyer et al. 1984). For the Lake City rhyolites,
Lipman et al. (1978b) have interpreted the Pb isotope data to
indicate derivation from an amphibolitic lower crustal source
with ThlU ratio of about 3.6. Another important feature of the
Pb-isotopic composition of both the Cortez and Lake City rhyolites is their similarity to chemically and temporally distinct rocks
in the same region. It appears that Pb isotope regionalization
noted by Zartman (1974) and others extends to these silicic volcanic systems as well.
Integrated isotopic (Pb, Sr, Nd, 0) and petrologic studies of
topaz rhyolites are of paramount importance for determining the
geochemical nature of the protolith of topaz rhyolites. Given the
probable crustal origin of topaz rhyolites, these studies might be
informative of crustal structure and composition when examined
on a regional scale, but studies of single volcanic complexes are
needed before the role of crustal contamination versus source
inheritance can be fIrmly established.
Magma-tectonic setting
The nature of igneous rock associations and contemporaneous tectonic activity gives some clues about the generation of
magmas. For example, lithospheric subduction at continental
margins is consistently associated with concurrent calc-alkaline
magmatism. However, any attempt to delineate the magmatectonic setting of topaz rhyolites in the western United States is
limited by our incomplete understanding of the complex evolu7
tion of Cenozoic tectonism across the region.
Extensional tectonism appears as a common denominator in
almost all areas where topaz rhyolites were erupted in the western
United States. Episodes of topaz rhyolite magmatism coincide
with periods of lithospheric extension; 1) in the eastern Great
Basin where basin and range faulting may have begun as early as
21-20 Ma (Rowley et al. 1978a) and then was renewed under a
different stress orientation about 10 Ma which has persisted to the
present (Zoback et al. 1981); 2) along Nevada's Cortez rift that
opened 16 Ma (Stewart et al. 1975); 3) in Montana where block
faulting began about 40 Ma (Chadwick 1978) and intra- or backarc graben formation may havebegun as early as 50 Ma (Armstrong 1978); 4) along the Rio Grande rift and its northern
extension into Colorado, which initially developed about 30 Ma
(Eaton 1979; Elston and Bornhorst 1979); and 5) in western
Arizona where block faulting was underway by about 9 Ma
(Sunneson and Luchitta 1983). Several groups of topaz rhyolites
lie on possible continental transform zones that developed as a
result of differential extension rates (Wah Wah Mountains, Kane
Springs Wash, Spor Mountain, Elkhorn Mountains). Lithospheric extension occurred in back- or intra-arc and post-arc
environments (Eaton 1979, 1984a; Elston and Bornhorst 1979).
The intimate association of extensional tectonics and topaz rhyolite magmatism in the western United States implies a strong
genetic connection, but the nature or existence of extension concurrent with topaz rhyolite magmatism in Mexico needs to be
clarifIed.
60
TABLE 10. MAGMATIC ASSOCIATIONS OF TOPAZ RHYOLITES
Calc-alkaline Suite
Bimodal Basalt-Rhyolite Suite
Alkaline Suite
Andesitic to dacitic lavas
and rhyolitic ash flows
associated with granodioritic
intrusions. Continuous Si02
variation diagrams.
Generally potassic basalts
of tholeiitic or alkaline
affinity.
Alkaline to peralkaline
tuffs, lavas, intrusions
Alkalic basalts may also
occur. Trachytes
present.
Chalk Mountain, CO
Nathrop CO
Tomichi Dome, CO
Black Range, NM
Sierra Madre Occidental
Mexico
Thomas Range, UT
Spor Mountain, UT
Smelter Knolls, UT
Honeycomb Hills, UT
Wah Wah Mountains, UT
L. Pliocene
Cortez, NV
Jarbidge, NV
Sheep Creek Mtns, NV
Lake City, CO
Boston Peak, CO
Burro Creek, AZ
Little Belt Mountains,MT
Silver Cliff, CO
Grants Ridge, NM
Probable
E. Miocene Wah Wah Mtns, UT
Climax-type Mo deposits
Henderson, CO
Climax, CO
Pine Grove, UT
Climax-type Mo deposit
Mt. Emmons, CO
The magmatic associations of topaz rhyolites are less
straightforward. Following Lipman et al. (1972) and Christiansen
and Lipman (1972), the magma-tectonic evolution of the western
United States may be divided into two fundamentally different
stages. An early suite of subduction-related calc-alkaline magmas
was time transgressive across the western United States and produced eruptions of andesitic lavas, dacites, and rhyolites; the latter
mostly occur as ash-flow tuffs associated with caldera collapse.
After about 20 Ma, the Basin and Range region experienced the
eruption of bimodal suites of basalt and rhyolite associated with
lithospheric extension. Elsewhere, the temporal relationships are
different, but the magmatic and tectonic products of each association can generally be identified. In the southern Rocky Mountains, Elston and Bornhorst (1979) defined an episode transitional
between these two associations, which is typified by eruptions of
basaltic andesite and high-silica rhyolite in a modified back-arc
extensional environment. We include these rocks with those of a
predominantly calc-alkaline character (cf. Ratte et al. 1984).
Topaz rhyolites appear to have been produced in all of these
igneous associations. In addition, several topaz rhyolites are contemporaneous with the emplacement of distinctively alkaline
magmas. Thus we have tentatively identified three principal
magma associations in which topaz rhyolites are found; 1) intermediate to silicic calc-alkaline volcanic suites; 2) bimodal basaltrhyolite associations; and 3) alkaline to (silica-saturated) peralkaline suites in which trachytes (or their intrusive equivalents) are
important. Table 10 shows these groups. Reference to Figure 42
may be helpful in this regard.
In some cases, the definition of the calc-alkaline group is
problematic. The Oligocene topaz rhyolites of Colorado are all
grouped with the calc-alkaline association because they are
broadly contemporaneous with volcanic rocks that show extended and continuous SiOz-variation diagrams, but lack basaltic
compositions. Some may be part of a "transitional" group
associated with the change from subduction-related magmatism
to extension-related magmatism (White et al. 1981; Bookstrom
1981). In addition, the early Miocene rhyolites of the Wah Wah
and Needle Ranges in Utah form a restricted bimodal
(trachyandesite-rhyolite) association within the mountain range,
but consideration of volcanism op. a broader scale (southwestern
Utah) could allow them to be grouped with the calc-alkaline
magmatic association.
The second group consists of topaz rhyolites clearly in bimodal association with more mafic rocks-the fundamentally
-basaltic group of Christiansen and Lipman (1972). The compositional variety of late Cenozoic "basalt" makes the definition of
this group somewhat arbitrary. The mafic members range in
composition from basalt to andesite and are variably alkaline. We
have also included the Oligocene rhyolites of the Elkhorn Mountains in the group because of their association with basaltic lavas
and the apparent absence of intermediate composition rocks.
A third type of magmatic association includes the alkaline
rocks associated with topaz rhyolites at Kane Springs Wash volcanic center, Nevada (trachyte to peralkaline rhyolite); Silver
Cliff-Rosita, Colorado (andesite, trachyte, rhyolite); and in the
Little Belt Mountains, Montana (trachybasalt to trachyte; quartz
monzonite to syenite). The Grants Ridge rhyolites in New Mexico are associated with the construction of an "andesitic" volcano
at Mount Taylor, which is surrounded by alkali basalt to trachyte
lava flows, and could be placed with the second group.
In spite of the vagaries of any such classification, it is obvious that topaz rhyolites are associated with a variety of igneous
Topaz Rhyolites
5
a
6
61
e
.
6
• 0
4
4
4
2
2
2
U
5
..
0
b
5
0
6
~
8
4
4
Q
2
2
~
~
~
N
N
U
N_6
C
~
~
0
N
0
::s::
::s::
2
+'
+'
+'
0
4
3
5
N
::s::
4
Specimen Mtn
0
5
k
4
2
011"-_-+-_-+-_-+-""""1
0
6
..
d
5
h
5
Q
4
4
2
2
ll'5
55
SiOZ
55
(ivt/.)
75
55
55
Si02 (wtiO
75
55
65
~+
75
5 i 02 (wt7.)
Figure 42. Potassium versus silica variation diagrams for mafic to intermediate composition volcanic
rocks associated with topaz rhyolites from the western United States. The original sources of the data are
given in the description of the individual occurrences. Representative analyses of topaz rhyolites are
indicated by crosses, other volcanic rocks as open boxes. Most topaz rhyolites occur in strongly bimodal
associations with variably potassic mafic rocks (e.g. basalt or basaltic andesite). Topaz rhyolites associated with trachytic magmas do not show strong SiOz gaps and have variation diagrams which extend
to mafic rocks (e.g. Little Belt Mountains, Kane Springs Wash, Grants Ridge/Mt. Taylor). A few, like
those in the Black Range (and the Oligocene of Colorado, not shown), have intermediate to silicic
calc-alkaline trends which include topaz rhyolites at their high silica ends. Lines are those used by Ewart
(1979) to define low-K, interrnediate-K, and high-K rock series.
rocks. Indeed, they show no consistent spatial or temporal relationship to a single magma series from which they could be
derived by differentiation. We suggest that none of these more
mafic magmas are parental to the F-rich rhyolites discussed here.
Instead, the observation that topaz rhyolites are associated with a
variety of more mafic magma suites suggests that the rhyolites
have a thermal relationship to the more mafic magmas. The
residence of these mafic magmas in the crust may have provided
the thermal energy required for melting to produce magmas parental to topaz rhyolites.
Ore deposits
Mineralization associated with topaz-rhyolite magmatism
generally consists of F, Be, Li, Cs, U, Sn, Mo(?), and W(?)
(Table 11). The marked magmatic· enrichment of these same
elements in topaz rhyolites strongly suggests that.the ore elements
were derived from the rhyolites or their intrusive relatives in the
case of Climax-type Mo-W deposits (Burt and Sheridan 1980;
Burt et al. 1982). Other types of mineralization (alunite, Hg,
Au-Ag) are spatially and temporally associated with some topaz
rhyolites. The association of these deposits with the rhyolites may
rely more on magmatic heat content and volcanologic style for
their generation than on any particular compositional feature of
topaz rhyolites. A variety of these ore deposit environments are
schematically indicated in Figure 44.
Beryllium. The most important ore deposit directly associated with a topaz rhyolite is the beryllium deposit at Spor
Mountain, Utah. It is currently (1985) the only important source
of Be in North America. Bertrandite (Be4Siz07) occurs in the
62
TABLE 11. MINERALIZATION ASSOCIATED WITH
TOPAZ RHYOLITES
Mineralization
Location
Thomas Range, UT
Spor Mountain, UT
Honeycomb Hills, UT
--~-
Be, U, Li, F
Be, Rb, Cs, Li
Smelter Knolls, UT
Keg Mountain, UT
Wah Wah Mtns. vicinity, UT
U, F, Hg, Au, Ag, Mo(?)
Cortez Mountains, NV
Jarbidge, NV
Au, Ag, Hg
Sn
Au, Ag
0 36 ••••
o
24
•
Ag, Pb, Zn, Mo(?)
Mo
Chalk Mountain, CO
Nathrop, CO
Mo, W
F(?)
Silver Cliff/Rosita, CO
Tomichi Dome, CO
Boston Peak, CO
Lake City, CO
Grants Ridge, NM
Black Range, NM
Saddle Mountain, AZ
Burro Creek, AZ
00
00
30~028
17~~
o
0
•
023
Ag, Au, Pb, Zn, Cu
U
Sn
Figure 43. The distribution and ages (Ma) of granitic Climax-type porphyry molybdenum deposits (open circles) and prospects (x) (Westra
and Keith 1981; White et al. 1981), compared to the locations of topaz
rhyolites (filled circles) in the western United States.
upper part of a tuff beneath a rhyolite lava flow. Likewise, ber- ened the alteration process in these intrusions to the formation of
trandite is probably the source of the Be-mineralization at the topaz-bearing lithophysae. Garnet, another aluminous mineral
Honeycomb Hills, Utah (Lindsey 1977). On the other hand, beryl common in topaz rhyolites, is found in other molybdenum deposoccurs in a number of other topaz-bearing lava flows including its or prospects (e.g. Pine Grove, Utah: Keith 1980; Mt. Hope,
some lavas in the Thomas Range, Utah, the Wah Wah Moun- Nevada: Westra 1982; and Henderson, Colorado: Gunow et al.
tains, Utah, and in the Taylor Creek Rhyolite in the Black Range 1980). The similarity of biotite compositions, which are sensitive
of New Mexico. In each case, beryl occurs within intensely devi- indicators of volatile fugacities, in both types of magmatic systems
trilled lavas. Late magmatic beryl segregations also occur in the has already been noted in terms of their F and CI contents and
topaz-bearing Sheeprock granite of west-central Utah (Williams ratios.
1954). The association of beryllium mineralization with topaz
2) The spatial distribution and ages of the major known
rhyolites strongly suggests that the magmatic enrichment of Be in Climax-type Mo deposits are illustrated in Figure 43. Their locathe rhyolites is important in the genesis of the deposits. Indeed, Be tions are shown superimposed on the distribution of topaz rhyoconcentrations average about 60 ppm in glassy specimens of the lites. The spatial and temporal correspondence of both types of
rhyolite from Spor Mountain, a value about 20 times that found magmatism in Colorado, Montana, and Utah have been noted in
in an average granite.
the individual descriptions and suggest some sort of genetic link
Climax-type molybdenum deposits. The association of exists between these distinctive groups of rocks.
topaz-rhyolite magmatism and "Climax-type" Mo-W deposits
3) The tectonic setting of both types of magmatism appears
has been noted by Burt and co-authors (1980,1982), Westra and to be in continental rifts or in zones of back-arc extension (e.g.
Keith (1981), White et al. (1981), and probably many others in Sillitoe 1980). Both magma types a.re emplaced in the upper crust
industry. By way ofjustifying this correlation we note the follow- as relatively small stock-like intrusions that explosively vent to
ing similarities.
the surface to emplace relatively small volumes of tuff and/or
1) Topaz rhyolites are mineralogically similar to the igneous lava (J. E. Sharp 1978; Keith 1982).
and metasomatic rocks associated with Climax-type Mo deposits.
4) Multiple intrusion/extrusion episodes are apparent for
Most molybdenite deposits associated with granitic (or rhyolitic) both types of magmas.
rocks cOl1tain topaz in their mineralized zones. Burt (1981) lik5) The chemical similarity of the magmas involved is shown
Topaz Rhyolites
63
Dome
4
Pyroclastic
Deposit
Pre-existing Sediment
Figure 44. Schematic cross-section showing the hypothetical structure of a small rhyolite dome complex
and some of the types of mineralization possibly associated with topaz rhyolite volcanism (after Burt et
al. 1982). 1 = hQt springs deposits (Ag, Au, W, Mn, etc.) within rhyolite or associated volcanic rocks; 2 =
clastic sedimentary rocks beneath tuffs (D); 3 = mineralized pyroclastic deposits (Be, D, F, Li, Cs, etc.); 4
=fractured and flow-banded lavas (Sn); 5 =vent and contact breccias (F, D, etc.); 6 =base and precious
metal veins (Ag, Pb, Zn, Au, Sn, W, etc.); 7 = mineralized breccia pipes (Mo, Ag, Au, F); 8 = stockwork
porphyry deposits (Mo, Sn, W, etc.); 9 = fluorite-rich skarn and/or sulfide-rich replacement ore bodies
in non-eruptive environment (Sn, W, Be, etc.); and 10 =greisen-bordered veins in non-eruptive environment (Sn, W, Cu, Zn, Be, etc.).
in Table 12. The granites of Climax-type Mo-systems and topaz
rhyolites share their major characteristics: high Si, K, Na, and F
and low Ti, Fe, Mg, and Ca (cf. Figure 35). The principal differencesbetween the rocks lie in the lower concentrations of P and
Mg in topaz rhyolites. Differences in the Na and K contents are
probably not meaningful because most Climax-type intrusions
have experienced some potassic alteration. The trace element
signature of topaz rhyolites (high D, Be, Sn, Li, Nb, Rb, and F
and low Sr, Ba, and Ti) is typical of Mo-related systems as well
(Westra and Keith 1981; Keith 1980; Mutschler et al. 1981;
White et al. 1981).
The similarities in distribution, age, tectonic setting, mode of
emplacement, chemistry, and mineralogy of fluorine-rich subalkaline rhyolites (with topaz or garnet) and Mo-mineralized
"rhyolite" stocks leads us to conclude that the eruption of topaz
rhyolites may be a surface manifestation of a potentially oreforming intrusive system.
Tin. The small deposits of cassiterite and wood tin that
occur in the Sheep Creek Range, Nevada, and in the Black
Range, New Mexico, are similar to the numerous small Sndeposits of Mexico. Burt and Sheridan (1984) and Duffield et al.
(1984) conclude that the high temperature vapor-phase deposition of cassiterite (and topaz) results from the extraction of Sn
from the rhyolitic glass or lava. The halogens appear to be effective complexing agents for Sn (Manning 1981b; Jackson and
Helgeson 1985). Breccias and other permeable zones (i.e. flow
bands and fractures) in the upper parts of these lavas are favorable locations for the accumulation, decompression, and cooling of
metal-bearing vapors or fluids, resulting in the common association of vapor-phase features increasing in abundance toward the
tops of rhyolite domes and flows. Low-temperature remobilization of Sn by circulating meteoric waters may lead to the deposition of wood-tin in narrow veinlets in cooling and flow fractures.
The small size of such deposits will probably prevent their successful exploitation in this country. Nonetheless, erupted topaz
rhyolites may be indicators of topaz granites that may develop
64
TABLE 12. AVERAGE COMPOSITION OF SILICIC INTRUSIONS
RELATED TO Mo-DEPOSITS
1
S.D.
2
S.D.
3
S.D.
Si02
Ti02
A1203
76.1
0.11
12.7
1.04
0.11
0.46
75.8
0.28
13.4
2.23
0.73
0.82
75.6
0.14
12.8
0.71
0.09
0.38
Fe203
FeO
MnO
0.84
0.55
0.06
0.36
0.26
0.05
0.71
0.42
0.04
0.83
0.38
0.03
1.12* 0.19
0.06
0.01
MgO
CaO
Na2 0
0.46
0.74
3.30
0.23
0.32
0.38
0.24
0.66
3.67
0.23
0.45
0.48
0.15
0.83
3.73
0.07
0.37
0.28
K20
P2 0 5
F
5.08
0.09
0.21
0.60
0.44
0.21
4.85
0.05
0.11
0.51
0.06
0.11
5.04
0.00
0.23
0.35
0.01
0.09
Note: All analyses in weight percent and recalculated
H20, C02 and S02 free.
S.D. is one standard deviation.
*Total Fe as Fe203
1. Average of 13 unaltered ore-related granite and
rhyolite porphyries from Mo deposits (Mutschler et
al. 1981).
2. Average of 50 granite and rhyolite porphyries from
unmineralized stocks near Mo deposits (Mutschler et
al. 1981).
3. Average of 14 rhyolite lavas from the Thomas Range,
Utah (Christiansen et al. 1984).
economic deposits of Sn (or W) in greisens or skarns (e.g., East
Kemptville, Nova Scotia: Richardson et al. 1982; and Anchor
Mine, Tasmania: Groves 1972).
Uranium. Small, generally sub-economic, deposits of uranium are associated with many topaz rhyolites. The enrichment
of U in topaz rhyolites probably accounts f<>r this association.
Notable examples include the rhyolites at Spor Mountain, Utah,
in the Wah Wah Mountains, Utah, and those near Lake City,
Colorado. The Be tuff member of the Spor Mountain Formation
contains low-grade U (and Th) mineralization that overlaps the
Be ore zone (Lindsey 1982; Bikun 1980). The U occurs in fluorite
and opal. Accumulations of U (as uranophane and weeksite) also
occm in a small lens of non-volcanic conglomerate at the base of
the Be tuff. The· U was probably leached by groundwater from
the U-rich tuff and is not associated with enrichments of Be or Li.
In the Wah Wah Range, U occurs in small altered horizons with
Fe~oxides and in conformable lenses with pyroclastic deposits
beneath topaz-bearing lavas (Christiansen 1980).
Fluorite. Fluorite deposits are closely associated with topaz
rhyolites from Spor Mountain, the Wah Wah Mountains, and
Needle Range, Utah, and perhaps with the Nathrop Volcanics of
central Colorado. Fluorite occurs in calcic rocks (sedimentary
carbonates or intermediate composition volcanic rocks) spatially
associated with topaz rhyolite vent complexes and in the Spor
Mountain district in tuff-lined breccia pipes. A hint to the origin
of the F. enrichment of topaz rhyolites and the generation of
fluQrite deposits in general lies in their distribution, as reiterated
most recently by Eaton (1984b). It has been known for several
decades (e.g., Peters 1958) that fluorite deposits are more common in the eastern part of the American Cordillera. With the new
understanding of the accretionary history of the continental crust
of western North America, it is clear that fluorite deposits (as well
as Be, Sn, and Climax-type Mo) are restricted to terranes underlain by Precambrian craton. Paleozoic and Mesozoic accreted
terranes consisting of fragments of ocean floor and island arcs are
not typified by fluorite deposits; instead, an association with Au
and Hg mineralization is indicated. Christiansen and Lee (1985)
have shown that the granitoids of the northern Great Basin show
differences in F-content that correlate with their location; granitoids in accreted terranes are F-poor, whereas granitoids rooted in
Precambrian sial are variably enriched in F. They interpret this
difference as resulting from different F concentrations in the
crustal component of their parent magmas. These disparate
observations point to the Precambrian continental crust as the
ultimate source of the F (and probably Be, Sn, and Mo as well) in
the ore deposits and in topaz rhyolites.
COMPARISONS WITH OTHER TYPES OF
RHYOLITIC ROCKS
The geochemical distinctiveness of topaz rhyolites is clearer
when contrasted with other types of rhyolitic volcanic rocks. In
western North America, topaz rhyolites are nearly contemporaneous with calc-alkaline rhyolites and peralkaline rhyolites.
The calc-alkaline rhyolites are part of an early to mid-Cenozoic
suite of intermediate .to silicic composition (e.g., Lipman et al.,
1972). The peralkaline rhyolites are part of a late Cenozoic bimodal suite of basalts and rhyolites (Christiansen and Lipman
1972). In spite of their close temporal and spatial association with
these silicic rocks, topaz rhyolites are distinct from both.
Calc-alkaline rhyolites
Calc-alkaline rhyolites are the silicic representatives of the
orogenic magma series characterized by a lack of iron-enrichment
during its differentiation. Calc-alkaline rhyolites are typically associated with andesitic volcanism on continental margins overlying subduction zones. They generally occur as small domes or
lava flows associated with composite volcanoes or calderas but
may form voluminous ash-flow sheets. Large volumes of high-K
calc-alkaline rhyolite were erupted during the mid-Cenozoic of
the western United States.
Ewart (1979) has reviewed the chemistry and mineralogy of
the silicic orogenic volcanic rocks, including the calc-alkaline
series. He points out that the calc-alkaline rhyolites generally
contain phenocrysts of plagioclase, Mg-augite, Mg-hypersthene,
Ca-Mg hornblende, Mg-biotite, Fe-Ti oxides, and occasionally
olivine. High-K varieties contain quartz and sanidine. Zircon,
apatite, titanite, and allanite are notable accessory minerals. Although generally not fluid-saturated before eruption, the common
presence of hornblende and biotite in these rhyolites indicates that
they are relatively hydrous. The T-fo2 relationships for some
Topaz Rhyolites
65
TABLE 13. COMPARISON OF CALC-ALKALINE RHYOLITES WITH
TOPAZ RHYOLITE
Si02
Ti02
A1203
1
Utah
2
Colorado
3
Guatemala
4
Talasea
5
Taupo
76.0
0.3
13.0
77 .1
0.19
14.3
74.4
0.12
11. 9
75.3
0.27
l2.G
74.2
0.28
13.3
76.0
0.13
12.8
6
Topaz
Rhyolite
Fe203*
MnO
MgO
1.2
0.04
0.05
0.42
0.04
0.19
0.83
0.07
0.2
2.55
0.07
0.24
1. 89
0.05
0.28
1.07
0.06
0.10
CaO
Na20
K20
P205
1.1
3.1
4.4
2.7
4.6
4.4
0.02
0.8
3.5
4.3
1. 25
4.02
3.82
0.05
1. 59
4.24
3.18
0.05
0.74
3.73
5.00
0.00
59
124
109
150
55
200
145
107
106
126
450
20
1135
14
4
645
859
12
3
41
49
19
Trace elements (ppm)
Zr
Rb
Sr
100
300
85
350
Ba
Th
U
1500
24
8
2000
6
2
--- Not reported.
* Total Fe as Fe203
1.
2.
3.
4.
5.
6.
Joy Tuff, Black Glass member (Lindsey 1981) •.
Rhyolite from Summer Coon volcano (Zielinski and Lipman 1976).
Los Chocoyos ash (Rose et al. 1979).
Rhyolite from New Britain (Lowder and Carmichael 1970).
Average rhyolitic lava, New Zealand (Ewart and Stipp 1968).
Average topaz rhyolite from the Thomas Range, UT (Christiansen
et al. 1984; Ba from Turley and Nash 1980).
calc-alkaline rhyolites are shown in Figure 30. A wide variety of
studies indicate that most orogenic silicic rocks crystallize under
relatively oxidizing conditions-2 to 3 log units above the QFM
buffer (Ewart 1979; Hildreth 1981; Gill 1981). This property is
expressed in the Mg-rich nature of the mafic minerals including
biotite and hornblende. Calc-alkaline batholithic rocks from the
Sierra Nevada and intermediate to silicic rocks of western Utah
and eastern Nevada have strikingly different compositions ofbiotites and hornblendes when compared to those found in topaz
rhyolites from the western United States (Figure 31).
Although there is substantial chemical variation among calcalkaline rhyolites, they are generally richer in AI, Ti, Fe, Mg, and
Ca, and poorer in totalll1kalies and F (although data are sparse)
than topaz rhyolites (Ewart 1979). In Table 13, five analyses
representative of calc-alkaline rhyolites are compared with the
"typical" topaz rhyolite composition described above. Few analyses of the halogens in glasses from calc-alkaline rhyolites exist,
but substantial differences in FICI are indicated. FICI ratios may
be less than 1 in magmas related to subduction processes (e.g.
Garcia et at. 1979; Coradossi and Martini 1981). In terms of their
trace element characteristics, calc-alkaline rhyolites generally
have lower concentrations of Rb, U, Th, Nb, Ta and other in-
compatible elements, and have higher concentrations of Ba, Sr,
and other compatible elements than topaz rhyolites. The K-Th
and Th-U concentrations of calc-alkaline volcanic rocks from
west-central Utah are compared with topaz rhyolites from the
same area in Figures 44 and 45; the enrichment of topaz rhyolites
in U and Th is obvious. Little REE data exists for the suite of
calc-alkaline rhyolites that preceded the eruptions of topaz rhyolites in Utah, New Mexico, and Colorado. The relationships
shown for the topaz rhyolites from Nathrop, Colorado, and a
slightly older calc-alkaline rhyolite from the San Juan volcanic
field may be typical (Figure 40). The topaz rhyolite from Nathrop
is relatively depleted in HREE compared to other topaz rhyolites,
but it is nonetheless enriched compared to the calc-alkaline rhyolite from Summer Coon volcano. Likewise the rhyolite from
Nathrop shows a striking negative Eu anomaly. These.important
differences notwithstanding, the differentiation trends of calcalkaline rhyolites appear to be similar to those of topaz rhyolites,
but the extreme enrichments and depletions noted above are not
observed.
It is generally agreed that rhyolitic magmas may originate by
fractional crystallization of plagioclase, pyroxenes, and Fe-Ti oxides from dacite or rhyodacite (e.g. Ewart 1979). Crustal fusion
66
5
_-----r-.----...,.r-------,
eruptions (Noble and Parker 1974). Peralkaline rhyolites
generally contain phenocrysts of anorthoclase or sodic sanidine,
quartz, sodic ferrohedenbergite, aenigmatite, and fayalite (Sutherland 1974). Arfvedsonite and riebeckite generally crystallize as
o
devitrification products. Zircon and apatite are common acces00 00
sory minerals. Fe-Ti oxides mayor may not be present (Nicholls
o Q O
00
and Carmichael 1969). The anhydrous nature of most Fe-Mg
o
silicates and the common enrichment in Cl suggests that most
CJ
I:
peralkaline rhyolites were not fluid-saturated before eruption
o
CJ
(Bailey 1980). Chlorine partitions strongly into hydrous fluids
and would be quantitatively extracted from a saturated magma
..c::
•
••
• 0 •••••
Eowhile F prefers to remain in a magma that coexists with a fluid
I: 3 ~
\
(Burnham 1979). Such a pattern does not appear in peralkaline
••
rhyolites (Figure 37). The common absence of hydrous mafic
silicates indicates a low water fugacity andlor high temperature.
• •• •
Mineral geothermometry indicates that crystallization occurs at
temperatures generally exceeding 800°C (e.g. Wolf and Wright
1981; Conrad 1984; Mahood 1981; Ewart 1981). The Fe-rich
I
2
•
character
of the mafic silicates and estimates of f0 2 from co4
1
3
2
existing ilmenite and magnetite suggest that crystallization occurs
at low f0 2 in many silicic peralkaline magmas (between QFM
In U cone. (ppm)
and WM; Ewart 1981; Wolf and Wright 1981; Conrad 1984;
2 r------....,.r------'.r---------, Mahood 1981).
The most important chemical features of peralkaline rhyoo
lites relative to other rhyolites are high Fe, Mn, Ti, F, and Cl,
along with low Al and Ca (Table 14). They are distinct from
topaz
rhyolites in each of these characteristics except their generCJ
I:
ally
high
fluorine content.
•
•
•
o 11•
The
F and Cl content of peralkaline rhyolites is compared to
CJ
•
•
that
of
topaz
rhyolites in Figure 37. Peralkaline rhyolites have
•
•
FICI
of
less
than
3 and are easily distinguished from topaz rhyoN
:::.:::
lites on this diagram. The affinity of fluorine for continental settings has been pointed out by Bailey (1980), who showed that a
I:
o 1--.......
L-1
......11.....
..-l FICI ratio of 1 divides oceanic from continental peralkaline
2
3
4
5 rhyolites. Schilling et al. (1980), however, have shown that glasses
from tholeiitic mid-ocean ridge basalts have high FICI ratios
In Th cone. (ppm)
(averaging 8.5), but at much lower concentrations than those disFigure 45. Geochemical comparison of Cenozoic volcanic rocks from cussed here. F ranges from 150 to 400 ppm. Plume-type magmas,
west-central Utah. (a) Logarithmic plot of K20 versus Th (b) Logarith- associated with oceanic peralkaline rocks, have lower FI Cl ratios.
mic plot of Th versus U. Open circles = late Tertiary topaz rhyolites;
Extreme enrichments and depletions of certain trace eleclosed circles = mid-Tertiary calc-alkaline rhyolites and rhyodacites.
ments
characterize both peralkaline and topaz rhyolites, but perData are from Lindsey (1982) and Christiansen et al. (1980).
alkaline rhyolites from the western United States generally have
lower concentrations of Rb, V, Th, Ta, and Ba and higher conand assimilation may also play an important role in the differenti- centrations of Zr, Hf, Nb, and Zn (Figure 41; Christiansen et al.
ation of silicic orogenic magmas (e.g. Myers and Marsh 1981; 1983a). The contrast between the two types of magmas is most
Grove et al. 1982; Hildreth 1981).
clearly seen in the nature of the negative Th-V "anomalies" and
absence of negative Zr anomalies in peralkaline rhyolites when
Peralkaline rhyolites
compared to subalkaline rhyolites. Peralkaline rhyolites also lack
Peralkaline rhyolites contain a molecular excess of NazO + or have small negative Cl anomalies. In addition, peralkaline
KzO over Alz03, expressed as normative acmite (for F- and rhyolites generally have higher concentrations of LREE than do
Cl-free analyses). They are most easily recognized by the presence F-rich aluminous rhyolites. As a consequence they have steeper
of sodic pyroxenes or amphiboles as phenocrysts or as vapor- chondrite-normalized REE patterns (Christiansen et al. "1983a).
phase minerals. Many peralkaline rhyolites in the Great Basin Peralkaline rhyolites show differentiation trends (indexed by inwere erupted during the late Cenozoic in large caldera-forming creasing (NazO + KzO)1 Alz03 and incompatible trace elements)
-
.
-
00cP1iPo
. .....-: .
-
-
.
-
o
-
Topaz Rhyolites
TABLE 14. COMPARISON OF PERALKALINE RHYOLITES WITH
TOPAZ RHYOLITE
Comendites (1)
Pantellerites(l)
Topaz
Rhyolite (2)
range
average
average
range
average
74.0
0.21
11.6
1. 25
69.4-75.0
0.09-0.87
10.2-13.4
0.40-3.22
71. 2
0.37
9.11
2.38
FeO
MnO
MgO
1. 88
0.08
0.04
0.80-3.70
0.01-0.17
0.0-0.22
4.52
0.21
0.09
1.60-6.73
0.03-0.36
0.0-0.75
0.06
0.15
CaO
Na20
K20
0.36
5.35
4.46
0.0-1.12
3.99-6.39
3.49-4.98
0.45
6.44
4.40
0.06-2.04
4.68-7.83
3.39-4.90
0.83
3.73
5.04
P205
F
Cl
0.02
0.37
0.24
0.0-0.08
0.06-0.76
0.05-0.41
0.05
0.30
0.28
0.0-0.28
0.11-1. 30
0.06-0.82
0.00
0.33
0.06
Si02
Ti02
A1203
Fe203
67.4-74.9
75.6
0.14-0.65 0.14
6.30-11. 3 12.8
0.44-5.60 1.12*
* Total Fe as Fe203
1. Macdonald 1974a.
2. Average of 11 rhyolite lavas from the Thomas Range
(Christiansen et al. 1984; Turley and Nash 1980, for
Cl) .
67
mendite to pantellerite transitions are rarely observed. In contrast,
Hildreth (1981) invokes the action of an "extraordinarily halogenrich" flux· released from crystallizing basalt to produce partial
melting of the lower crust or an earlier accumulation of underplated gabbro to yield the parental magmas for peralkaline
rhyolites.
Aluminous bimodal rhyolites
Ewart (1979) established that (non-peralkaline) rhyolites of
bimodal associations are distinctive from most orogenic rhyolites
in their mineralogy and chemistry. He notes that they have
strongly Fe-enriched mafic silicates and appear to have crystallized at temperatures in excess of about 800°C and at low f0 2
(between QFM and NNO) as compared to calc-alkaline rhyolites. In addition, Ewart points out that these rhyolites generally
.exhibit fractionated trace element patterns-Ba, Sr, Cr, Ni, and V
are depleted and Nb, Pb, and La are enriched. Although we
regard topaz rhyolites as part of this group, a variety of rhyolite
types exists within it. At one extreme lie the high temperature,
pyroxene (and commonly one-feldspar) rhyolites of the Snake
River Plain region (Hildreth 1981; Hildreth and Christiansen
1984; Leeman 1982a; Wilson et al. 1983). These rhyolites are
typified by anhydrous mafic silicates such as pyroxene and fayalite. At another extreme lie two-feldspar rhyolites with low equilibration temperatures and biotite as the principal mafic phase, as
in topaz rhyolites. The bimodal rhyolites of the Coso Range,
California (Bacon et al. 1981) and at Twin Peaks, Utah (Crecraft
et al. 1981), are chemically similar to topaz rhyolites. Notable
contrasts between the two groups include the relatively high
KINa, Zr, Fe, and Ti of the first group, coupled with less extreme
enrichments of incompatible trace elements, including F. Most
investigators derive the parental magmas for bimodal rhyolites by
partial melting of the (lower) continental crust (Hildreth 1981;
Leeman 1982a; Ewart 1982; Christiansen et al. 1983a). Subsequent fractionation (near the minimum in the granite system) of
plagioclase, alkali feldspar, quartz, biotite or pyroxene, Fe-Ti oxides, and accessories (apatite, zircon, allanite, and monazite) leads
to the characteristic compositions of these high-silica rhyolites.
Fractionation may occur enroute to the surface or in (relatively
shallow) magma chambers. These rhyolites have a close thermal,
but not chemical, relationship to contemporaneous basalts that
appear to have provided the heat for crustal melting.
that differ markedly from topaz rhyolites. Trachyte to pantellerite
transitions are marked by decreasing AI, Ca, Ba, Sr, Mg, Sc, Ti,
Ni, and Co that correlate with increasing Na, Cl, Mn, Fe, Zn, Hf,
Zr, Ta, Y, Nb, REE, U, Th, Rb and occasionally Eu and P
(Macdonald and Bailey 1973; Noble et al. 1979; Civetta et al.
1984). Differentiation trends involving trachytes and comendites
are similar to those characteristic of more aluminous magmas
with decreasing Fe, Ti, and P; but Zr, Hf, and the REE (excluding
Eu) remain incompatible (e.g., Ewart 1982; Conrad 1984). Some
of these characteristic trace element features can be explained by
the high solubility of Zr in peralkaline melts and the consequent
lack of zircon fractionation. In a similar fashion, the absence of
stable REE-rich aluminosilicates like allanite, and perhaps phosphates like monazite as well, may explain the high concentrations
of LREE in fractionated peralkaline rhyolites.
Silicic peralkaline rocks occur predominantly in continental
rift environments or rift-like settings (Macdonald 1974b) and are
prominent members of bimodal volcanic suites. Peralkaline rhyolites also occur in oceanic islands and late orogenic suites but the
common feature linking all of the geologic environments is lithospheric extension. We have shown that this strong association Ongonites
with extension is also typical of topaz rhyolites.
Several authors (Burt and Sheridan 1981; Turley and Nash,
Peralkaline rhyolites can be derived by fractionation of alkali basalt through an intermediate trachytic composition. Most 1980; Christiansen et al. 1983a) have suggested that topaz rhyoquantitative major and trace element models invoke fractionation lites are similar to the so-called ongonites that occur in Mongolia
of plagioclase, andlor ternary alkali-feldspar, Fe-rich pyroxene, and the Trans-Baikal region of the U.S.S.R. (Kovalenko and
magnetite, apatite, and olivine from trachyte (or high-alkali da- Kovalenko 1984). Ongonites are defined as topaz-bearing
cite) to produce a vertically zoned chamber with comendite or "quartz keratophyres." They occur in subvolcanic dikes, stocks,
pantellerite residing in the upper part of the chamber (Barberi et and as lava flows with underlying pyroclastic depositS (Kovaal. 1975; Civetta et al. 1984; Parker 1983; Middlemost 1981; lenko et al. 1971; Kovalenko and Kovalenko 1976; Kovalenko et
Bevier 1981; Souther and Hickson 1984; Conrad 1984). Co- al. 1979). The primary minerals of ongonites are albite, potas-
68
300
,.------------------;;~--___"7I
200
100
E0.....00-
+
80
- ......
60
Z
40
30
20
200
100
300 400
a
600 800 1000
2000
Figure 46. Rb-Nb-Ta concentrations in ongonites from central Asia
compared to other rhyolitic rocks from western North America. a) Logarithmic plot ofRb and Nb. b) Logarithmic plot ofTa and Nb. Arrows
show differentiation sequences in topaz rhyolite dome complexes (WW
= Wah Wah Mountains; SM =Spor Mountain; TR =Thomas Range)
and from the Coso Range, California, rhyolites which are not known to
contain topaz (COSO - Bacon et al. 1981). The compositions of 4 ongonites that do not lie in the same field as others are shown with crosses.
The average compositions of three types of orogenic rhyolites from the
western United States (Ewart 1979) are also shown (open circles), along
with the composition of the Bishop Tuff, California (BT - Hildreth 1977)
and the mildly peralkaline Tala Tuff, Mexico (TT - Mahood 1981).
Most ongonites with low Nb/Ta and Nb/Rb ratios are from a dike near
Ongon, Mongolia, which contains columbite. The fractionation of this
mineral may have buffered evolving liquid compositions to Nb concentrations of approximately 60 ppm in much the same manner that zircon
saturation controls Zr concentrations. Data for topaz rhyolites and ongonites are from sources cited in text.
Rb(ppm)
200
,.-------------------r--~---____::>-----____:II
100
~
80
-
60
.0
40
E
Ongonites
c.
c.
Z
---
~~
--
..,., /
I
"
I
~
30
.<,.fl>o
::.Q\
~
20
10
L.-
1
b
---I:....-....:.-.......
_-L----I'--..L..-..L..-.L.....1~
2
3
5
10
Ta(ppm)
..L._.._
20
____J'______L_
30
40
_L_...................L....L....I
60
80 100
Topaz Rhyolites
69
TABLE 15. AVERAGE MAJOR ELEMENT COMPOSITION OF
ONGONITES FROM CENTRAL ASIA
1
Ongon
Mongolia
Si02
Ti02
A1203
2
3
Baga-Gazryan Arybulak
USSR
Mongolia
71. 4
74.4
16.9
15.3
4
Teeg-Uula
Mongolia
73.0
0.09
14.8
74.1
0.10
13 .5
5
Spor Mtn.
Utah
73.2
0.05
13.5
Fe203
FeO
MnO
0.27
0.26
0.18
0.21
0.82
0.05
1. 14
0.54
0.09
1. 20
1. 29*
0.58
0.05
0.06
MgO
CaO
Na20
0.20
0.34
5.29
0.19
1. 02
0.23
0.54
4.17
0.20
0.86
4.24
0.11
0.61
3.95
K20
P205
F
3.34
0.07
2.01
-3.67
0.05
0.82
4.14
0.05
0.96
4.60
0.04
0.52
4.86
0.00
1. 14
4.49
"
Note: All analyses recalculated H2O-free.
1. Average of 53 analyses of the Amazonitov dike (Kovalenko
and Kovalenko 1976).
2. Average of 6 ongonites (Kovalenko and Kovalenko 1976).
3. Average of 9 ongonites (Antipin et al. 1980).
4. Average of 3 volcanic ongonites (Kovalenko et al. 1979).
5. Average of 11 topaz rhyolites (Christiansen et al. 1984).
* FeTotal reported as Fe203.
sium feldspar, and quartz. Micas (biotite, muscovite, or lithiumphengite) occur as phenocrysts and in the groundmass along with
topaz. Accessory minerals include fluorite, gamet, zircon, Fe-Ti
oxides, columbite-tantalite, cassiterite, Li-phosphates, pyrite, and
sometimes tourmaline. The relative importance of magmatic versus vapor-phase crystallization in developing this mineralogy is
unclear; but the relatively high Rb/Nb ratios in ongonite (Figure
46) suggest that columbite may be a fractionating magmatic
phase. Kovalenko and Kovalenko (1976) also regard topaz and
mica to be magmatic, but the samples do not appear to be glasses.
The average chemical composition of four ongonites is
shown in Tables 15 and 16. Although there are some obvious
differences between topaz rhyolites and these ongonites (ongonites have higher Al and P and lower Si), we feel that their
similarities are greater. Figures 41 and 46 show the overall similarity in the chemical features of ongonites and topaz rhyolites
from the western United States. Ongonites are markedly enriched
in F, Li, Rb, Cs, Nb, Ta, Be, and other incompatible lithophile
elements and are depleted in Zr, Ba, Sr, and Eu just like the
fluorine-rich rhyolites described here. Kovalenko et al. (1983)
published the REE concentrations of one ongonite from Mongolia (Figure 47) that shows a familiar negative Eu anomaly and
extreme HREE enrichment, so much so that LaN/YbN is less
than one. Overall, the REE are depleted when compared with
topaz rhyolites. Based on phenocryst/matrix partition coefficients
and .experimental studies; Kovalenko (1977), Kovalenko et al.
(1978), and Antipin et al. (1980a,b) have suggested that these
extreme geochemical features are the result of protracted frac-
tional crystallization of crustally-derived magmas with 0.2. to
0.5% F. Expansion of the stability field of quartz by elevated
fluorine contents leads to Na and Al enrichment with Si depletion
during crystal fractionation (Manning 1981) as noted for the Spor
Mountain rhyolite, Utah. The simultaneous fractionation of
quartz, feldspars, and REE-rich accessory minerals seems to be
required (Kovalenko et al. 1983). Some fluid-phase transport of
F and fluorophile elements to the upper parts of evolving magma
chambers may have aided in their differentiation. However, to
enrich the melt in these elements by this process would require
the volume of the fluid to exceed the volume of melt by up to
several hundred times. The suggested origin by fractional crystallization is strengthened by the association of ongonites with large
granitic massifs from which they may have differentiated. Their
spatial and temporal association with basaltic lavas and a description of a mixed (basalt-ongonite) magma (Kovalenko et al. 1975)
suggest that ongonites, like topaz rhyolites, are products ofbimodal magmatic processes. Regarded as the volcanic analogs of Li-F
rare-metal granites, ongonites are associated with Wand other
types of rare-metal mineralization (Mo, Li, and F).
PETROGENETIC MODEL FOR TOPAZ RHYOLITES
Based on the information reviewed here, Christiansen et al.
(1983a) have formulated a petrogenetic model that accounts for
the principal features of topaz rhyolites. We summarize this
model in the context of the nature and composition of the crustal
source of topaz rhyolites; the net power input, as represented by
70
TABLE 16. AVERAGE TRACE ELEMENT CONCENTRATIONS IN
ONGONITES FROM CENTRAL ASIA
1
Ongon
Mongolia
2
3
Baga-Gazryan Arybulak
Mongolia
USSR
4
Teeg-Uula
Mongolia
5
Spor Mtn.
Utah
Li
Rb
Cs
1670
1876
121
186
842
6
417
1024
71
340
975
32
80
971
56
Be
Pb
Nb
19
41
69
2
35
56
20
50
55
25
147
63
40
120
Ta
Zr
Hf
67
66
11
22
78
11
24
83
6
30
138
9
26
107
7
Mo
Sn
Ba
Sr
101
166
115
1
71
69
32
2
37
25
20
30
62
47
6
Note: All analyses in ppm.
1. Average of 53 analyses of the Amazonitov dike (Kovalenko
and Kovlenko 1976).
2. Average of 6 ongonites (Kovalenko and Kovalenko 1976).
3. Average of 9 ongonites (Antipin et al. 1980).
4. Average of 3 volcanic ongonites (Kovalenko et al. 1979).
5. Average of 2 topaz rhyolite vitrophyres (Christiansen et
al. 1984).
the flux of mafic magma, to the base of the continental crust; and
Ongonite
the nature and magnitude of stress in the lithosphere. According
100
to Hildreth (1981), these are the principal controls on the nature
of continental igneous rock associations and their eruption styles.
60
The distribution of topaz rhyolites in the western United
States points strongly to the importance of a magmatic componen.t derived from the Precambrian craton. of North America. The ~
10
notibn that a distinctive crustal reservoir is the source of the -6c0
F-enrichment found in topaz rhyolites is supported by the distri- 0
6
bution of fluorite deposits (Eaton 1984b) and F-rich granitoids
(Christiansen and Lee 1985) in. the western United States as C.E
described above. In the absence of Precambrian crust in. the en
northwestern part of the Great Basin., topaz rhyolites are not
generated. Instead, a sUbtly different bimodal rhyolite, emplaced
0.6
in small extrusive domes with high Na/K ratios and with lower
concentrations of incompatible trace elements, is widely distributed in. western Nevada and eastern Oregon (Wilson et al. 1983;
0.1
E. H. Christiansen, in. preparation).
La Ce
Nd
8m Eu Gd
Dy Ho
Yb Lu
In spite of this inferred crustal origin, topaz rhyolites are not
Figure 47. Rare earth element pattern for a Mesozoic ongonite from
evolved volcanic equivalents of S-type granites derived by partial Mongolia as reported by Kovalenko et al. (1983).
melting of pelitic metasedimentary rocks. The rhyolites possess
distinctly lower initial 87Sr/86Sr, 207Pb/ 204Pb, and 180/ 160 ratios than the S-type granites of the western United States (Wilson equilibration with metasedimentary graphite. Likewise, wholeet al. 1983; Lee et al. 1981; Farmer and DePaolo 1983, 1984). rocks are not strongly peraluminous and in many cases are metaThe compositions of biotites from topaz rhyolites are distinctly luminous. Equilibration of igneous melts with muscovite or
less aluminoUS than those in. muscovite- or garnet-bearing S-type aluminosilicates produces liquids with 3 to 8% normative corungranites. The relatively oxidized conditions under which some dum (Thompson and Tracy 1977; Clemens and Wall 1981). In
topaz rhyolites crystallized (QFM or greater) is inconsistent with addition, the relatively high temperatures of some lavas (up to
~
Ql
11I
L.-~--'-----'-----'-_L.-~--'-----'-----'_-'----'-----'-----'-_L.-~...J
Topaz Rhyolites
71
850°C) and their rise to shallow crustal levels suggest that mus- melting of a low Rb (30-50 ppm) and U (1.5 ppm) protolith will
produce magmas that could fractionate toward compositions typcovite decomposition was not involved in their genesis.
The low to moderate Sr-isotope ratios of topaz rhyolites ical of topaz rhyolites such as the Spor Mountain rhyolite. Such
suggest that their protoliths had Rb/Sr ratios of 0.04 to 0.08. small proportions of partial melting are a natural consequence of
These are relatively low ratios for a source in the continental water-undersaturated melting of high-grade metamorphic rocks
crust, which is typified by Rb/Sr ratios in excess of 0.2 (Taylor such as granulites. The small degree of melting required by in1964). However, such low ratios are typical of granulitic terranes compatible element enrichments could occur with the complete
that experienced Rb depletion during metamorphism and/or ana- decomposition of less than about 10% biotite at lower crustal
texis. This sort of protolith is also consistent with the oxygen pressures (cf. Clemens 1984; Burnham 1979). The decomposition
isotope ratios, but sparse Pb isotope ratios suggest that the Th/U of biotite and its replacement with residual pyroxene (plus melt)
ratio of the protolith must have been "normal" rather than high, would also lower the bulk partition coefficient between rhyolitic
as is found in many granulitic terranes. Small amounts of upper melt and restite enhancing Rb enrichment in the melt. Thus it
crustal contamination are suggested by the high initial Sr-isotope appears that the concentrations of Rb, U, and by analogy other
ratios found at some complexes (e.g. the Thomas Range and trace elements enriched in topaz rhyolites, need not be higher
Nathrop) and suggest caution in attributing measured isotopic than those found in average continental crust. The inferred concentrations are in fact significantly lower than average for contivalues to magmatic sources.
Another important indication of a high-grade metamorphic nental crust. Likewise, there is no requirement that the crustal
protolith for topaz rhyolites are their elevated F concentrations sources of topaz rhyolites contained anomalously low concentraand low F/Cl ratios. Hydrous minerals from high-grade meta- tions of Sr, Ba, Eu, Ni, and other compatible elements. Instead,
morphic rocks are F-rich. As shown by Holloway (1977; see also the small degrees of partial melting in the lower crust, followed
Holloway and Ford 1975; and Manning and Pichavant 1983), by substantial crystal fractionation (of sanidine, quartz, plagiohigh F/(F+OH} ratios increase the thermal stability of biotite and clase, biotite, Fe-Ti oxides, and accessories) enroute to the surface
amphibole. Others have shown that F/(F+OH) ratios in hydrous and in small magma chambers have produced the characteristic
mafic silicates increase with increasing metamorphic grade, ex- compositional features of topaz rhyolites.
The proposed granulitic nature of their protoliths, their extending to granulite facies (Fillippov et aI. 1974; Janardhan et aI.
1982; E. R. Padovani, oral communication, 1984) or to the onset tensional tectonic setting, and their geochemical features imply
of melting (White 1966). Thus, although the absolute amount of that topaz rhyolites may be the extrusive equivalents of A-type or
biotite may decrease with increasing grade of metamorphism, it anorogenic granites (Loiselle and Wones 1979; Collins et aI.
probably becomes more F-rich. The decomposition of small 1982). There are two "species" of anorogenic granites that may
amounts of F-rich biotite would therefore produce small amounts have contrasting origins and evolutionary histories. One type is
of aluminous F-rich melt (probably on the order of 0.2 wt% F) metaluminous to slightly peraluminous (analogous to topaz rhyothat could evolve to produce a topaz rhyolite. Such melts are lites) and the other is peralkaline (analogous to peralkaline rhyoprobably less viscous than their dry equivalents (Dingwell et aI. lites). In granitic complexes both types may coexist, one intruding
1985). It is perhaps noteworthy that scapolites from granulite the other or grading into the other (e.g. the Arabian shield, Stuckgrade rocks are Cl-poor relative to those found in amphibolite less et aI. 1982; or the Younger granites of Nigeria, Bowden et aI.
grade metamorphic rocks (Hoefs et al. 1981). A depletion of Cl is 1984). A similar situation is apparent for the comenditic ash-flow
expected in granulites whether they are formed by reaction with a tuffs that preceded the eruption of the topaz rhyolite at Kane
C02-rich fluid with consequent dehydration or by the removal of Springs Wash, Nevada.
Collins et aI. (1982) and Christiansen et aI. (1983a) have
a silicate melt. In either case CI would preferentially partition into
the escaping fluid/melt. Thus, igneous rocks derived from granu- emphasized the role of F and CI in the evolution of A-type
lites would be expected to have high F/CI ratios as found in topaz granites. The fractional crystallization histories of these magmas
rhyolites.
may be controlled in part by their characteristic F/Cl ratios. For
Although consistent with a granulitic source, the require- example, Manning et aI. (1980) have suggested that F and Al
ment that the sources of topaz rhyolites have relatively low have a strong affinity in granitic melts-so much so that Al is
Rb/Sr ratios is in sharp contrast to the remarkably high Rb/Sr removed from tetrahedral coordination in the aluminosilicate
ratios of topaz rhyolites themselves. By analogy with high-grade framework of the melt and placed in interstitial sites in octahedral
metamorphic rocks, the presumed lower crustal protoliths should coordination. This effect may result in the lowering of the activity
also be depleted in other elements characteristically enriched in of aluminum in the melt and coexisting minerals maintaining an
topaz rhyolites such as U, Th, K, Cs, Li, Be, Nb, Ta, and Y aluminous composition throughout the fractionation history of
(Collerson and Fryer 1978; Sheraton et al 1984; Condie et aI. the magma. In contrast, Cl-rich magmas may experience en1982). This "dilemma" can be resolved if the degree of partial hanced Ca-plagioclase fractionation as a result of Na-CI commelting that produces topaz rhyolites is low. Using estimated bulk plexes in the melt. This process could lead to the production of
partition coefficients for granulitic restite, Christiansen et aI. peralkaline rhyolites through the plagioclase effect of Bowen
(1983b) suggest that values of approximately 10% batch partial (1928). Any process such as volatile escape, which would signifi-
72
Rhyolite Dome
Mafic Lava
Ash-flow Sheet
Stock
)
Caldera
System
Crystallizing Magma
-
(;)
Pre-eruption
Cumulates
~
, \ , ' . .,\.
Exten.sion
or
Shear
Partial
) .. -
~(Q.,':
Melting \
Crust
Mantle
Ponding,
Differentiation
Contamination
f
a
Chamber
i~( \
I', '\
/
"
I
\~eltlng a d Mixing
./' ." '9 9" . . . i
~ . . ' ,'9' 'I t . . .
,
..".
~
.............
Mantle-derived
Basaltic Magma
b
Rhyolite Dome
Ash-flow Sheet
Variably
Cont'aminated
Basalt
,
Ca.ldera-re la ted
Pluton
Crust
Mantle
c
. I \\\
/j.' .. \' \\
I \
/..
.
\
.\
Crust
Mantle
d
Alkaline Lavas
AI.\\1\
\
Silicic
Hybrid Magma
«F/CI)
Cry s ta lliz a t ion,
Degassing,
Contamination of Basaltic Magma
I
Stratovolcano
Rhyolite
Crust
Mantle
Residual
Reworked,,/
Crust
/
Dike Injection
/ '
Basaltic
Magma
t
\ I ,'. '\,
....
:>' ~
I
.
Gabbro
Continued(?) Mafic Magma Input
Figure 48. Hypothetical cross-sections of magma systems (modified in
large part from Hildreth 1981), showing a variety of environments in
which topaz rhyolites are produced. The relative sizes of volcanic edifices
are exaggerated. a) The intrusion of hot mantle-derived basalt into the
continental crust may result in partial melting of felsic granulites in the
lower to middle crust upon decomposition of hydrous minerals. Extensional tectonism and a diffuse focus of dike injection (caused perhaps by
distributed extension) favor the separate rise and eruption of silicic and
mafic magmas. Some silicic magmas may accumulate in small high-level
chambers and experience wall-crystallization (small arrows show the
direction of movement for the crystallization front) and vertical stratification (e.g. Sheeprock granite pluton of west-central Utah; Christiansen et
al. 1983b). Periodic eruptions from the top of such an evolving chamber
may produce large fields of rhyolite domes (e.g. Thomas Range, Utah)
contemporaneous with variably fractionated and/or contaminated mafic
lavas. Some rhyolites fractionate on their passage through the cooler
crust obviating the requirement for their eruption from a sizable shallow
magma reservoir. b) Where the zone of mafic magma injection is wellfocused or the rate of injection is high, hybridization of mantle and crust
materials could be enhanced. Mixing might be unavoidable in such an
environment. The magmas produced could fractionate along a calcalkaline basalt-andesite-dacite (BAD) trend in a shallow magma reservoir feeding a stratocone (gradational to larger caldera-related systems
described below). Alkalic basalts and trachytes also appear to be common in this sort of environment. Examples include the Mt. Taylor volcanic field, New Mexico, and the volcanic systems at Silver Cliff/Rosita,
Colorado. Hybridization should be limited on the flanks of the thermal
focus and partial melting of the crust in these areas could lead to the
eruption of topaz-bearing rhyolite lavas under favorable stress regimes
(e.g. Grants Ridge). Similar high-silica rhyolites may be produced in
advance of crustal penetration by mafic magmas, or after decline of the
mafic magma input when crustal temperatures were still high but opportunities for mixing were small. c) In a variety oftectonic environments,
large collapse caldera systems may develop as the result of the sustained
injection or ponding of mafic magma in the crust. Fractionation of mafic
magmas or hybridization of mafic and silicic crustal melts may have
preceded diapiric(?) separation of moderately silicic magmas to shallower levels. A strongly modified residual crust composed of restite
phases remains in the lower crust (perhaps bearing a mixed mantle and
crustal isotopic signature). Residual hydrous phases should be F-rich.
Decomposition of these phases during a later heating event (represented
by late basalts) could produce the parental magmas for topaz rhyolites
Topaz Rhyolites
cantly alter the F/Cl ratio of the melt, might change the fractionation path followed by the remaining magma and explain the
association of both types of anorogenic magmas within one igneous complex.
In their major and trace element compositions, topaz rhyolites are similar in most respects to the aluminous A-type granites
of southeastern Australia. Notable exceptions are that topaz rhyolites are not LREE rich (125 to 200 ppm Ce versus less than 100
ppm in topaz rhyolites) or Sc rich (greater than 10 versus less
than 3 ppm for topaz rhyolites), but topaz rhyolites from the
western United States contain higher concentrations ofNb (20 to
30 ppm versus 30 to 120 ppm in topaz rhyolites). These differences may relate to higher temperatures inferred for Australian
examples (most are hypersolvus one-feldspar granites with relatively high Zr concentrations) or to differences in the composition
of their crustal sources. The A-type granites of Australia, the
Younger granites of Nigeria, and the Precambrian anorogenic
granites of the southwestern United States (Anderson 1983) are
also F-rich like topaz rhyolites. Subalkaline A-types generally
contain Fe-rich biotite and/or amphibole; fluorite is a common
phase as well.
Anorogenic granites are thought to result from differentiation of variably contaminated alkali basalts (Loiselle and Wones
1979) or from small degrees of partial melting of "residual"
crustal materials from which earlier water-rich magmas had been
removed during granulite-grade metamorphism. It is this later
model that we prefer for the origin of aluminous anorogenic
rhyolites/granites, but it is difficult to imagine how Cl-rich peralkaline magmas could come from granulites without the introduction of CI (and possibly other volatiles) from another source
(presumably mantle-derived basalts). As pointed out by Collins et
al. (1982) and Christiansen et al. (1983a), the protoliths of aluminous anorogenic granites probably consist of felsic granulites
with potassium feldspar, plagioclase, clinopyroxene and orthopyroxene (after decomposition of biotite, which we presume to
be consumed by the melt forming reaction) and quartz. The
which could then erupt through an olcier caldera-related magma system
(e.g. Kane Springs Wash, Nevada, or SW Colorado). Alternatively,
residual pockets of silicic melt might be retained in the lower crust to rise
and erupt slightly later. Their separation and rise could be induced by a
change in stress orientations (e.g. Colorado topaz rhyolites follow voluminous calc-alkaline magmatism during transition to extensional tectonism
and development of the Rio Grande rift) or regional adjustments to the
redistribution of mass in the lithosphere following the development of
granitic batholiths. d) An alternative explanation for the association of
topaz rhyolites with silicic caldera systems holds that topaz rhyolites are
produced by fractionation of the residue left in the magma chamber after
eruption. Such residues might have the high F/Cl ratios typical of topaz
rhyolites if de-volatitization during an earlier eruption effectively extracted CI in preference to F from a portion of the magma remaining in
chamber. Such a process would require the streaming of substantial
amounts of water vapor through the unerupted portion of the magma
chamber. Examples that could be studied with this process in mind
include the Kane Springs Wash, Nevada, and Black Range, New Mexico, rhyolites that are intimately related to caldera systems.
73
experimental studies of Naney (1983) on a synthetic granite and
granodiorite suggest that 10w-SiOz rhyolitic melts would coexist
with this biotite-free phase assemblage at 850 to 900°C and 8 kb.
At lower pressures the requisite temperature is also lower. Residual accessory minerals are probably zircon and apatite, as indicated by their low solubilities in granitic melts (summarized in
Watson and Harrison 1.984) and by the low concentrations of Zr
and P in aluminous anorogenic rocks. Magnetite, ilmenite, or
titanite may also be important residual phases depending on f0 2•
Monazite or some other REE-rich phase is probably residual to
the melting process and holds REE content of the melt to acceptable levels.
In short, we suggest that the most important component in
topaz rhyolites is derived from felsic granulites of the lower or
middle crust. According to this model, high heat flow, resulting
from the emplacement of mafic magmas in or at the base of the
crust, elevated temperatures sufficiently to produce the decomposition of hydrous silicates. More mafic magmatism of a variety of
types is typically associated with the eruption of topaz rhyolites
(Table 10). In Figure 48 we have illustrated some of the geologic
environments in which topaz rhyolites occur.
Topaz rhyolites are commonly found in extensional environments in association with contemporaneous basalts (Figure
48a). Because of density contrasts, hot, mantle-derived basaltic
magma may pond at the base of the crust where it differentiates
by· fractional crystallization and assimilation of crustal materials.
Partial melting of continental crust occurs when temperatures
become high enough to induce the breakdown of hydrous minerals such as biotite. Small quantities of rhyolitic partial melt would
be formed. The character of the erupted mafic magma depends
on its original composition and upon the extent of fractionation
andinteraction with crust materials. In an extensional stress field,
these buoyant melts could rise, fractionate, and erupt to produce
topaz rhyolites which are coeval with variably contaminated (potassic) basalts-a typical bimodal suite (Christiansen and Lipman 1972). Mixing of the contrasting magma types is inhibited
by the efficient separation of the silicic magma in an extensional
setting. Subsequent fractionation in a shallow magma chamber of
moderate size is indicated for some topaz rhyolites by the eruption of moderate volumes of rhyolites over short periods of time
and by the existence of at least one small pluton of Cenozoic
topaz-bearing granite. Eruptions from the tops of vertically zoned
magma chambers may explain the strongly fractionated character
of topaz rhyolites. Many other rhyolite domes may not require
shallow magma chambers because fractionation would probably
occur enroute to the surface.
In an environment where extension is less pronounced or
where the flux of mantle-derived basalt is strongly focused, hybridization of mafic and crust-derived silicic melt might be more
common (Figure 48b). The magmas rising from this zone could
fractionate to produce basalt-andesite-dacite sequences and stratovolcanoes. On the spatial or temporal flanks of the thermal
"focus," independent batches of silicic partial melts and variably
contaminated and fractionated basalt could develop and erupt, as
74
for example at Grants Ridge, New Mexico, which is related to the
Mt. Taylor volcanic field.
Some topaz rhyolites are erupted during the development of
overlapping caldera cycles. This has been interpreted to be the
case for the Mexican topaz rhyolites and for the rhyolites in New
Mexico's Black Range. The eruption of voluminous silicic ashflow tuffs and the formation of collapse calderas suggest the
existence of large shallow-level magma chambers. In examples
from the United States, topaz rhyolites are commonly erupted
after the latest ash-flow tuff in a given area. We suggest two
possible alternatives for the production of topaz rhyolites in this
environment. Our preferred explanation calls on the late rise of
small pockets of melt trapped in the reworked, residual crust, or
the later generation of small volumes of partial melt of the residual crust (Figure 48c). As described by Hildreth (1981), this
modified, perhaps granulitic, crust is thought to have developed
as the result of the injection of mantle-derived magma into the
continental crust and the subsequent extraction of silicic melts
that coalesce to form a large magma body. This model is consis-
i
tent with the coincident rise of mafic and rhyolitic magmas
through the sub-caldera magma chamber, indicating that the
magma was solidified and susceptible to brittle fracture and the
propagation of dikes. An alternative explanation for the occurrence of topaz rhyolites in caldera settings, which might apply to
the Kane Springs Wash caldera, invokes fractional crystallization
of the unerupted portion of an ash-flow-producing magma
chamber (Figure' 48d). The ash flows and the F-rich rhyolite
lavas are seen as being co-genetic in a partially open magma
system. The volatile saturation and eruption of the early magma
are critical for the development of the high FICI ratios observed
in topaz rhyolites relative to earlier magmas. The melt's preferentialloss of CI relative to F (Burnham 1979) during volatile exsolution associated with a large plinian eruption might be able to
produce this change. Subsequent fractional crystallization could
elevate F concentrations to the levels required for the formation
of topaz during post-eruption devitrification and vapor-phase alteration of the evolved residual magma.
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The Geology and Geochemistry of Cenozoic Topaz Rhyolites from

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