JOURNAL OF PETROLOGY
VOLUME 40
NUMBER 12
PAGES 1771–1802
1999
Petrogenesis of the 1·43 Ga Sherman
Batholith, SE Wyoming, USA: a Reduced,
Rapakivi-type Anorogenic Granite
C. D. FROST∗, B. R. FROST, K. R. CHAMBERLAIN AND
B. R. EDWARDS†
DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF WYOMING, LARAMIE, WY 82071, USA
RECEIVED OCTOBER 13, 1998; REVISED TYPESCRIPT ACCEPTED MAY 28, 1999
The 1·43 Ga Sherman batholith, southeastern Wyoming, USA,
shows extreme A-type petrochemical characteristics compared with
other Mid-Proterozoic granite batholiths of North America. It
consists of: (1) the Sherman granite, a coarse-grained biotite
hornblende granite that locally contains fayalite and pyroxenes; (2)
the Lincoln granite, a medium-grained biotite granite; (3) a porphyritic biotite hornblende granite that probably formed by interaction
of granitic and mafic magmas; and (4) iron-enriched mafic dikes
and pods. The ilmenite-series, metaluminous Sherman granite
exhibits extreme values of FeOt/(FeOt + MgO) and is rich in
K, REE, Nb and Y. It crystallized at temperatures exceeding
900°C and a pressure of ~2·5 kbar, with water activity of 0·7
and Dlog fO2 of –0·1 to –0·5. The Lincoln granite, which is
peraluminous and has less extreme A-type geochemical characteristics,
crystallized at temperatures as low as 750°C and Dlog fO2 of
around 0·5 units above FMQ (fayalite–magnetite–quartz). The
rocks of the Sherman batholith are chemically equivalent to lavas
from the Yellowstone hotspot. Like the Yellowstone magmas, the
Sherman batholith probably originated by partial melting of underplated, mantle-derived mafic rocks.
Granites and rhyolites with high K contents and extreme
Fe enrichment are a distinctive rock type of problematic
origin that are found throughout the Proterozoic ‘anorogenic’ granite provinces of the southwestern USA, Adirondacks, eastern Canada, southern Greenland and the
Fennoscandian Shield (Anderson, 1983). Phanerozoic
examples include Paleozoic granites of southeastern Australia (Collins et al., 1982) and of Corsica (Poitrasson et
al., 1995), and the Mesozoic Nigerian Younger granites
( Jacobsen et al., 1958) and the White Mountain Magma
Series (Foland & Allen, 1991). ‘Anorogenic granite’ refers
to plutons that are not associated with compressional
structures, and that were emplaced long after any known
orogenic event. Anorogenic, or ‘A-type’ granites (Loiselle
& Wones, 1979) are characterized by low H2O and O2
fugacities along with high FeOt/(FeOt + MgO), K2O/
Na2O and K2O contents. A-type granites typically are rich
in incompatible elements, including rare earth elements
(REE), Zr, Nb and Ta, but poor in Co, Sc, Cr, Ni, Ba,
Sr, and Eu. Minimum magma liquidus temperatures are
900–1000°C (Clemens et al., 1986; Creaser & White,
1991). Some workers suggest these rocks are derived
from melting of tonalitic or more felsic crust (Anderson
& Cullers, 1978; Collins et al., 1982; Clemens et al., 1986;
Creaser et al., 1991), whereas others relate the granites
to mantle-derived tholeiitic magmas ( Jacobsen et al.,
1958; Turner et al., 1992; Frost & Frost, 1997), or favor
a combination of crustal and mantle sources (Barker et
al., 1975; Foland & Allen, 1991).
The goal of this study is to understand the origin of
the Sherman batholith of southeastern Wyoming. This
intrusion, a ‘reduced, rapakivi-type granite’ (Frost &
Frost, 1997), is composed of ilmenite-series metaluminous
granites that commonly exhibit rapakivi texture. They
∗Corresponding author. Telephone: +1-307-766-6254. Fax: +1-307766-6679. e-mail: [email protected]
†Present address: Department of Geology, Grand Valley State University, Allendale, MI 49401, USA.
 Oxford University Press 1999
KEY WORDS:
A-type; anorogenic; granite; rapakivi; Proterozoic
INTRODUCTION
JOURNAL OF PETROLOGY
VOLUME 40
have among the highest K2O and FeOt/(FeOt + MgO)
contents and the lowest f O2 and f H2O of any MidProterozoic, anorogenic granite in North America (Anderson, 1983; Frost & Frost, 1997), which limits possible
source rocks, and restricts the tectonic environments in
which it could be produced.
Geologic setting of the Sherman batholith
About 1300 km2 of Mid-Proterozoic Sherman batholith
is exposed in the southern Laramie Mountains in southeastern Wyoming and the Front Range of northern
Colorado, with smaller exposures in the southern Medicine Bow Mountains (Fig. 1). The batholith lies near the
Cheyenne belt, a complexly deformed 1·76–1·78 Ga
suture between Proterozoic island arc rocks and the
Archean Wyoming province (Hills & Houston, 1979;
Karlstrom & Houston, 1984; Duebendorfer & Houston,
1987; Chamberlain, 1998). The Cheyenne belt is exposed
in the Medicine Bow Mountains, but in the Laramie
Mountains its inferred trace has been obliterated by the
1·43 Ga Laramie anorthosite complex. Most of the
Sherman batholith cuts Proterozoic igneous and metamorphic rocks that lie south of the inferred trace of
the Cheyenne belt. Only the Mule Creek lobe, the
northeasternmost exposure of the Sherman batholith,
cuts Archean granitic gneiss (Fig. 1).
The Laramie Mountains form an asymmetrical Laramide uplift in which Precambrian rocks have been
thrust eastward over Phanerozoic sedimentary rocks and
have been unconformably overlain in the west by Paleozoic rocks. The Sherman batholith cuts Early Proterozoic supracrustal rocks along its southern margins in
the Laramie Mountains and the Colorado Front Range.
Proterozoic country rocks are also found as a belt on the
east side of the batholith that extends from Granite
Village to Virginia Dale (Fig. 1). Most northern contacts
cut the 1·43 Ga Laramie anorthosite complex (Scoates
& Chamberlain, 1995) or the 1·76 Ga Horse Creek
anorthosite complex (Scoates & Chamberlain, 1997).
LITHOLOGIC UNITS
Our study was primarily of the Sherman Mountains area
of the Sherman batholith, directly east of Laramie. This
small area contains fresh exposures of the rock types
throughout the batholith (Fig. 2). From our work in this
area and reconnaissance mapping elsewhere we have
identified four major units: (1) the Sherman granite, a
coarse-grained biotite hornblende granite that locally
contains fayalite and pyroxenes; (2) the Lincoln granite,
a medium-grained biotite granite; (3) porphyritic biotite
hornblende granite; (4) iron-rich mafic rocks. Rarer are
NUMBER 12
DECEMBER 1999
sodic granitoid rocks, such as the Pole Mountain gneiss,
which we interpret as the oldest unit of the batholith.
Sherman granite
The dominant rock type of the Sherman batholith is
coarse-grained, biotite hornblende granite. This reddish
orange rock commonly weathers deeply to a thick grus.
The Sherman granite is subporphyritic, with a seriate,
hypidiomorphic granular texture. Locally it is an augen
gneiss (Fig. 2), indicating late-stage deformation. Major
phases are microcline, plagioclase, quartz, hornblende,
biotite, and ilmenite. Accessory phases are zircon and
apatite with rarer allanite and fluorite. Augite, pigeonite,
fayalite, and magnetite are found in some samples. The
more hydrous samples contain titanite, produced by
7 CaFeSi2O6 + 5 SiO2 + 3 FeTiO3 + H2O =
in pyroxene
quartz
ilmenite
fluid
2 Ca2Fe5Si8O22(OH)2 + 3 CaTiSiO5.
in hornblende
titanite
Microcline is megacrystic, perthitic, and in places it is
rimmed by plagioclase to create a rapakivi-textured
mantle. The mafic minerals are locally glomerocrystic.
The rock is a granite sensu stricto in the Sherman Mountains area (Table 1; Fig. 3) but in the Virginia Dale area
the Sherman granite consists of quartz syenite and quartz
monzonite in addition to granite (Eggler, 1968).
The Sherman granite locally contains fayalitic olivine
or its alteration products, clinopyroxene, and (in one
sample) pigeonite. Pyroxenes are typically rimmed by
hornblende. Olivine is rimmed by grunerite, which in
turn is rimmed by hornblende; biotite is sparse. On
fresh surfaces, olivine and/or pyroxene-bearing Sherman
granite is green to black. Olivine-bearing samples contain
both orthoclase and microcline; the order–disorder transition in K-feldspar evidently was sluggish in these relatively dry rocks (see Vorma, 1971). The fayalite granite
does not crop out boldly. We were able to sample it only
in blasted roadcuts along I-80. The contact between
fayalite-bearing and fayalite-absent granite appears gradational, and apparently reflects variations in water activity. In many areas the Sherman granite weathers to
dark grus, suggesting fayalite-bearing granite may be
more abundant than we have documented.
Lincoln granite
This medium-grained, red–orange to orange–gray biotite
granite was named after the monument that marks the
summit of the old Lincoln Highway, US 30 (Edwards,
1993). The Lincoln granite occupies much of the area
directly south of the summit of the Sherman Mountains,
1772
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Fig. 1. Geologic map of southeastern Wyoming and a portion of northern Colorado showing the extent of the Sherman batholith. Inset shows
location of study area along the SE margin of the Wyoming province (WY, west of the Superior province (SP). The box outlines the Pole
Mountain area shown in detail in Fig. 2. Open circles give locations of samples that occur outside the area of Fig. 2.
where it crops out as sub-horizontal sheets (Fig. 2). North
of the Sherman Mountains, the Lincoln granite caps hills
and knolls. The granite also occurs as dikes, in which it
is locally commingled with monzodiorite. It is also found
as inclusions in the Sherman granite. Lineated Lincoln
granite occupies a small area at the northwestern end of
the Sherman Mountains, suggesting that the unit was
emplaced throughout the history of the batholith. Houston & Marlatt (1997) equated medium-grained to porphyritic facies in the eastern portion of the batholith to
the Lincoln granite of Edwards (1993). Smith (1977)
described a medium-grained granite in the Mule Creek
lobe of the Sherman batholith near Iron Mountain,
which our reconnaissance work indicates is the Lincoln
granite.
The Lincoln granite is composed of quartz, plagioclase,
microcline, perthite, biotite, apatite, zircon, and locally
traces of hornblende, ilmenite, and fluorite. It contains
more modal quartz than does the Sherman granite
(Table 1; Fig. 3). The rock is generally equigranular,
with an allotriomorphic granular texture. Some samples
display isolated alkali feldspar megacrysts that rarely
make up more than 1% of the rock.
Porphyritic granite
Orange–gray granite with 1–2 cm, orange–pink alkali
feldspar phenocrysts is most abundant north of highway
210 in the Pole Mountain area. This unit was described
by Harrison (1951) as the dominant constituent of the
central part of the batholith, but also resembles the Inner
and Outer Cap Rock quartz monzonite of Virginia
Dale (Eggler, 1968). The porphyritic granite commonly
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JOURNAL OF PETROLOGY
VOLUME 40
NUMBER 12
DECEMBER 1999
Fig. 2. Geologic map of the Sherman Mountains showing the extent of various intrusive phases of the Sherman batholith. Open circles give
the location of samples analyzed in this study.
contains oblate mafic enclaves and schlieren, and xenoliths of Sherman granite, and shows gradational contacts
with monzodiorite and with Lincoln granite. However,
most contacts between the porphyritic granite and the
Sherman granite are sharp.
The major phases in the porphyritic granite are perthitic microcline, plagioclase, quartz, biotite, and hornblende. Titanite, ilmenite, apatite, and zircon are minor
phases. Textures are porphyritic and hypidiomorphic
granular, with megacrystic alkali feldspar. Modal compositions of porphyritic granite samples overlap those of
the Sherman and Lincoln granites, but Eggler (1968)
showed that the porphyritic granite suite displays greater
proportions of plagioclase (Table 1; Fig. 3).
Mafic rocks
Mafic rocks constitute less than 1% of the total area of
the Sherman batholith. In the Virginia Dale area, the
mafic rocks are gabbroic (Vasek & Kolker, 1999), whereas
in the Sherman Mountains area ferrodiorite, monzonite
and monzodiorite are present. Ferrodiorite sampled near
the summit of I-80 contains plagioclase, pigeonite, magnetite and ilmenite, quartz, pigeonite and augite, biotite,
and hornblende, thus resembling ferrodiorites in the
Laramie anorthosite complex (Mitchell et al., 1996). Contacts are poorly exposed. Olivine- and pyroxene-bearing
monzonite from the same locality occurs as a 100 m
diameter, dark bluish purple enclave within Sherman
granite. This rock contains fayalitic olivine, ferro-
1774
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Table 1: Modal percentages of alkali
feldspar (A), plagioclase (P) and
quartz (Q) determined for representative
samples of granitic rocks of the
Sherman batholith
Sample number
A
P
Q
Sherman granite
91SMW28
0·55
0·22
0·23
91PH1
0·54
0·19
0·27
90PH2
0·46
0·25
0·29
91SMW5
0·50
0·29
0·21
90DC2
0·37
0·36
0·27
0·44
0·29
0·27
Porphyritic granite
91PH3
91SMW7
0·47
0·28
0·25
90SME2
0·34
0·32
0·34
90SMW19
0·45
0·32
0·23
90PH4
0·34
0·42
0·24
91SMW13
0·39
0·30
0·31
90SMW1
0·40
0·29
0·31
91SMW3
0·41
0·29
0·30
91PH6
0·40
0·31
0·29
91SMW9
0·37
0·35
0·28
91SMW27
0·28
0·64
0·09
97SMW1
0·28
0·58
0·14
91SMW4
0·10
0·49
0·41
Fig. 3. Modal composition of rocks from the Sherman batholith in
the Sherman Mountains area, compared with modes of samples from
the Virginia Dale area (Eggler, 1968).
whereas other contacts are gradational with porphyritic
granites and lobate–cuspate with the Lincoln granite.
Monzodiorite locally contains enclaves of Sherman granite. We have observed commingling and hybridization
relations between monzodiorite and granite within the
Sherman batholith immediately south of the Sherman
Mountains. Similar relationships have been reported in
Virginia Dale (Eggler, 1968; Vasek & Kolker, 1999) and
near Granite Village (Houston & Marlatt, 1997).
Lincoln granite
Sodic rocks
Sodic rocks of the Sherman batholith
Modal proportions determined from stained slabs; a minimum of 400 points counted.
hedenbergite, alkali feldspar, plagioclase, amphibole, biotite, magnetite and ilmenite, and is similar to monzonites
of the Maloin Ranch and Red Mountain plutons, Laramie anorthosite complex (Kolker & Lindsley, 1989;
Anderson, 1995). It is less siliceous yet has a higher FeOt/
(FeOt + MgO) ratio than the Sherman granite.
The most common mafic rock type in the Sherman
batholith is monzodiorite. Monzodiorite samples contain
plagioclase, hornblende, biotite, quartz, and orthoclase,
with minor titanite, magnetite and ilmenite, apatite, and
zircon. Monzodioritic dikes and pods commonly contain
irregularly distributed alkali feldspar megacrysts that display thin rims of plagioclase, along with plagioclase
xenocrysts, and/or rounded quartz grains rimmed by
hornblende. Contact relations vary. In some areas monzodiorite bodies sharply crosscut the Lincoln granite,
Gray to pink hornblende biotite gneiss (the Pole Mountain
gneiss, Harrison, 1951) makes up most of the crest of the
Sherman Mountains (Fig. 2). It contains plagioclase,
quartz, biotite, sodic hornblende and microcline, and
consists of granodiorite and quartz monzodiorite (Fig. 3).
It is more sodic than the other units of the Sherman
batholith (see geochemistry section below). Foliations of
the gneiss vary in orientation, and feldspars show no
evidence of subsolidus deformation; we therefore interpret
the foliations to be magmatic. The gneiss encloses xenoliths of amphibolite. Both the Lincoln and porphyritic
granites contain enclaves of the Pole Mountain gneiss.
Another member of the sodic series forms an isolated
outcrop along I-80 (locality 91SMW27). This quartz
monzodiorite looks like a whiter example of Sherman
granite and is undeformed. However, it contains sodic
hornblende and much less quartz and alkali feldspar than
does the typical Sherman granite, and chemically appears
to be closely related to the Pole Mountain gneiss. We
interpret it as an enclave of the older, sodic rock series
included in the Sherman granite.
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JOURNAL OF PETROLOGY
VOLUME 40
U–Pb ZIRCON GEOCHRONOLOGY
Previous U–Pb age determinations of zircon from the
Sherman batholith exist only for porphyritic granite from
a drillhole ~5 km north of Buford. Analyses of multiple
fractions of zircon yielded Pb/Pb ages of 1425–1377 Ma
(Aleinikoff, 1983). Interpretation of the zircon data was
complicated by inheritance and lead loss, but Aleinikoff
(1983) favored an age of 1415–1435 Ma.
We selected four samples for U–Pb geochronology
(Table 2): Sherman granite (91PH1), Lincoln granite
(91PH6a), monzonite inclusion (90SMW4), and Pole
Mountain granite gneiss (94SMW2 and 90SMW13).
These samples appear to span the intrusive history of
the Sherman batholith. U–Pb data for sample 91PH1
were obtained from five fractions of zircon, including
two air-abraded fractions (Table 2). On a concordia
diagram, these fractions range from 3 to 35% discordant,
and yield an intercept age of 1433 ± 1·5 Ma (Fig. 4).
Common lead was corrected based upon the Pb isotopic
composition of coexisting feldspars. There is no evidence
of inherited zircon or other complexities in the U–Pb
systematics, which indicates that the intercept age is the
crystallization age of the Sherman granite.
Seven fractions of Lincoln granite sample 91PH6a
were analyzed, including six air-abraded fractions
(Table 2). The air-abraded fractions are nearly concordant, ranging from 0·6 to 6% discordant, whereas
the unabraded fraction is 31% discordant. Together, the
seven fractions yield an intercept age of 1430 ± 2·6 Ma
(Fig. 4). The correction for common lead was made using
the Pb isotopic composition of coexisting feldspar. Zircons
from the Lincoln granite show higher U and Th/U ratios
than zircons from the other samples, and contain the
most common Pb of the samples. Common Pb is associated with the outer portions of the grains, because it
was reduced by air-abrasion.
Monzonite inclusion sample 90SMW4 was dated to
limit the age of the Sherman granite that engulfs this
10 m diameter enclave. U–Pb data for six fractions of
zircon from this sample yield a non-linear array on a
concordia diagram, with a small range in Pb/Pb ages of
1436–1440 Ma (Table 2; Fig. 4). All of the analyses
are concordant or nearly concordant, with a maximum
discordance of 1·4%. The range in Pb/Pb ages reflects
small amounts of inheritance, and the youngest Pb/Pb
age of 1436·3 ± 1·3 Ma is a maximum age for the
monzonite.
Two samples of the Pole Mountain gneiss were analyzed. Sample 94SMW2 was collected from an outcrop
free of enclaves of older country rock. The U–Pb systematics of this sample were reasonably simple. Four
fractions of zircon are nearly concordant, ranging from
1 to 4% discordant on a U–Pb concordia diagram, and
yield a linear array with an upper intercept age of
NUMBER 12
DECEMBER 1999
1437·8 ± 3·2 Ma (Table 2; Fig. 4). This is probably the
crystallization age of the Pole Mountain gneiss. The
second sample was collected from an outcrop with abundant enclaves of older rocks, including amphibolite and
quartzite. Twelve fractions of zircon from this sample
(90SMW13) yield Pb/Pb ages ranging from 1450 to 1689
Ma (Fig. 4). The range indicates varying proportions of
inherited components in the zircon. The zircon fraction
with the youngest Pb/Pb age of 1450 Ma has the least
inheritance. Assuming that no fraction is completely free
of inheritance, the results from this sample are consistent
with those for sample 94SMW2, and imply that the Pole
Mountain gneiss has a maximum age of 1438 Ma, and
that the gneiss contains a component of assimilated older
crust. In summary, the Pole Mountain gneiss is the oldest
unit of the Sherman batholith, with a maximum age of
1439 Ma. The volumetrically dominant lithologies, the
Sherman and Lincoln granites, were emplaced at 1433
± 1·5 and 1431 ± 2·6 Ma, respectively.
GEOCHEMICAL CHARACTERISTICS
Major element geochemical data for the Sherman bathoith are available from several sources: the majority is
from the Pole Mountain area (present study, Tables 3
and 4), but additional analyses exist for rocks from
Virginia Dale (Eggler, 1968), the Granite Village area,
and from Sheep Mountain in the Medicine Bow Mountains (Houston & Marlatt, 1997). Because published
sample descriptions did not allow us to consistently
identify the lithologic units analyzed, we plot only data
from the Pole Mountain area in Figs 5–9.
Most analyses form coherent arrays on major element
Harker diagrams (Fig. 5). The mafic rocks and the
Sherman, porphyritic and Lincoln granites decrease in
TiO2, FeOt, MgO, CaO and P2O5 with increasing SiO2.
Al2O3 and Na2O decrease and increase, respectively, in
the most siliceous samples. K2O generally increases with
silica except for Lincoln granite samples with the highest
SiO2 contents. The Sherman batholith is subalkalic,
and spans the metaluminous–peraluminous boundary
(Fig. 6a). In contrast to calc-alkalic rocks, they exhibit
extreme iron enrichment (Fig. 6b), and have large K2O
values (Fig. 5). Rocks of the Sherman batholith contain
abundant large ion lithophile elements (LILE; Rb, Ba,
REE) and high field strength elements (HFSE; Zr, Y, and
Nb), which are typical of anorogenic granites. Analyses of
Sherman batholith rocks lie mostly in the within-plate
granite field of Pearce et al. (1984), although Lincoln
granite samples lie between the within-plate and volcanic
arc fields (Fig. 6c).
The older, volumetrically minor sodic rocks and the
monzonite enclave do not follow the geochemical trends
of the rest of the Sherman batholith. The sodic rocks
1776
Weight
(mg)
U
(ppm)
1777
20·61
233
0·01
0·01
0·03
0·00
0·01
0·04
51
34·7
29·2
41·9
28·9
32·1
0·92
33·21
233
26·6
26·92
264
6·37
29·06
235
82·8
46·09
0·11
19·2
110
2·60
1·32
0·82
1·29
comPb
(ppm)
9·6
33·7
17·8
54·2
Pb
(ppm)
90SMW4 monzonite inclusion
d-2 +100 0·047
204
aa single
No. 2
d-2 +100 0·600
116
aa
d-2 +100 0·108
138
aa single
No. 1
d-3
0·670
169
d-1
0·560
115
d-2 +100 0·940
127
91PH6a Lincoln granite
d-1 single 0·005
363
grain
nm10 aa 0·050
852
pink No. 1
nm10 aa 0·014
977
pink No. 2
m6 aa
0·070
825
small
nm10 aa 0·014
860
pink No. 3
nm10 aa 0·040
311
colorless
nm6 aa
0·040
103
large
91PH1 Sherman granite
d-3
0·126
35
d-1 −200 0·080
201
d-2
0·080
96
d-1 aa
0·173
233
−100 to
+200
d-2 aa
0·190
77
Sample
name
Pb/238U
(rad.)
206
%err
630284
220296
48507
69150
130837
43293
1659
703
602
361
518
418
13·1
11·5
11·9
12·0
11·4
12·9
12·5
8·4
7·9
8·3
8·0
7·5
0·2467
0·2475
0·2492
0·2485
0·2483
0·2494
0·2484
0·2360
0·2353
0·2328
0·2322
0·2302
0·33
0·33
0·37
0·33
0·33
0·33
0·37
0·34
0·34
0·50
0·34
0·43
3·0825
3·0937
3·1151
3·1047
3·0989
3·1129
3·0907
2·9209
2·9103
2·8778
2·8660
2·8358
1·9691
2·3766
1·9723
2·1517
2·7713
1430 ± 2·6 Ma MSWD = 0·85
93
5·9
0·1638 0·63
0·2412
Pb/235U
(rad.)
207
3·0031
9·7
1·5 Ma MSWD = 1·3
7·4
0·1943 0·45
13·0
0·1601 0·33
10·8
0·1743 0·35
11·0
0·2229 0·30
Pb/
208
Pb
(rad.)
206
0·33
10596
1433 ±
172
1479
1230
2420
Pb/
204
Pb
(corr.)
206
Corrected atomic ratios∗
Table 2: U–Pb zircon data for samples from the Sherman batholith
0·33
0·33
0·37
0·34
0·33
0·34
0·39
0·41
0·40
0·79
0·43
0·54
3·01
0·34
1·69
0·35
0·40
0·30
%err
0·0906
0·0907
0·0907
0·0906
0·0905
0·0905
0·0903
0·0898
0·0897
0·0896
0·0895
0·0893
0·0872
0·0903
0·0887
0·0894
0·0895
0·0902
0·06
0·06
0·06
0·06
0·06
0·07
0·12
0·20
0·19
0·56
0·24
0·30
2·68
0·08
1·49
0·12
0·19
0·12
%err
Pb/206Pb
(rad.)
207
Pb/
U
Age
( Ma)
1421
1425
1434
1431
1430
1436
1430
1366
1362
1349
1346
1336
978
1393
1144
957
1036
1297
238
206
Pb/
U
Age
( Ma)
1428
1431
1436
1434
1432
1436
1430
1387
1385
1376
1373
1365
1105
1408
1236
1106
1166
1348
235
207
Pb/
Pb
Age
( Ma)
1438·9
1439·7
1439·7
1438·3
1436·4
1436·3
1430·9
1420·2
1419·4
1418·0
1415·6
1411·3
1364·5
1431·8
1398·4
1412·1
1415·2
1429·1
206
207
0·98
0·98
0·99
± 1·2
± 1·2
± 1·2
0·95
± 2·3
0·98
0·87
± 3·8
± 1·2
0·88
± 3·6
0·98
0·71
± 11
± 1·2
0·84
± 4·5
0·98
0·84
± 5·6
± 1·3
0·59
0·97
± 1·7
± 52
0·56
0·94
0·88
0·93
Rho
28
2·3
3·6
2·3
±
±
±
±
±
1
1
0
1
1
0
0
4
4
5
5
6
31
3
20
35
29
10
%
discordance
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
1778
86·1
291
252
939
80·5
191
178
686
643
339
71·9
142
91·7
64·1
104
356
253
415
332
487
32·9
98·9
76·6
311
gneiss
157
382
104
26·6
414
105
46·1
Pb
(ppm)
Pb/238U
(rad.)
206
%err
2970
13975
21564
578
2541
8769
665
42904
8355
16831
9905
14623
7170
2433
4220
14·6
16·5
13·7
15·1
13·2
13·3
13·2
10·1
11·1
10·8
14·1
10·3
13·1
12·3
11·3
0·2859
0·2366
0·2650
0·2507
0·2679
0·2146
0·2639
0·2508
0·2482
0·2461
0·2096
0·2516
0·2434
0·2436
0·2471
0·56
0·33
0·31
0·34
0·34
0·33
0·38
0·34
0·38
0·35
0·36
0·36
0·33
0·34
0·55
1437·8 ± 3·2 Ma MSWD = 1·3
3242 14·1
0·2386 0·34
Pb/
208
Pb
(rad.)
206
Pb/235U
4·0820
3·3471
3·7115
3·4979
3·7481
2·8112
3·6249
3·2106
3·1858
3·1628
2·6343
3·2183
3·0349
3·0359
3·0808
2·9696
(rad.)
207
0·65
0·34
0·32
0·37
0·44
0·34
0·47
0·35
0·39
0·36
0·38
0·39
0·34
0·35
0·63
0·35
%err
0·1035
0·1026
0·1016
0·1012
0·1015
0·0950
0·0996
0·0928
0·0931
0·0932
0·0911
0·0928
0·0904
0·0904
0·0904
0·0903
0·31
0·07
0·08
0·14
0·26
0·08
0·25
0·07
0·09
0·07
0·11
0·12
0·07
0·08
0·29
0·08
%err
Pb/206Pb
(rad.)
207
Pb/
U
Age
( Ma)
1621
1369
1516
1442
1530
1253
1510
1443
1429
1418
1227
1447
1404
1406
1423
1379
238
206
Pb/
U
Age
( Ma)
1651
1492
1574
1527
1582
1359
1555
1460
1454
1448
1310
1462
1416
1417
1428
1400
235
207
Pb/
Pb
Age
( Ma)
1688·5
1671·4
1653·0
1646·4
1651·4
1528·4
1617·1
1484·8
1489·8
1492·1
1449·5
1483·3
1435·0
1433·6
1434·7
1431·3
206
207
0·97
0·85
0·93
0·82
0·97
0·98
0·88
± 1·3
± 4·6
± 2·5
± 4·7
± 1·5
± 0·87
± 5·3
0·98
0·97
0·98
± 1·4
± 1·7
± 1·4
0·98
± 1·4
0·96
0·95
0·97
0·89
± 1·6
± 5·5
± 2·3
± 2·3
0·97
Rho
± 1·6
±
5
20
9
14
8
20
7
3
5
6
17
3
2
2
1
4
%
discordance
NUMBER 12
1·64
0·33
0·67
17·86
3·98
0·47
11·55
0·12
0·44
0·35
0·19
0·38
0·63
2·47
0·36
0·83
comPb
(ppm)
Pb/
204
Pb
(corr.)
206
Corrected atomic ratios∗
VOLUME 40
∗Ratios corrected for blank and mass discrimination.
Zircon fraction: d— , nm— represent angles of diamagnetic and paramagnetic susceptibility on a barrier style Frantz separator. aa, air abraded; 100, 200, mesh
size; comPb (ppm) corrected for laboratory blank of 10 pg; % err, 2r in percent; rho, 206Pb/238U vs 207Pb/235U error correlation coefficient. Heavy minerals were
concentrated from crushed samples by standard density and magnetic separation procedures. Concentrated zircon grains varied in color, cores and size. The
grains selected were devoid of inclusions and each fraction was composed of a single population of distinct size and/or color. The grains were cleaned in 2 N
HNO3 before dissolution in concentrated HF and HNO3 in Teflon microbombs following the procedure of Parrish (1987), and converted to chlorides. Aliquots of
sample solutions were spiked with a mixed 235U/208Pb spike. Pb and U were purified on anion exchange columns using HCl chemistry modified from Krogh (1973).
The total Pb blank was decreased significantly during the course of this study from 80 pg to 5 pg. The blank composition is 206Pb/204Pb = 19·09, 207Pb/204Pb =
15·652, 208Pb/204Pb = 38·31. Pb and U samples were loaded onto single rhenium filaments for isotopic analysis by thermal emission mass spectrometry. The H3PO4
silica gel technique was used for Pb. U samples were loaded with H3PO4 and graphite and run as metal ions. The samples were analyzed isotopically at the
University of Wyoming on a multi-collector VG Sector mass spectrometer. Mass discrimination factors for Pb and U were determined by multiple analyses of
NBS SRM 983 and U-500, respectively, and were 0·048 ± 0·06% per a.m.u. for Pb and 0 ± 0·06% per a.m.u. for U. PBDAT (Ludwig, 1988) was used to reduce the
raw mass spectrometer data, correct for blanks, and calculate uncertainties. Ages and uncertainties were calculated by ISOPLOT (Ludwig, 1991), and errors are
quoted at 2r.
90SMW13 Pole Mt
nm10
0·028
d-2 aa
0·010
No. 3
d-1 aa
0·020
nm10 aa 0·014
d-1 aa
0·014
no cores
d-1
0·080
d-1 & d-2 0·029
aa
d-3
0·040
d-2 aa
0·027
No. 1
d-1 aa
0·027
grains
with cores
d-1
0·120
resorbed
grains
d-2 aa
0·007
No. 2
0·041
0·041
0·013
Pole Mt gneiss
0·041
190
94SMW-2
d-4 2
grains
nm1
d-4 1
grain
d-3
U
(ppm)
Weight
(mg)
Sample
name
Table 2: continued
JOURNAL OF PETROLOGY
DECEMBER 1999
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Fig. 4. U–Pb concordia diagrams for zircons from samples of the Sherman batholith.
(crosses in Fig. 5) show total alkali contents identical to
younger units, but contain higher Na2O and lower K2O.
Sodic rocks are also distinguished by much higher Al2O3
contents and lower FeOt and TiO2 contents. Analyses of
the monzonite enclave lie off both these trends.
Mafic rocks
The mafic rocks of the Sherman batholith display high
Ti, P, Fe, Zr, Ba, Ga and Nb contents relative to calcalkalic basalts (Table 3). An MgO-rich group (MgO >
3·0 wt %), which contains biotite as the main mafic
mineral, is found in the western Sherman Mountains.
An MgO-poor group (MgO < 3·0 wt %) consists of
pyroxene-bearing ferrodiorite and monzodiorite with
subequal amounts of hornblende and biotite. The MgOpoor monzodiorites occur as dikes and commingled bodies in the eastern portion of the Sherman Mountains
area.
Sherman granite
The SiO2 contents of Sherman granite samples range
from 64·4 to 71·3 wt %. The Sherman granite has the
highest K2O content of any unit in the batholith, along
with the highest Nb and Y contents (Fig. 6c), as well as
Zr and Ga contents as high as those in the mafic rocks
(Table 3). The light REE (LREE) contents of Sherman
granite sample (Ce = 282 ppm) are the highest of any
of the rocks analyzed from the Pole Mountain area, and
its heavy REE (HREE) contents are exceeded only by
those of the monzodiorite sample (91SMW20; Fig. 7).
TiO2, MnO, P2O5, and Rb are low compared with
porphyritic granites with the same SiO2 contents. For
these elements, the Sherman granites lie slightly below
the trends defined by the monzodiorites, porphyritic
granites, and Lincoln granites (Fig. 5).
In the Sherman Mountains area, all Sherman granite
samples are metaluminous (Fig. 6a). However, those
exposed in the Sheep Mountain area of the Medicine
Bow Mountains are peraluminous (Houston & Marlatt,
1779
14·06
11·22
0·15
3·68
6·46
3·45
2·66
1·64
0·92
Al2O3
FeOt
MnO
MgO
CaO
Na2O
K2 O
P2 O 5
LOI
1780
18
Pb
177
Zn
2
27
Cu
11
39
Cr
U
37
Ni
Th
44
Nb
25
623
Zr
203
3855
Ba
La
1008
Sr
11
2
3
119
24
246
6
19
11
35
599
2406
378
181
68
0·664
13
2
7
142
24
170
15
62
30
35
520
2000
796
88
43
0·720
0·739
6·15
100·11
0·51
1·14
2·44
3·71
6·59
3·88
0·14
11·00
14·92
2·5
6
0
1
20
25
143
19
40
20
13
213
1627
528
22
39
0·790
0·831
5·61
100·00
0·00
0·71
1·47
4·14
6·75
2·36
0·24
11·58
16·34
1·71
54·70
10
2
5
56
21
122
27
53
35
11
240
881
335
60
42
0·761
0·772
5·4
98·36
0·14
0·44
2·1
3·3
6·09
3·24
0·16
10·97
14·29
1·79
55·85
18
2
11
179
25
192
17
7
7
47
734
2568
345
126
75
0·735
0·846
6·82
99·00
1·08
1·14
3·45
3·37
5·03
2·05
0·19
11·25
13·55
1·99
55·90
19
10
19
67
26
138
8
7
8
40
621
1766
269
150
72
0·798
0·851
7·11
98·70
0·14
0·90
3·86
3·25
4·06
1·62
0·16
9·26
13·49
1·60
60·36
21
4
18
151
25
125
8
10
8
37
653
1707
229
95
80
0·850
0·902
8·83
99·31
0·18
0·25
4·77
4·06
2·98
0·69
0·12
6·36
14·68
0·80
64·42
Sherman
px-
90smw9
90smw5
17
5
7
102
24
116
10
7
8
53
515
1197
176
107
114
0·842
0·867
7·92
99·08
0·28
0·26
4·28
3·64
3·05
0·95
0·11
6·21
13·61
0·92
65·80
27
430
1400
185
117
45
0·916
0·914
9·71
100·22
0·47
0·09
5·75
3·96
1·82
0·41
0·07
4·33
14·70
0·42
68·20
Sherman
Sherman px-
90dc2
90ph2
91ph1
90ph1
24
4
20
136
23
78
6
6
4
33
468
1355
182
114
50
0·928
0·905
9·26
99·94
0·41
0·12
5·50
3·76
1·94
0·40
0·07
3·81
14·54
0·44
68·95
20
4
15
80
23
54
3
6
4
23
383
1283
165
162
50
0·930
0·892
9·96
99·17
0·28
0·08
5·99
3·97
1·53
0·34
0·07
2·82
14·69
0·36
69·04
25
6
20
170
26
113
4
27
10
67
572
992
103
174
117
0·941
0·946
9·02
97·69
0·04
5·61
3·41
1·30
0·17
0·06
3·00
13·22
0·27
70·62
21
6
27
128
23
78
2
8
6
53
476
896
124
186
94
0·952
0·907
8·89
99·89
0·93
0·06
5·22
3·67
1·50
0·32
0·06
3·11
13·72
0·33
70·97
63
540
1100
106
185
140
0·916
0·948
9·52
100·39
0·39
0·04
5·76
3·76
1·33
0·20
0·05
3·67
13·60
0·29
71·3
Sherman Sherman Sherman Sherman Sherman
91smw28 91smw5
NUMBER 12
Ga
49
73
Rb
0·693
Y
A/CNK
0·844
7·17
98·8
0·71
1·52
3·38
3·79
5·69
2·29
0·22
2·39
52·94
Fe-mzdi
91smw17 91smw20
Fe-mzdi
Sherman granite
VOLUME 40
(FeOt+MgO)
6·11
0·753
Na2O+K2O
FeOt/
98·44
12·43
2·52
TiO2
Total
13·44
51·68
SiO2
53·28
Mg-mzdi
summ fdi Mg-mzdi
91smw11 91smw30 91VD1
Mg-mzdi
Description:
Fe-mzdi
Sample no.: 90smw17 91smw6
Mafic rocks
Table 3: Major and trace element analyses of the Sherman batholith
JOURNAL OF PETROLOGY
DECEMBER 1999
91smw16
91ph3
90sme2
91smw7
91smw19
91ph2
0·33
0·35
99·21
LOI
Total
1781
10
Ni
Cr
23
146
22
9
24
Ga
La
Th
U
Pb
3
8
Nb
114
42
Zr
Zn
620
Ba
Cu
189
1426
Sr
80
247
Rb
0·909
Y
A/CNK
(FeOt+MgO)
0·884
4·83
K2 O
P2 O 5
8·1
3·27
Na2O
Na2O+K2O
4·73
2·50
CaO
FeOt/
3·23
0·75
MgO
5·39
MnO
24
6
24
117
23
119
5
6
5
37
564
1121
167
205
76
0·941
0·882
7·96
99·07
0·56
0·30
2·26
0·72
0·11
5·74
0·10
FeOt
13·69
0·73
67·35
24
8
26
132
23
99
8
6
4
36
544
1111
166
178
74
0·919
0·886
8·27
99·14
0·79
0·28
4·79
3·48
2·14
0·66
0·08
5·14
13·60
0·70
67·48
23
8
26
107
22
92
6
14
8
35
517
1091
166
207
72
0·902
0·883
8·52
99·72
0·64
0·27
4·97
3·55
2·23
0·66
0·09
5·00
13·77
0·68
67·86
22
11
26
101
22
91
6
11
6
37
532
1097
137
183
69
0·986
0·864
8·02
99·48
1·59
0·27
4·55
3·47
1·73
0·73
0·08
4·66
13·59
0·66
68·14
20
340
910
212
193
68
0·935
0·841
8·12
100·73
0·93
0·20
4·51
3·61
2·34
0·93
0·07
4·93
14·10
0·61
68·5
20
7
41
145
21
62
3
10
7
28
351
1006
202
184
51
0·964
0·854
8·74
99·49
0·58
0·18
5·25
3·49
1·90
0·59
0·05
3·46
14·34
0·49
69·16
18
280
1000
209
197
57
0·975
0·853
8·42
100·10
1·00
0·14
4·82
3·60
1·83
0·62
0·06
3·59
14·10
0·44
69·90
26
6
45
162
24
83
2
12
8
42
452
854
129
202
86
0·963
0·904
8·48
99·42
0·79
0·16
5·18
3·30
1·54
0·43
0·08
4·04
13·33
0·48
70·09
24
6
39
153
21
75
3
8
6
39
406
866
167
196
60
0·982
0·855
8·55
99·42
0·43
0·15
5·12
3·43
1·51
0·57
0·06
3·36
13·68
0·46
70·65
27
260
910
160
220
72
0·975
0·856
8·85
100·46
0·77
0·08
5·15
3·70
1·54
0·54
0·06
3·20
14·10
0·32
71·00
28
11
54
153
23
82
2
6
5
38
385
667
97
244
101
0·974
0·909
8·82
99·66
0·74
0·10
5·25
3·57
1·09
0·33
0·08
3·28
13·18
0·36
71·68
28
5
29
59
16
44
2
6
4
22
190
1015
217
203
16
1·060
0·792
8·4
99·57
0·66
0·06
5·17
3·23
1·23
0·42
0·04
1·60
14·06
0·22
72·88
21
4
19
37
17
19
6
6
2
19
148
986
165
175
28
1·033
0·870
9
99·38
0·46
0·03
5·77
3·23
0·76
0·16
0·03
1·07
13·37
0·11
74·39
13·77
90ph5
Al2O3
91smw2
0·78
90ph4
66·78
90smw20
TiO2
91smw10
SiO2
91smw14
Description:
91smw22
90smw19
porphyritic porphyritic porphyritic porphyritic porphyritic porphyritic porphyritic porphyritic porphyritic porphyritic Buford pg porphyritic Buford pg porphyritic
Sample no.:
Porphyritic granite
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
0·23
0·30
13·90
Al2O3
3·46
5·18
0·06
0·62
Na2O
K2 O
P2 O 5
LOI
1782
2
45
18
71
38
10
30
Cu
Zn
Ga
La
Th
U
Pb
4
16
233
6
20
Nb
Cr
220
Zr
585
118
216
31
1·013
14
220
1300
209
208
30
1·051
0·859
8·78
100·50
0·77
0·05
5·28
3·50
1·07
0·40
31
8
26
59
18
39
2
6
3
15
191
830
146
225
31
1·076
0·856
8·50
99·87
0·89
0·05
5·17
3·33
0·92
0·29
0·05
1·73
13·71
0·20
73·53
Lincoln
91ph6
11
21
110
570
105
235
0·975
0·881
9·47
100·40
0·62
0·02
5·26
4·21
0·79
0·19
0·03
1·40
13·70
0·12
73·70
Lincoln
90smw6
23
31
12
44
29
17
37
2
6
3
22
126
423
94
192
1·028
0·903
8·84
100·00
0·51
0·02
5·35
3·49
0·79
0·15
0·04
1·40
13·33
0·13
74·79
Lincoln
31
26
9
36
50
20
44
3
6
1
27
189
478
114
181
1·012
0·877
8·62
100·90
0·56
0·05
4·81
3·81
0·98
0·25
0·04
1·79
13·41
0·18
75·02
Lincoln
30
8
33
24
19
34
1
6
2
12
103
445
99
215
22
1·043
0·870
8·68
99·97
0·57
0·02
4·97
3·71
0·70
0·18
0·03
1·20
13·30
0·11
75·18
Lincoln
19
9
27
9
18
19
2
5
1
21
110
365
74
204
20
1·112
0·885
8·72
99·64
0·62
0·01
5·46
3·26
0·45
0·11
0·03
0·85
13·45
0·07
75·33
Lincoln
34
12
42
20
18
34
1
6
3
27
113
269
64
237
20
1·047
0·893
8·68
99·87
0·43
0·01
4·69
3·99
0·58
0·13
0·04
1·08
13·29
0·09
75·54
Lincoln
25
6
22
17
18
20
2
6
3
11
71
458
87
202
27
1·050
0·869
8·86
99·77
0·66
0·01
5·39
3·47
0·50
0·13
0·03
0·86
13·08
0·07
75·57
Lincoln
37
13
25
8
21
46
1
6
1
29
121
202
58
248
30
1·002
0·914
8·67
99·94
0·76
0·01
4·30
4·37
0·60
0·10
0·04
1·06
12·96
0·08
75·66
Lincoln
41
12
5
18
152
428
86
281
27
1·043
0·881
8·97
101·19
0·66
0·03
5·21
3·76
0·57
0·20
0·05
1·48
13·41
0·14
75·68
Lincoln
24
10
29
12
20
16
2
6
2
32
104
266
74
236
33
1·052
0·902
8·95
100·07
0·69
0·01
5·04
3·91
0·33
0·10
0·01
0·92
13·14
0·09
75·83
Lincoln
91smw26 91smw12 91smw25 91smw21 91smw23 91smw18 91smw15 91smw13 91ph5
41
9
16
3
25
11
3
5
0
35
92
65
15
325
65
0·977
0·833
9·64
103·10
0·43
0·00
4·89
4·75
0·32
0·04
0·01
0·20
13·37
0·02
76·07
Lincoln
29
6
26
7
20
18
0
6
3
21
74
281
75
211
21
1·008
0·890
8·79
100·09
0·51
0·01
4·95
3·84
0·68
0·09
0·02
0·73
13·01
0·07
76·18
Lincoln
91smw9 91smw8
NUMBER 12
Ni
216
1200
Ba
Rb
Sr
15
197
Y
1·042
A/CNK
0·868
8·96
99·31
0·61
0·05
1·08
0·32
0·04
2·43
14·10
0·26
72·60
Lincoln
90smw3
VOLUME 40
(FeOt+MgO)
8·64
0·866
Na2O+K2O
FeOt/
99·58
5·53
1·12
CaO
Total
3·43
0·39
MgO
2·10
MnO
0·04
2·51
0·04
FeOt
13·77
72·15
72·00
TiO2
Lincoln
Description: Lincoln
SiO2
91smw3
Sample no.: 90smw1
Lincoln granite
Table 3: continued
JOURNAL OF PETROLOGY
DECEMBER 1999
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
batholith. Fourteen of 16 samples are peraluminous. The
normative Ab–An–Or values for Lincoln granite samples
lie along cotectics calculated by Nekvasil & Lindsley
(1990) for the system Ab–An–Or–H2O (Fig. 9), and trend
to higher Na2O and lower K2O values in the most
SiO2-rich samples. Rb/Ba increases dramatically with
increasing SiO2 (Fig. 8), a feature consistent with crystallization of feldspar. Zr, Y and REE contents are lower
than for the other granites in the batholith; these are
probably controlled by the accessory phases zircon, allanite and apatite (Figs 7 and 10). The low Zr, REE and Y
values in the Lincoln granite samples may result because
the Lincoln granite is (1) a partial melt in which the
accessory minerals were left in the residua, or (2) a
differentiate that separated from an assemblage that
contained these accessory minerals (i.e. monzodiorite and
Sherman granite). The lowest REE contents are also
found in the most siliceous units of the zoned monzoniteto-granite Red Mountain pluton of the Laramie anorthosite complex, in which it was demonstrated that REEs
were sequestered in allanite and zircon during differentiation of a monzonitic magma (Anderson, 1995).
Table 3: continued
Older rocks
Sample no.:
90smw4
91smw27
Description:
monzonite
sodic group sodic group sodic group
97smw1
91smw4
SiO2
60·06
62·91
67·66
TiO2
0·83
0·41
0·25
70·89
0·16
Al2O3
16·51
19·56
17·70
15·93
FeOt
7·68
3·20
1·90
1·40
MnO
0·17
0·06
0·04
0·03
MgO
0·38
0·37
0·71
0·58
CaO
3·14
4·36
2·83
2·80
Na2O
4·57
6·61
5·80
6·07
K2 O
5·58
2·28
3·26
1·50
P2O5
0·25
0·19
0·084
0·04
LOI
0·33
0·09
0·24
0·47
Total
99·49
100·04
100·46
99·87
Na2O+K2O
10·15
FeOt/
8·89
9·06
7·57
0·953
0·896
0·728
0·707
0·857
0·920
0·972
0·954
(FeOt+MgO)
A/CNK
Y
37
68
Rb
77
49
8
12
Sr
159
286
1104
432
Ba
4933
324
1636
637
Zr
651
356
141
74
Nb
12
21
Ni
14
4
7
4
Cr
38
6
8
10
Cu
12
6
236
6
Zn
98
74
29
28
Ga
35
34
La
31
67
7
Th
1
11
1
Porphyritic granite
47
Porphyritic granite samples range from 66·8 to 74·4 wt
% SiO2, the largest variation in the granites of the
Sherman batholith. They occupy intermediate major and
trace element compositions between Lincoln, mafic rock
and Sherman units. The REE pattern of porphyritic
granite sample 91SMW2 lies between Sherman sample
91SMW28 and Lincoln sample 91PH6 (Fig. 7), but
is also LREE enriched with a negative Eu anomaly.
Porphyritic granite samples lie between the MgO-poor
monzodiorite and the Lincoln granite (Figs 5 and 10).
These geochemical features, along with petrographic
and field evidence, indicate that porphyritic granite is
probably a product of magma mixing or interaction of
magmas and feldspar-rich crystal mush.
13
16
U
1
2
1
Pb
10
17
9
Samples were analyzed by X-ray fluroescence spectrometry
at the University of Southern California, except sample
97SMW1, which was analyzed by XRAL Activation Services,
Inc.
1997). Although there is a 4 wt % difference in SiO2
content of the least siliceous Sherman granite and the
most siliceous monzodiorite, the two groups are geochemically similar. Both are metaluminous and rich in
Zr, Nb and Y, and of low Rb/Ba (Fig. 8).
MINERAL CHEMISTRY
Mineral compositions were determined on the JEOL
Superprobe using natural and synthetic minerals for
standards (Tables 6–9). Oxygen abundance in silicate
minerals was based upon stoichiometry. Fe–Ti oxide
minerals were analyzed as weight per cent elements, but
oxygen was analyzed along with the cations. For these
Fe–Ti oxide minerals, ferric iron was calculated assuming
that all elements apart from Fe have fixed valence.
Lincoln granite
Olivine
Samples of Lincoln granite vary from 72·0 to 76·2 wt %
SiO2, the highest for any rock type in the Sherman
We analyzed fayalitic olivine from a fayalite monzonite
inclusion in the Sherman granite (90SMW4) and two
1783
JOURNAL OF PETROLOGY
VOLUME 40
NUMBER 12
DECEMBER 1999
Table 4: Rare earth element contents of selected Sherman batholith samples
91SMW20
91SMW30
91SMW28
91SMW2
91PH6
91SMW4
monzodiorite
ferrodiorite
Sherman
porphyritic
Lincoln
sodic series
73·5
70·9
SiO2
60·4
54·7
La
66·9
20·3
136
145
59·1
48·3
282
270
95·3
32·1
113
Ce
133
Nd
66·9
Sm
13·8
Eu
3·19
68·95
69·2
94·0
40·9
6·98
14·1
6·76
7·77
18·2
15·9
7·64
1·48
4·2
3·0
2·4
1·15
0·362
12·6
6·31
1·36
Gd
16·5
8·49
15·6
Dy
13·2
7·97
11·5
9·77
5·11
1·34
Er
6·64
3·96
5·61
4·64
2·52
0·813
Yb
5·89
3·67
4·36
4·44
2·59
0·89
Lu
0·86
0·562
0·783
0·735
0·401
0·17
Rare earth elements were determined by ICP-MS at the University of Nebraska. Errors estimated from replicate analyses
are less than 8% for La–Gd, 12% for Dy, and 20–25% for Er, Yb and Lu.
Fig. 5. Variation of major element contents of rocks from the Sherman batholith.
samples of the Sherman granite (90SMW5, 90SMW9;
Table 5). The olivine in the fayalite monzonite sample
is richer in iron (Fa95Tp3Fo2) than that from the Sherman
granite (Fa92Tp2Fo6) (Table 6; Fig. 11). The olivine from
the fayalite monzonite is similar in composition to that
from the monzosyenitic plutons of the Laramie anorthosite complex (Fuhrman et al., 1988; Kolker & Lindsley,
1989; Anderson, 1995).
1784
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Fig. 7. Diagram showing the variation in normalized rare earth
abundance for various units of the Sherman batholith. Normalizing
values from Hanson (1980).
Fig. 8. Plot of Rb/Ba against SiO2 for rocks of the Sherman batholith.
The extreme enrichment in Rb/Ba in the Lincoln granites is indicative
of feldspar fractionation.
Pyroxene
Fig. 6. (a) Plot of Al2O3/(CaO + Na2O + K2O) calculated on a
molecular basis against SiO2. (b) Plot of FeOt/(FeOt + MgO) against
SiO2. The Sherman batholith lies within the tholeiite field. Calcalkalic–tholeiitic (CA–TH) boundary from Miyashiro (1974). (c) Plot
of Nb against Y showing how the Sherman granite plots well within
the within-plate granites field, whereas the Lincoln and a few of the
porphyritic granites trend into the field for volcanic arc granites. This
could reflect either a different source for the Lincoln or the effect of
Y depletion during differentiation. ORG, ocean ridge granites; VAG
+ syn-COLG, volcanic arc granites and syn-collisional granites; WPG,
within-plate granites. Fields from Pearce et al. (1984).
We analyzed augite from a sample of ferrodiorite
(90SMW30) and of fayalite monzonite (90SMW4), and
two samples of the Sherman granite (90SMW4;
90SMW9) (Table 3; Fig. 11). Inverted pigeonite occurs
in the ferrodiorite and in 90SWM9; the latter sample
also contains primary orthopyroxene that displays few
augite lamellae. The orthopyroxene composition was
obtained directly from the microprobe analyses. In contrast, compositions of highly exsolved pigeonite and augite
were reconstructed using image analysis.
1785
JOURNAL OF PETROLOGY
VOLUME 40
Fig. 9. Plot showing variation of normative Ab–An–Or in the Sherman
batholith. It should be noted that the Lincoln granite lies along the
cotectics for this system as determined by Nekvasil & Lindsley (1990).
NUMBER 12
DECEMBER 1999
contain grunerite (and also fayalite) display hornblende
compositions that extend to lower values of total Al and
(Na + K). This probably reflects reaction of fayalite with
feldspar to produce hornblende, and suggests that, like
the grunerite, at least some hornblende in the fayalitebearing Sherman granite samples formed under subsolidus conditions.
Application of the Al-in-hornblende barometer (Anderson & Smith, 1995) to samples with the limiting assemblage yields 1·8 kbar at 850°C. Because the
temperature uncertainty is ±50°C, the uncertainty in the
pressure estimate is ±1·8 kbar. Although the hornblende
composition does not constrain pressure, its composition
resembles that of hornblendes that have crystallized from
high temperature, and low-pressure magmas (see Anderson & Smith, 1995).
Amphibole
Biotite
There are two types of amphibole in the Sherman granite.
All rocks, apart from most of the Lincoln granites, contain
hornblende. Fayalite-bearing Sherman granite also contains minor grunerite that has replaced fayalite. Hornblende contains 1·8 atoms p.f.u. of total Al in all rock
types of the Sherman granite (Table 7). Samples that
Biotite from the Sherman batholith shows textures indicating both primary and secondary formation. The Ti
contents of both generations of biotite are identical,
suggesting that formation of biotite from olivine and
ilmenite was a late-stage magmatic reaction. Both are
rich in iron (XFe = 0·75–0·90; Table 8; Fig. 12). The
Fig. 10. Harker diagrams comparing the Sherman batholith with evolved rocks of the Laramie anorthosite complex for critical elements. Scoates
et al. (1996) interpreted the LAC monzonitic rocks as differentiates of ferrodiorite. The high Ba and Zr in some monzodiorites and in some of
the Sybille pluton is the result of accumulation of feldspar and zircon. The geochemical similarities of Sherman and LAC dioritic rocks extend
to REE characteristics. The REE pattern of a Sherman ferrodiorite sample is similar to LAC ferrodiorites that Mitchell et al. (1996) suggested
had accumulated plagioclase, and the REE pattern of an MgO-poor monzodiorite sample resembles REE patterns for more evolved diorites at
Virginia Dale (Vasek & Kolker, 1999). Data from Anderson (1995), Mitchell et al. (1996), Scoates et al. (1996) and this study.
1786
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Table 5: Mineral assemblages of samples for which mineral chemistry was obtained
Sample
Rock type
Q
Ksp
Plag
Bio
Hb
Aug
Pig
91SMW30
ferrodiorite
90SMW4
monzonite
X
90SMW5
Sherman granite
X
90SMW9
Sherman granite
91SMW5
Sherman granite
O
X
L
L
X
X
O,M
X
L
L
X
X
O,M
X
L
L
X
X
X
O,M
X
L
L
X
X
M
X
X
X
X
90PH1
90PH4
Sherman granite
X
M
X
X
X
X
X
X
porphyritic granite
X
M
X
X
X
X
X
90SME2
porphyritic granite
X
O,M
X
X
X
X
X
90SMW19
porphyritic granite
X
O,M
X
X
X
X
90SMW1
Lincoln granite
X
M
X
X
91SMW12
Lincoln granite
X
M
X
X
91SMW26
Lincoln granite
X
M
X
X
X
Ol
Ti
Flu
X
X
X
X
Zir
Ilm
Mag
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
X
Ksp, alkali feldspar; Plag, plagioclase; Bio, biotite; Hb, hornblende; Aug, augite; Pig, pigeonite; Ol, olivine; Ti, titanite; Flu,
fluorite; Zir, zircon; Ilm, ilmenite; Mag, magnetite; O, orthoclase; M, microcline; L, late magmatic.
Table 6: Representative pyroxene and olivine analyses from the Sherman batholith
Opx
Pigeonite
Pigeonite
Augite
Augite
Augite
Augite
Olivine
Olivine
90SMW9
90SMW9
90SMW30
90SMW4
90SMW5
90SMW9
90SMW30
90SMW4
90SMW9
Sherman
Sherman
ferrodiorite monzonite Sherman
Sherman
ferrodiorite
monzonite Sherman
SiO2
47·60
47·85
48·73
49·49
49·93
48·91
50·88
30·04
29·86
Al2O3
0·18
0·23
0·49
0·79
1·28
0·82
0·90
—
—
TiO2
0·11
0·07
0·13
0·20
0·23
0·19
0·19
—
—
FeO
42·82
40·61
35·38
27·68
24·51
24·64
19·80
67·02
65·79
MnO
1·41
1·28
0·86
0·77
0·63
0·70
0·48
2·34
1·56
MgO
5·91
5·75
8·86
3·44
4·73
5·42
7·87
0·78
2·19
CaO
1·54
3·70
5·48
17·55
18·34
18·49
19·50
0·03
Na2O
0·03
0·08
0·08
0·30
0·51
0·29
0·23
Total
99·61
99·57
100·03
100·24
100·17
99·45
99·85
—
100·22
0·02
—
99·43
Cation proportions calculated on the basis of 6 oxygens for pyroxenes and 4 oxygens for olivine
Si
1·995
1·995
1·975
1·997
1·988
1·971
1·989
Al
0·009
0·011
0·023
0·038
0·060
0·039
0·042
—
—
Ti
0·003
0·002
0·004
0·006
0·007
0·006
0·006
—
—
Fe
1·500
1·418
1·205
0·935
0·816
0·833
0·648
1·880
1·844
Mn
0·050
0·045
0·030
0·026
0·021
0·024
0·016
0·067
0·044
Mg
0·369
0·358
0·536
0·207
0·281
0·326
0·459
0·039
0·110
Ca
0·069
0·164
0·232
0·758
0·782
0·795
0·816
0·001
Na
0·002
0·006
0·006
0·024
0·040
0·022
0·018
Wo
0·036
0·085
0·118
0·399
0·416
0·407
0·424
Fa
0·947
0·923
En
0·190
0·184
0·272
0·109
0·149
0·167
0·238
Tp
0·034
0·022
Fs
0·774
0·731
0·111
0·492
0·434
0·426
0·337
Fo
0·020
0·055
X Fe
0·803
0·799
0·692
0·819
0·744
0·719
0·586
X Fe
0·980
0·944
1787
1·007
—
1·001
0·001
—
JOURNAL OF PETROLOGY
VOLUME 40
NUMBER 12
DECEMBER 1999
iron contents of biotite correlate with the FeOt/(FeOt +
MgO) ratio of whole rocks: the most iron-rich biotite is
from the fayalite monzonite, whereas the most magnesian
biotite is from porphyritic granite. The Al content of
biotite also reflects the whole-rock composition. Biotite
with low Al contents comes from the low-SiO2 Sherman
granite samples that contain pigeonite and fayalite, and
hence are saturated in annite by the equilibrium
3 Fe2SiO4 + 2 KAlSi3O8 + H2O =
fayalite
kspar
fluid
2 KFe3AlSi3O10(OH)2 + 3 SiO2.
biotite
quartz
Fig. 11. Pyroxene quadrilateral showing the compositions of pyroxenes
and olivine from the Sherman batholith.
The most aluminous biotite comes from the Lincoln and
porphyritic granites, both of which show higher Al/(Ca
+ Na + K) than the Sherman granite.
Table 7: Representative hornblende analyses from Sherman batholith
90SMW30
90SMW4
90SMW5
90SMW9
90SMW19
91SMW5
90PH4
90PH1
ferrodiorite
monzonite
Sherman
Sherman
porphyritic
Sherman
porphyritic
Sherman
SiO2
41·15
40·96
40·53
41·13
40·46
40·71
40·90
39·53
Al2O3
10·65
9·07
9·06
8·97
9·89
9·51
9·43
10·20
TiO2
1·72
1·11
2·21
2·05
1·35
2·17
1·58
1·63
FeO
25·57
31·42
28·70
27·73
28·67
27·93
26·29
29·52
MnO
0·21
0·57
0·29
0·36
0·66
0·65
0·74
0·78
MgO
5·59
2·45
3·55
4·07
3·06
3·24
4·31
2·13
CaO
10·98
10·54
10·68
10·50
11·10
11·59
11·32
11·35
Na2O
1·20
1·44
1·87
1·80
1·42
1·70
1·33
2·00
K 2O
1·46
1·13
1·29
1·18
1·55
1·16
1·36
1·61
H 2O
1·81
1·76
1·58
1·61
1·65
1·62
1·72
1·62
F
0·01
0·02
0·44
0·42
0·34
0·32
0·27
0·24
Cl
0·41
0·39
0·35
0·31
0·24
0·46
0·13
0·52
F,Cl=O
−0·10
−0·10
−0·26
−0·25
−0·20
−0·24
−0·14
−0·22
Total
100·68
100·77
100·28
99·87
100·19
100·81
99·24
100·91
Cation proportions calculated on the basis of 23 oxygens
Si
6·418
6·566
6·473
6·547
6·472
6·456
6·525
Al
1·958
1·715
1·705
1·682
1·865
1·778
1·773
1·931
AlIV
1·582
1·434
1·527
1·453
1·528
1·545
1·475
1·651
AlVI
0·375
0·281
0·178
0·229
0·337
0·233
0·298
0·280
Ti
0·202
0·134
0·266
0·245
0·162
0·259
0·190
0·197
Fe
3·335
4·213
3·833
3·691
3·835
3·704
3·506
3·966
Mn
0·027
0·078
0·039
0·049
0·089
0·088
0·100
0·106
Mg
1·300
0·586
0·844
0·965
0·729
0·766
1·025
0·510
Ca
1·835
1·810
1·827
1·790
1·903
1·969
1·935
1·953
Na
0·363
0·449
0·580
0·556
0·439
0·524
0·411
0·622
K
0·291
0·232
0·263
0·240
0·316
0·235
0·278
0·331
H
1·885
1·882
1·683
1·706
1·765
1·717
1·830
1·735
F
0·007
0·012
0·221
0·211
0·170
0·160
0·135
0·123
Cl
0·109
0·106
0·096
0·083
0·064
0·123
0·034
0·142
X Fe
0·720
0·878
0·820
0·793
0·840
0·829
0·774
0·886
1788
6·349
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Table 8: Representative biotite analyses from the Sherman batholith
90SMW5
90SMW1
91SMW5
90PH1
91SMW12 90PH4
Sherman
Lincoln
Sherman
Sherman
Lincoln
90SMW19 90SMW4
90SMW9
porphyritic porphyritic monzonite Sherman
90SME2
91SMW26
porphyritic Lincoln
SiO2
34·05
35·18
34·78
34·91
34·89
35·72
35·47
34·00
35·00
35·53
34·83
Al2O3
13·10
15·07
13·55
13·30
13·70
13·95
13·46
13·26
13·73
14·37
14·42
TiO2
3·80
3·56
3·06
3·42
2·75
3·25
2·85
3·61
3·24
2·39
3·33
FeO
32·85
27·10
31·45
31·54
30·61
27·44
29·65
33·95
30·09
28·18
30·88
MnO
0·15
0·38
0·40
0·44
0·42
0·44
0·39
0·12
0·38
0·45
0·49
MgO
2·38
4·40
2·99
2·86
4·23
5·35
4·54
2·15
3·82
4·97
2·53
K 2O
8·66
9·34
8·95
9·08
8·94
8·98
9·17
8·69
9·02
9·18
9·01
Na2O
0·04
0·06
0·03
0·06
0·05
0·05
0·04
0·03
0·05
0·07
0·04
H 2O
3·59
3·57
3·53
3·41
3·64
3·60
3·24
3·59
3·23
3·27
3·55
Cl
0·19
0·14
0·24
0·23
0·07
0·06
0·23
0·13
0·17
0·18
0·29
F
0·00
0·14
0·18
0·46
0·12
0·32
0·91
0·08
0·30
0·41
0·17
−0·04
−0·09
−0·13
−0·25
−0·07
−0·15
−0·44
−0·06
−0·17
−0·21
−0·14
98·77
98·85
99·05
99·47
99·36
99·01
99·52
99·54
98·86
98·77
99·40
O=F,Cl
Total
Cation proportions calculated on the basis of 24 oxygens
Si
5·604
5·611
5·669
5·675
5·638
5·692
5·703
5·576
5·660
5·692
Al
2·541
2·832
2·604
2·548
2·610
2·619
2·552
2·563
2·513
2·713
2·751
AlIV
2·396
2·389
2·331
2·325
2·362
2·308
2·297
2·424
2·340
2·308
2·365
AlVI
0·145
0·444
0·273
0·223
0·248
0·311
0·255
0·139
0·173
0·405
0·386
Ti
0·471
0·428
0·375
0·418
0·334
0·389
0·345
0·445
0·404
0·288
0·405
Fe
4·521
3·615
4·287
4·287
4·137
3·657
3·987
4·657
4·161
3·776
4·179
Mn
0·020
0·052
0·056
0·060
0·058
0·060
0·053
0·016
0·020
0·060
0·067
Mg
0·584
1·047
0·727
0·694
1·019
1·271
1·088
0·527
1·024
1·187
0·609
K
1·818
1·900
1·861
1·883
1·843
1·826
3·882
1·819
1·788
1·876
1·860
Na
0·012
0·019
0·010
0·018
0·015
0·015
0·014
0·008
0·009
0·022
0·013
H
3·948
3·894
3·842
3·699
3·919
3·822
3·474
3·924
3·538
3·744
3·833
Cl
0·053
0·037
0·066
0·064
0·019
0·017
0·062
0·036
0·057
0·048
0·079
F
0·000
0·069
0·092
0·237
0·062
0·161
0·464
0·040
0·405
0·209
0·088
X Fe
0·886
0·775
0·855
0·861
0·802
0·742
0·786
0·898
0·803
0·761
0·873
Biotite compositions from Sherman batholith samples
lie on the iron-rich end of a trend of biotite compositions
from 1·4 Ga granites in the SW USA (Anderson &
Bender, 1989; Fig. 12), although they are less iron rich
than biotite from the 1·1 Ga Pikes Peak batholith. This
indicates that the Sherman batholith is both the most
iron rich and the most reduced of the 1·4 Ga anorogenic
granite batholiths in the SW USA.
Fe–Ti oxide minerals
Back-scattered electron images show that most ilmenite
grains are altered to a cryptocrystalline mixture of phases,
probably magnetite and rutile. Only a few grains preserve
the oxygen to cation ratio typical of ilmenite. These had
hematite contents of 0·04–0·06 (Table 9), values typical
5·635
of ilmenite from fayalite granites (Frost et al., 1988). Mn
contents of ilmenite from the Sherman granite range
from <1 wt % to >16%; much of this range may be
displayed within a single sample. Because olivine in these
rocks is not rich in MnO, the Mn contents of ilmenite
probably result from the oxidation of ilmenite to magnetite + rutile, in which Mn released by this reaction
was sequestered in residual ilmenite.
Magnetite shows typical oxyexsolution lamellae of
ilmenite, as well as low-temperature oxidation (probably
to maghemite). Using the ilmenite composition in the
rock, assuming the magnetite was stoichiometric magnetite, and taking into account the abundance of the
ilmenite lamella in magnetite, the primary titanomagnetite composition from sample 90SMW9 was calculated to be Usp45. Despite uncertainties in this
calculation, this composition is consistent with the silicate
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JOURNAL OF PETROLOGY
VOLUME 40
NUMBER 12
DECEMBER 1999
Fig. 12. Plot comparing the composition of biotite from the Sherman batholith with biotite from other Mid-Proterozoic A-type granites in
southwestern USA. Data from Barker et al. (1975) and Anderson & Bender (1989).
assemblage of the rock. Usp45 titanomagnetite, fayalite
and quartz equilibrated at ~800°C, consistent with the
other thermometers applied to this sample.
INTENSIVE FACTORS
Pressure
The assemblage olivine–pigeonite–augite–quartz present
in sample 90SMW9 (Table 5) is both a geothermometer,
because temperature can be determined from the pyroxene solvus or from the low-T limit of pigeonite, and a
geobarometer, because the iron contents of ferromagnesian silicates increase with pressure. The olivine
and pyroxene in 90SMW9 (Table 6) equilibrated at 624
± 15°C and 3130 ± 720 bar, as determined from the
QUILF program (Andersen et al., 1993). This temperature
reflects subsolidus re-equilibration. The uncertainties
given from the QUILF expression are the calculated
precision of the various equilibria involved. Where only
a single equilibrium is used there is no uncertainty given.
The accuracy of the thermometer is probably ±30°C
or less.
In these samples, mineral compositions changed in
four ways during cooling: (1) by exchange of Ca and (Fe
+ Mg) between augite and pigeonite; (2) by inversion
of pigeonite to orthopyroxene; (3) by redistribution of Fe
and Mg among olivine, orthopyroxene and augite, which
probably continued after Ca ceased to exchange between
pyroxenes; (4) during late hydration reactions by which
hornblende, biotite, and grunerite were produced.
By analyzing the effect of exchange or hydration reactions on the XFe of minerals in a given rock, one can
qualitatively estimate the changes in mineral compositions as the rock cools. Fe–Mg exchange should cause
the phases with high XFe to become progressively richer
in iron, and those with low XFe to become more magnesian
because Fe–Mg distribution is more extreme at lower
temperatures. As a hydration reaction progresses, the
parent will be enriched in Fe if it has a higher XFe than
the hydrated product; it will be enriched in Mg if it has
a lower XFe (see Thompson, 1976).
In order of decreasing XFe, the minerals in 90SMW9
are: fayalite > orthopyroxene > biotite > hornblende >
grunerite > augite (if a melt were present, it would have
XFe > fayalite). During cooling, ion exchange reactions
1790
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Table 9: Representative ilmenite analyses from the
Sherman batholith
90SMW4
90PH1
90SMW9
90SMW19
monzonite
Sherman
Sherman
porphyritic Lincoln
90SMW1
Ti
30·10
29·73
30·46
30·31
Al
0·05
0·05
0·05
0·04
0·04
Cr
0·00
0·00
0·03
0·01
0
Fe
37·57
29·09
36·99
22·39
33·08
Mn
0·87
8·43
1·12
16·03
4·50
Zn
0·03
1·80
0·01
0·47
0·11
Mg
0·02
0·01
0·01
0·06
0·00
31·71
31·75
31·79
31·97
31·71
100·35
100·85
100·46
101·29
100·04
O
Total
30·61
Fig. 13. Pyroxene quadrilateral plot showing the composition of pyroxenes and olivine analyzed from 90SMW9 (filled circles) with calculated compositions. Ruled circles give presumed high-T compositions
as dictated by the Fe/Mg ratios of orthopyroxene and pigeonite;
shaded circles give the compositions if re-equilibration continued to
the temperature of hydration (665°C). (See text for discussion.)
Cation proportions calculated on the basis of 3 oxygens
Ti
0·951
0·938
0·960
0·950
0·968
Al
0·003
0·003
0·003
0·002
0·002
Cr
0·000
0·000
0·001
0·000
0·000
Fe
1·018
0·787
1·000
0·602
0·897
Fe3+
0·100
0·118
0·088
0·086
0·074
Fe2+
0·918
0·669
0·912
0·516
0·823
Mn
0·024
0·232
0·031
0·438
0·124
Zn
0·001
0·042
0·000
0·011
0·002
Mg
0·001
0·000
0·001
0·004
0·000
O
3·000
3·000
3·000
3·000
3·000
cations
1·998
2·002
1·995
2·006
1·994
MnTiO3
0·024
0·232
0·031
0·438
0·124
ZnTiO3
0·001
0·042
0·000
0·011
0·002
FeTiO3
0·918
0·669
0·912
0·516
0·823
Fe2O3
0·050
0·059
0·044
0·043
0·037
Al2O3
0·001
0·001
0·001
0·001
0·001
TiO2
0·009
−0·005
0·017
−0·015
0·018
1·002
0·998
1·005
0·994
1·006
reaction, indicating that during cooling the relative
abundances of olivine and orthopyroxene remained the
same, and that Fe–Mg exchange predominated.
During cooling, the minerals in 90SMW9 followed
reaction (1) to higher XFe (Lindsley & Frost, 1992). Thus,
pressure estimates for sample 90SMW9 depend on the
temperatures at which Fe–Mg exchange ceased. If the
Fe/Mg of orthopyroxene was fixed when pigeonite inverted to orthopyroxene, the pressure estimate is 3000
bar. If cooling continued down to 665°C (when hydration
ceased), then the estimate is 1700 bar (Fig. 13). Because
Fe–Mg exchange probably ceased before the texturally
latest hydration reactions, the Sherman granite probably
was emplaced at 2500 bar ± 500 bar.
Sum
Sum
oxides
Temperature and water activity
will make olivine and orthopyroxene richer in Fe, as
augite and hornblende become richer in Mg. Hydration
of olivine to grunerite and orthopyroxene to biotite will
enrich the remaining olivine and orthopyroxene in iron.
Because the weighted XFe of Opx + Cpx resembles that
of hornblende, hydration of pyroxenes to hornblende
only slightly changes XFe of the minerals.
Mineral compositions in the assemblage olivine–
orthopyroxene–quartz are governed by two reactions:
the Fe–Mg reaction between olivine and orthopyroxene,
and the mass transfer reaction
orthopyroxene = olivine + quartz.
There is no textural evidence for the mass transfer
Geothermometers record a range of temperatures during
the crystallization and cooling of the Sherman batholith.
The highest temperature, 968°C, comes from the composition of reconstructed pigeonite in ferrodiorite sample
90SMW30 [as determined by the QUILF program of
Andersen et al. (1993)]. This temperature critically depends on how closely the reconstructed pigeonite composition represents the original pigeonite composition.
However, the temperature at which pigeonite inverts is
a robust thermometer that depends only on the XFe of
the pyroxenes (Lindsley & Frost, 1992). The pigeonite in
90SMW30 inverted at 874°C, and Fe–Mg exchange
between pigeonite and augite continued down to 789°C.
The small ferrodiorite body from which 90SMW30 was
sampled is surrounded by the far more voluminous
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JOURNAL OF PETROLOGY
VOLUME 40
Fig. 14. Crystallization history of fayalite granite from the Sherman
batholith, sample 90SMW9. QUILF surface calculated from Andersen
et al. (1993), biotite-bearing equilibria from TWQ program of Berman
(1991), and the low oxygen fugacity limits of titanite from Xirouchakis
& Lindsley (1998).
Sherman granite, and therefore 874°C is also the minimum temperature of the granite.
Reconstructed pigeonite compositions from fayalite
granite sample 90SMW9 record a T of 869°C; it inverted
to orthopyroxene at 806°C (Fig. 14). These temperatures
are based on exsolution textures of the pigeonite, which
probably formed late in the crystallization history of the
rock. Because orthopyroxene also occurs in 90SMW9,
crystallization of this sample continued to temperatures
at least as low as 800°C. Because the Ca content of
orthopyroxene is nearly what would be predicted for
orthopyroxene in equilibrium with pigeonite, Ca exchange among the pyroxenes must have ceased at temperatures of ~800°C.
The temperature of hydration from sample 90SMW9
was calculated from the equilibria
3 Fe2Si2O6 + 2 KAlSi3O8 + 2 H2O =
orthopyroxene
kspar
fluid
2 KFe3AlSi3O8(OH)2 + 6 SiO2
biotite
quartz
7 Fe2SiO4 + 9 SiO2 + 2 H2O = 2 Fe7Si8O22(OH)2.
olivine
quartz
fluid
grunerite
Using the TWQ program and biotite model of Berman
(1991, 1992; version 2.02), the feldspar solution model
of Elkins & Grove (1990), the program SOLVCALC
(Wen & Nekvasil, 1994), and the grunerite thermodynamic data of Evans & Ghiorso (1995) and Ghiorso et
al. (1995), we calculate a sanidine activity of 0·76, a water
activity of 0·7, and that hydration occurred at 665°C.
Two other thermometers are available, the zirconsaturation temperature (Watson & Harrison, 1983) and
the apatite-saturation temperature (Harrison & Watson,
1984). These thermometers yield temperatures at which
a melt of a given composition would become saturated
NUMBER 12
DECEMBER 1999
Fig. 15. Plot comparing the zircon saturation temperature and apatite
saturation temperature of rocks from the Sherman batholith using the
models of Watson & Harrison (1983) and Harrison & Watson (1984).
Samples for which zircon saturation temperatures equal apatite saturation temperatures lie along dashed line.
with each of these two accessory minerals, and are valid
only if the rock represents a liquid composition. If zircon
and apatite both crystallize from a magma early in its
history but do not accumulate, their saturation temperatures should be nearly identical. This is the case with
the Lincoln granite, but not for the rest of the Sherman
batholith (Fig. 15). Samples of the Lincoln granite that
have the lowest zircon and apatite saturation temperatures lie along the lowest temperature portions of
the An–Ab–Or cotectic (Fig. 9). These zircon and apatite
saturation temperatures, 750–875°C, are similar to the
liquidus temperatures determined for haplogranite melts
of the same composition at 5 kbar water pressure (Nekvasil
& Lindsley, 1990). The lower crystallization temperature
of the Lincoln granite compared with the Sherman
granite and ferrodiorite reflects its more siliceous bulk
composition (Nekvasil & Lindsley, 1990) and greater
water activity than the Sherman granite.
Oxygen fugacity
Because the Fe–Ti oxide minerals in the Sherman granite
are altered, its oxygen fugacity is difficult to characterize.
The oxygen fugacity of the two samples of Sherman
granite that contain fayalite + quartz can be determined
because these rocks must have crystallized along the
QUILF surface (Frost et al., 1988). Their final oxygen
fugacities lay between 0·1 and 0·5 log units below those
of the fayalite–magnetite–quartz (FMQ) buffer (Fig. 16).
Oxygen fugacity also can be determined for sample
90SMW1 of Lincoln granite that contains biotite–
orthoclase–plagioclase–magnetite–ilmenite. The ilmenite
and magnetite in this rock are both altered, but because
the magnetite is saturated in Ti, the location of reaction
(3) can be calculated using the TWQ program:
1792
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Fig. 16. Plot of the deviation of the log of the oxygen fugacity from that of the FMQ buffer (Dlog f O2) against T for the Sherman batholith
and other A-type granites of North America. Also shown for comparison is the oxygen fugacity of melts generated in the experiments of Skjerlie
& Johnston (1993). Data are from this study, Anderson (1983), Frost et al. (1988), and Anderson & Bender (1989).
2 Fe3O4 + 2 KAlSi3O8 + 2 H2O =
magnetite
sanidine
fluid
2 KFe3AlSi3O10(OH)2 + O2.
in biotite
These calculations were made using the composition of
biotite in 90SMW1 and assuming the activity of water
as 1·0, which yields the highest oxygen fugacity likely for
the assemblage, 0·5 log units above the FMQ buffer.
The range of input temperatures varied from 750 to
665°C, which encompasses the likely range of solidus
temperatures.
Pb, Nd AND Sr ISOTOPIC
COMPOSITIONS
Initial Pb, Nd and Sr isotopic compositions of samples
from each unit of the Sherman batholith constrain possible magma sources (Fig. 17). Initial Pb isotopic compositions were estimated by the least radiogenic fraction
of stepwise dissolved feldspars (Table 9); Nd and Sr initial
isotopic ratios were calculated from present-day wholerock values, corrected for radiogenic growth since 1·43
Ga (Table 10).
Except for the Pole Mountain gneiss, different rock
units of the Sherman batholith cannot be distinguished
by their initial Pb isotopic compositions (Fig. 17a). The
only variation outside of error is in the 206Pb/204Pb ratio,
which could reflect modest radiogenic growth after 1·43
Ga, rather than true variation in initial composition. The
Pb isotopic data from Sherman batholith feldspars plot
above the model mantle evolution curve but below the
model upper-crust evolution curve of Zartman & Doe
(1981). The Sherman batholith feldspar compositions
have much less radiogenic 207Pb/204Pb ratios than Archean rocks from the Wyoming province at 1·43 Ga
(Laramie Mts, Verts et al., 1996; Wind River Range,
Frost et al., 1998; Beartooth Mts, Wooden & Mueller,
1988). Initial Pb isotopic compositions of feldspars from
the Sherman batholith also have distinctly less radiogenic
207
Pb/204Pb ratios than the 1·76 Ga Horse Creek anorthosite complex immediately to the north (Chamberlain,
1998), and display slightly less radiogenic to similar 207Pb/
204
Pb ratios compared with the ~1·78 Ga granitoids of
the Colorado province that lie 25–100 km to the south
(Aleinikoff et al., 1993; Fig. 17a).
In contrast to their initial Pb isotopic compositions,
initial Nd and Sr isotopic ratios vary outside of error
(Fig. 17b). eNd values for Sherman batholith samples at
1·43 Ga vary from 1·1 to –1·5; the range of initial 87Sr/
86
Sr is larger, from 0·701 to 0·731 (Table 11). Samples
of the Pole Mountain gneiss, mafic rocks, and Sherman
granite define the tightest cluster, with eNd –0·4 to 1·1
and initial Sr of 0·7024–0·7126. The variation in Sr and
Nd initial compositions of the Lincoln and porphyritic
granites is larger, and includes the most negative eNd
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JOURNAL OF PETROLOGY
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NUMBER 12
DECEMBER 1999
Table 10: Alkali feldspar Pb
analytical results from the Sherman
batholith
Sample
Rock type
206
204
Pb/
207
Pb/
208
Pb/
Pb
204
Pb
204
Pb
91PH6 step 4
Lincoln
16·656
15·429
35·762
91SMW9 step 4
Lincoln
16·993
15·478
36·021
91SMW28 step 5
Sherman
16·790
15·444
35·852
91PH1 step 5
Sherman
16·796
15·455
35·859
90SMW9 step 4
Sherman px gr
16·759
15·427
35·810
90PH5 step 3
porphyritic granite 16·650
15·437
35·912
90PH5meg step 4
porphyritic granite 16·710
15·439
35·944
91SMW10 step 3
porphyritic granite 16·761
15·436
35·833
91SMW10meg step 3 porphyritic granite 16·775
15·441
35·863
91SMW11 step 3
Mg-mzdi
16·790
15·415
36·035
GM15a∗
Mzdi Maloin
16·834
15·478
35·853
90SMW4 step 3
Summit monz
17·188
15·506
36·064
91SMW4 step 5
Pole Mtn gneiss
16·661
15·390
35·660
∗GM15a data from Kolker et al. (1991).
meg, K-feldspar megacryst. Isotopic compositions of feldspar
come from analyses of the least radiogenic dissolution step
dissolved in 5% HF after Ludwig & Silver (1977). Lead was
purified in the feldspar and whole-rock samples using HCl–
HBr chemistry, modified from Tilton (1973). Mass spectrometry was performed as described in Table 1.
Uncertainties on 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb are
0·12%, 0·18% and 0·24%, respectively.
values and the most radiogenic initial Sr isotope ratios
(Fig. 17b).
DISCUSSION AND CONCLUSIONS
Petrogenesis of the Sherman batholith
The Sherman batholith (1430–1438 Ma) and the Laramie
anorthosite complex (1431–1436 Ma; Scoates & Chamberlain, 1995; Verts et al., 1996) are part of a widespread
magmatic event that affected much of the southwestern
USA. The Sherman batholith and Laramie anorthosite
complex crystallized at similar levels in the crust: the
apparent pressure at which the Sherman batholith was
emplaced, ~2·5 kbar, is close to pressures of 3–4 kbar
determined for the Laramie anorthosite complex
(Fuhrman et al., 1988; Kolker & Lindsley, 1989; Grant
& Frost, 1990; Spicuzza, 1990). Crystallization temperatures for the Sherman batholith (>900°C to 750°C)
are only slightly lower than those for the monzosyenitic
plutons of the Laramie anorthosite complex of ~950°C
Fig. 17. (a) Pb isotopic compositions of alkali feldspar from the Sherman
batholith and of potential sources at 1·43 Ga. Upper-crust and mantle
model curves are from Zartman & Doe (1981). Fields of nearby crust
are growth curve projections to 1·43 Ga from measured feldspar Pb
isotopic compositions (see text for references). Typical felsic rocks have
238
U/204Pb (l) values of 8–16. Symbols for different phases of the
Sherman batholith follow Figs 3 and 5–12. The Pb data from the
Sherman batholith are consistent with derivation from mafic, mantlederived rocks mixed to varying degrees with Proterozoic crustal sources
similar to the Horse Creek anorthosite complex (HCAC) that crops
out immediately to the north or the ~1·78 Ga Colorado Province to
the south. Archean Wyoming Province sources probably did not
contribute significantly to these samples. (b) Plot comparing the eNd
and 87Sr/86Sr at 1430 Ma for whole-rock samples from the Sherman
batholith, along with Sr and Nd isotopic compositions of possible source
rocks. Northern Colorado province metasedimentary rocks are from
the Idaho Springs formation north of Ft Collins. Symbols for different
phases of the Sherman batholith follow Figs 3 and 5–12. Data are
from Hedge et al. (1967), Peterman et al. (1968), DePaolo (1981),
Zielinski et al. (1981), Nelson & DePaolo (1984), Vasek (1999), and our
unpublished data.
(Fuhrman et al., 1988; Kolker & Lindsley, 1989; Anderson, 1995). The range of oxygen fugacities determined
for the Sherman batholith, from Dlog f O2 of –0·5 to
+0·5, also are similar to oxygen fugacities calculated for
the Laramie anorthosite complex (Dlog f O2 = 0 to –2;
Frost et al., 1996). As is typical of reduced, rapakivitype granite batholiths and coeval, proximal anorthosite
intrusions worldwide, the initial Pb, Sr and Nd isotopic
compositions of rocks from the Sherman batholith and
the Laramie anorthosite complex are comparable. Both
have initial eNd near zero, and initial 87Sr/86Sr ratios
1794
1795
monzodiorite
monzodiorite
ferrodiorite
monzonite
Pole Mt gneiss
91SMW11
90SMW17
91SMW30
90SMW4
91SMW4
43·10
65·70
19·82
89·62
65·20
228·6
207
241·7
198·2
167·8
189·6
193
84·2
112·5
102·7
166·4
303·3
102·8
160·3
Rb (ppm)
418·6
160·3
515·9
1065
740·4
225·7
166
333·2
278·6
128·5
330·9
212
214·2
212·0
180·0
119·4
13·5
88·3
173·7
Sr (ppm)
Rb/86Sr
0·300
1·193
0·112
0·245
0·256
2·948
3·630
2·110
2·069
3·816
1·668
2·650
1·144
1·549
1·661
4·057
65·40
3·399
2·688
87
Sr/86Sr
0·70935
0·73463
0·70925
0·70744
0·70948
0·76916
0·77697
0·74892
0·75857
0·80915
0·76052
0·76864
0·72764
0·74434
0·74054
0·79150
2·21949
0·79448
0·77401
87
0·70320
0·71015
0·70696
0·70242
0·70422
0·70869
0·70077
0·70564
0·71613
0·73087
0·72631
0·71426
0·70417
0·71257
0·70646
0·70827
0·87786
0·72377
0·71887
Initial Sr
1·329
7·20
6·998
20·10
16·33
16·48
16·98
13·54
0·318
27·18
21·61
17·66
15·66
19·82
3·165
8·877
14·69
Sm (ppm)
6·49
34·17
29·81
126·14
98·08
91·23
117·2
76·25
1·880
151·9
121·5
99·36
89·02
109·2
6·454
50·79
90·60
Nd (ppm)
Sm/144Nd
0·12390
0·12735
0·14194
0·09634
0·10066
0·10921
0·08756
0·10738
0·10230
0·10822
0·10750
0·10749
0·10640
0·10968
0·29668
0·10566
0·09802
147
Nd/144Nd
0·512000
0·511979
0·512104
0·511678
0·511733
0·511830
0·511652
0·511771
0·511691
0·511789
0·511853
0·511846
0·511795
0·511848
0·513669
0·511729
0·511649
143
0·510846
0·510783
0·510770
0·510773
0·510787
0·510804
0·510827
0·510762
0·510730
0·510796
0·510843
0·510836
0·510795
0·510817
0·510881
0·510736
0·510728
Initial Nd
1·10
−0·13
−0·38
−0·33
−0·05
0·28
0·82
−0·55
−1·18
0·13
1·04
0·91
0·11
0·53
1·79
−1·06
−1·21
Initial eNd
Analytical details: ~80 mg of sample were dissolved in HF–HNO3. After conversion to chlorides, one-third of the sample was spiked with 87Rb, 84Sr, 149Sm, and
146
Nd. Rb, Sr, and REE were separated by conventional cation-exchange procedures. Sm and Nd were further separated in di-ethyl-hexyl orthophosphoric acid
columns. All isotopic measurements were made on a VG Sector multi-collector mass spectrometer at the University of Wyoming. An average 87Sr/86Sr isotopic
ratio of 0·710246 ± 23 (2r) was measured for NBS987 Sr, and an average 143Nd/144Nd ratio of 0·511846 ± 11 (2r) normalized to 146Nd/144Nd = 0·7219 was measured
for the La Jolla Nd standard. Uncertainties in Sr isotopic ratio measurements are ±0·00002 and uncertainties in Nd isotopic ratio measurements are ±0·00001
(2r). Blanks are <50 pg for Rb, Sr, Nd, Sr, and no blank correction was made. Uncertainties in Rb, Sr, Sm and Nd concentrations are ±2% of the measured value;
uncertainties on initial eNd = ±0·5 epsilon units. Initial Sr and Nd isotopic ratios and initial eNd values are calculated for 1·43 Ga.
porphyritic granite
megacryst
91SMW10meg
megacryst
91SMW2meg
91SMW10
porphyritic granite
Sherman granite
90SMW9
porphyritic granite
Sherman granite
90SMW5
91SMW2
Sherman granite
91SMW28
90SME2
Sherman granite
91PH1
porphyritic granite
Lincoln granite
91SMW9
megacryst
Lincoln granite
91PH6
90PH4meg
Lincoln granite
90SMW3
90PH4
Unit
Sample
Table 11: Nd and Sr isotopic data for the Sherman batholith
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
JOURNAL OF PETROLOGY
VOLUME 40
mostly between 0·703 and 0·710. Initial Pb isotopic
compositions of the Sherman batholith and Laramie
anorthosite complex range from 207Pb/204Pb ratios slightly
above the model mantle curve (for the Sherman batholith
and southern Laramie anorthosite complex) to ratios
more like those of the Archean Wyoming province (for
rocks of the northern Laramie anorthosite complex; this
study; Geist et al., 1989, 1990; Kolker et al., 1991; Mitchell
et al., 1995, 1996; Scoates & Frost, 1996).
The Pb, Nd and Sr isotopic compositions of the
Sherman batholith rule out some potential mantle and
crustal sources. Archean crust, which is present at depth
beneath northern portions of the Sherman batholith
(Chamberlain, 1998), has eNd ~ –15, 87Sr/86Sr = 0·74–
0·80, and 207Pb/204Pb = 15·7 for 206Pb/204Pb = 16·5,
and cannot be a major component of the Sherman
magmas (see Fig. 17). Model depleted mantle, with eNd =
+4 to +6 at 1433 Ma, 87Sr/86Sr of ~0·701, and a
less radiogenic Pb isotopic composition than Sherman
granitoid rocks, cannot be the sole source of Sherman
batholith magmas. Neither the Laramie anorthosite complex nor the Sherman batholith has these model mantle
isotopic compositions, not even the high-alumina gabbros
and anorthosites interpreted as mantle-derived melts
(Mitchell et al., 1995). Perhaps the mantle beneath the
Sherman batholith has less extreme isotopic compositions: a mantle with eNd = ~+2 and 87Sr/86Sr =
0·703 best describes the source of the Laramie anorthosite
complex gabbros and anorthosites, and could also have
been a source for the Sherman batholith. Although a
slightly depleted mantle or mafic lower-crustal source is
compatible with the isotopic data, the isotopic data
also permit a Proterozoic-age felsic crustal source. The
samples display less radiogenic 207Pb/204Pb ratios than
the Horse Creek anorthosite complex immediately to the
north and lie only slightly below a 1·78–1·43 Ga reference
isochron that originates in the field of ~1·78 Ga Colorado
province granitoids (Fig. 17a). The samples also exhibit
Nd and Sr isotopic compositions indistinguishable from
Colorado province volcanic rocks (Fig. 17b).
However, the high liquidus temperatures of the
magmas, the geochemical composition and mineralogy
of the rocks, and, most importantly, oxygen fugacity near
or below FMQ rule out both typical felsic calc-alkalic
and pelitic rocks as sole sources of reduced, rapakivi-type
granites (see Frost & Frost, 1997). These oxygen fugacities
are distinctly lower than those of other A-type granites
and calc-alkalic granites in general (Fig. 16; Frost &
Lindsley, 1992). The relatively low oxygen fugacity is
important because the f O2 of a magma probably reflects
that of its source (Carmichael, 1991). The low f O2 of
the Sherman granite and other reduced, rapakivi-type
granites is evidence for their derivation from a tholeiitic
source, because calc-alkalic sources are more oxidized and
pelitic sources, although appropriately reduced, produce
NUMBER 12
DECEMBER 1999
peraluminous, not metaluminous melts (Frost & Frost,
1997). Although partial melts of tonalite or granodiorite
mimic the bulk compositions of reduced, rapakivi-type
granites, such sources are not likely to have appropriately
low f O2: calc-alkalic rocks typically have Dlog f O2 in the
range of +1 to +3 (Frost & Lindsley, 1991). For
example, the A-type granite melt produced by melting
of tonalite in Skjerlie & Johnston’s (1993) experiments
had f O2 > 1 log unit above FMQ (Fig. 16). Melting
of magnetite-free tonalite could produce magmas of
appropriately low oxygen fugacities, but such magmas
appear to be uncommon. Only some of the tonalites
from the Sierra Nevada are ilmenite-series (Ague &
Brimhall, 1988), but those from the Japanese arc are
magnetite-series (Ishihara, 1979).
Experimental results indicate that A-type granite can
be produced by differentiation or by partial melting of
ferrodiorite (Scoates et al., 1996). Although the Sherman
granite could have been produced by extreme differentiation of a tholeiitic magma as has been proposed for
the Sybille monzosyenite of the Laramie anorthosite
complex (Scoates et al., 1996), this is not likely for several
reasons. First, the volume of Sherman granite is large
compared with the amount of mafic and monzonitic
rocks in the Sherman batholith. The opposite is true in
the Laramie anorthosite complex and other batholiths
for which this model has been proposed. The batholith
lacks the continuum of rock types observed in the Laramie
anorthosite complex, and is instead bimodal. The vicinity
of the Sherman batholith lacks gravity anomalies that
might indicate the presence of a large mass of mafic
cumulates (see Scoates et al., 1996). Instead, the Sherman
batholith was probably derived from partial melting of
pre-existing tholeiites or their differentiates. Tholeiiteseries rocks have low f O2 and f H2O, and evolved compositions have iron and LILE enrichment that are typical
of reduced rapakivi-type granites.
Sherman ferrodiorite and monzodiorite resemble dioritic rocks of the Laramie anorthosite complex (Mitchell
et al., 1996) in their geochemistry (Fig. 10). They are
typical of iron-rich diorites found in many anorogenic
batholiths, often mingled with granite (Wiebe, 1980;
Noblett & Staub, 1990; Eklund et al., 1994; Salonsaari
& Haapala, 1994; Vasek & Kolker, 1999). The mafic
rocks intruded the Sherman batholith while some granitic
magmas were still molten. We suggest that, as with
dioritic rocks of the Laramie anorthosite complex, the
mafic rocks of the Sherman batholith represent samples
of mantle-derived mafic magmas and their differentiates
that were variably contaminated by continental crust
during ascent (see Mitchell et al., 1996).
In the Sherman batholith, Lincoln granite records
incorporation of felsic continental crust. Lincoln granite
is more oxidized than Sherman granite, is peraluminous
whereas the Sherman granite is metaluminous, and has
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FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
biotite richer in Mg than that of the Sherman granite.
In addition, Lincoln samples exhibit radiogenic initial Sr
isotopic compositions (87Sr/86Sr > 0·710). The only known
sources with such radiogenic Sr isotopic compositions
are Proterozoic and Archean metasedimentary rocks,
which had 87Sr/86Sr of 0·715 and higher at 1·43 Ga
(Fig. 17b). Although their Sr isotopic compositions are
uniformly high, the eNd of the metasedimentary rocks
appears to correspond to distance from the edge of the
Archean Wyoming province: pelitic rocks at the south
edge of the Wyoming province have eNd of –16, metasedimentary rocks 25 km south of this boundary have
eNd of –8, and metasedimentary rocks along the Colorado
Front Range have eNd of 0 to +3, all at 1·43 Ga.
Interpolating these data, metasedimentary rocks near the
Sherman batholith probably had eNd of –3 or –4 at 1·43
Ga. A component with an isotopic composition similar
to that of such metasedimentary rocks is present in the
Lincoln granite. Lincoln granite is very similar to the 1·4
Ga Silver Plume and St Vrain batholiths of the Colorado
Front Range in mineralogy, geochemistry, crystallization
conditions and radiogenic Sr isotopic compositions (Anderson & Thomas, 1985).
Much of the porphyritic granite may have formed by
the interaction of mafic and felsic magmas. Feldspar
megacrysts are not in textural equilibrium with the host
rock, as evidenced by plagioclase rims on K-feldspar
megacrysts, and K-feldspar with euhedral cores and
inclusion-rich rims. Clots and schlieren of mafic material
are present in the groundmass of some porphyritic granites. On Harker diagrams, the porphyritic granites invariably lie between the mafic rocks and the Lincoln
granite, and the Sr isotopic compositions of porphyritic
granite samples vary from the least to the most radiogenic,
both of which also suggest magma mixing. To search for
isotopic evidence of disequilibrium, we obtained Rb–Sr
isotopic data for three samples of porphyritic granite,
and for feldspar megacrysts separated from these samples.
In two samples the megacryst is more radiogenic Sr than
the bulk rock, and in the third the megacryst is less
radiogenic (Table 10). One megacryst for which we also
obtained Sm–Nd isotopic data has an initial eNd that is
slightly more than one e unit lower than that of the bulk
rock. These data corroborate the field, petrographic and
geochemical evidence for magma mixing.
Petrogenetic model for A-type granitoids
Our petrogenetic model for the Sherman batholith begins
with the emplacement of mantle-derived mafic magmas
at or near the base of the crust. This basaltic melt
underwent differentiation along a tholeiitic trend at low
f O2, producing ferrodiorite with high FeOt/(FeOt +
MgO) and rich in REE and Zr. These underplated
gabbroic to ferrodioritic rocks represent newly formed
mafic lower continental crust. The presence of such mafic
material beneath southeastern Wyoming is suggested by
the seismic wide-angle studies of Gohl & Smithson (1994).
Those workers interpreted alternating low- and highvelocity layers in the lower crust beneath the Sherman
batholith as evidence of mafic and ultramafic rocks
interlayered with more felsic crust. This mafic lower crust
may have been formed during 2·0 Ga rifting (Cox et al.,
1999), at 1·76 Ga in association with the Horse Creek
anorthosite complex (Frost et al., 1999), or at 1·43 Ga
when the Laramie anorthosite complex and Sherman
batholith formed.
Partial melting of anhydrous underplated mafic material can produce granitic melt with extreme A-type
compositions (Frost & Frost, 1997). We suggest that
ferrodioritic portions of this underplated material are the
most appropriate source rock: a ferrodiorite source will
yield larger volumes of granitic melt than basalt, and
ferrodiorite also has a lower melting point than its basalt
parent. Such a melt can form the Sherman granite, the
unit with the lowest oxygen and water activities, highest
temperatures, and highest incompatible element contents
in the batholith. The composition of this granitic magma
may be altered by assimilation of crustal wall rocks.
Assimilation of metasedimentary rocks could produce the
peraluminous Lincoln granite, for example. Coeval mafic
magmas that commingled with the granitic magmas could
yield the hybrid porphyritic granites.
The Sherman batholith contains both potassic and
sodic granitoids, as does the 1·1 Ga Pikes Peak batholith
of Colorado (Barker et al., 1975), the classic rapakivi
terrane of southern Finland and the Ragunda massif of
Sweden (Rämö & Haapala, 1995). Alkalic basalt and its
differentiates are potential sources for the sodic rocks
(Barker et al., 1975). Like tholeiites, alkali basalts evolve
residual melts that have low oxygen fugacity, but these
magmas have higher Na/K ratios than potassic A-type
granites (Frost & Lindsley, 1991).
The Sherman batholith is the most reduced of the
~1·4 Ga A-type granitoids in the southwestern USA. Its
magmas may have ascended via fundamental crustal
structures related to the Cheyenne belt, which marks the
suture between Archean and Proterozoic crust. Moreover, Archean crust traversed by Sherman magmas was
probably more refractory than Proterozoic crust, so that
A-type granitoids farther south were more contaminated
with oxidized, felsic crust than were magmas of the
Sherman batholith.
Phanerozoic analogs to Proterozoic
anorogenic granites
Phanerozoic A-type granites and rhyolites occur in three
different tectonic settings: (1) they are associated with
1797
JOURNAL OF PETROLOGY
VOLUME 40
mantle plumes, such as the fayalite rhyolites of the
Yellowstone–Snake River Plain province; (2) they occur
in rifted continental settings, such as granite complexes
associated with the opening of the Atlantic Ocean in
Africa, South America and New England; (3) they are
found in areas of large-scale continental extension, such
as the Basin and Range province.
(1) Mantle plumes: the Yellowstone–Snake River Plain
province
The lavas of the Yellowstone–Snake River Plain province
exhibit remarkable petrochemical similarities to reduced,
rapakivi-type granites. The fayalite rhyolites of the main
caldera show extreme high iron and K2O contents, low
f O2 and high incompatible contents, just like the Sherman
granite (Hildreth et al., 1991; Frost & Frost, 1997). The
lavas of the Snake River plain include abundant tholeiitic
basalts and small volumes of differentiated lavas—
icelandites—the eruptive equivalents of ferrodiorites (Leeman et al., 1976). The tectonic setting of the
Yellowstone–Snake River Plain is clearly extensional,
related to a propagating mantle plume (Smith & Braile,
1993).
Hildreth et al. (1991) concluded that the Yellowstone
lavas were produced by partial melting of slightly older
Cenozoic basalt, which generated the large volumes of
fayalite rhyolite that erupted from the main caldera. The
isotopic composition of these rhyolites limited the role of
Archean crust to <15 wt %. Outside the main caldera,
the rhyolites are poorer in iron, and show greater degrees
of assimilation of Archean crust. This model is very
similar to the one we propose for the Sherman batholith
and for other reduced, rapakivi-type granites (Frost &
Frost, 1997).
(2) Continental rifting: the opening of the Atlantic Ocean
Anorogenic magmatism is associated with the break-up
of Gondwanaland, during which bimodal magmatism
was focused along older lineaments or other major zones
of weakness. In Africa, such anorogenic magmatism took
place throughout the Phanerozoic, peaking during the
Mesozoic opening of the Atlantic Ocean (Kinnaird &
Bowden, 1987; Bowden et al., 1990). The similarity
of the Jurassic Nigerian–Niger province to Proterozoic
anorogenic granites was noted by Kisvarsanyi (1981) and
Van Schmus et al. (1993). Both suites contain alkaline to
subalkaline granitic and rhyolitic rocks, with or without
fayalite and Fe-rich pyroxene. The ~410 Ma Aı̈r complex
of Niger consists of ~30 ring complexes. In seven of
these, granite is associated with anorthosite and ferrodioritic rocks (Demaiffe et al., 1991). Moreau et al. (1994)
concluded that the emplacement of the Aı̈r ring complexes was controlled by pre-existing lineaments. Their
NUMBER 12
DECEMBER 1999
tectonic model links a transtensional tectonic regime with
anorogenic magmatism.
(3) Continental extension: Basin and Range province, western
USA
In the Basin and Range province, broadly distributed
extensional deformation began in earliest Oligocene time,
and continues today (Eaton, 1982). Extension-related
magmas include tholeiitic and alkali basalts, basaltic
andesite, and high-silica rhyolite (Eaton, 1982). Many of
the rhyolites are fayalite bearing and reduced (Dlog f O2
near or below the FMQ buffer; Frost et al., 1988). These
fayalite-bearing rhyolites typically contain K2O > Na2O
and K2O > 5 wt %, and include, for example, the Kane
Spring Wash rhyolite, Nevada (Novak & Mahood, 1986),
the Twin Peaks rhyolite, Utah (Crecraft et al., 1981),
lavas of the McDermitt Caldera, Nevada–Oregon (Conrad, 1984) and the Coso volcanic field (Bacon et al.,
1981).
Tectonic environment for A-type granites
Three extant tectonic models for the generation of MidProterozoic anorogenic granites are: (1) large-scale mantle
upwelling beneath a Mid-Proterozoic supercontinent
(Hoffman, 1989); (2) extension and fragmentation of
a Mid-Proterozoic supercontinent (Windley, 1993); (3)
synorogenic gravitational collapse of the hinterland of a
contractional orogeny (Nyman et al., 1994).
Hoffman (1989) concluded that bimodal intraplate
magmatism could result from mantle upwelling, and that
the resultant doming would explain the lack of synplutonic alluvial and lacustrine sediments. He suggested
that a Middle Proterozoic supercontinent effectively insulated the underlying mantle, leading to a very large
area of mantle upwelling and partial melting that he
termed a ‘superswell’. Although a mantle plume of such
large scale may be a unique Proterozoic phenomenon,
we maintain that a smaller-scale version of this upwelling
is represented by the Yellowstone–Snake River Plain
province, which is sourced by a mantle hotspot. The
rhyolites and icelandites of this province are petrochemically equivalent to A-type granites. Mantle upwelling events such as the one Hoffman (1989) described
should also produce magmas with A-type characteristics.
Windley (1993) postulated that Mid-Proterozoic anorogenic granites were generated during extensional collapse
of thickened crust after the assembly of Laurentia. In
this model, Laurentia split to form a southern continent,
which was reattached to Laurentia by the Grenvillian
orogeny. Such intracontinental rifting may lead to magmatism across a broad area (e.g. the Basin and Range)
or in narrow belts (such as in Africa). A-type granite
compositions are found in both environments. However,
1798
FROST et al.
PETROGENESIS OF SHERMAN BATHOLITH, WYOMING
Hoffman (1989) noted that stretched continental lithosphere subsides upon cooling and that thick sedimentary
deposits, such as those in the Basin and Range, should
be found in regions of extensional collapse. Such contemporaneous sediments are not found in the southwestern USA, but their absence may reflect current levels
of exposure.
Several Mid-Proterozoic granite plutons in the southwestern USA record syn-intrusive or post-intrusive deformation (Nyman et al., 1994). Although some of this
deformation may reflect local strain related to emplacement of the plutons, part has been interpreted
as reflecting regional stress fields. Nyman et al. (1994)
proposed that deformation associated with ~1·4 Ga plutons resulted from intraplate strains associated with a
distant contractional orogeny along the southern margin
of Proterozoic North America. Their model is somewhat
analogous to the tectonics of the Tibetan plateau, in
which extension and crustal thinning are taking place at
the same time as convergence (Inger, 1994). The Tibetan
plateau is roughly similar in size to the Mid-Proterozoic
anorogenic granite province of the southwestern USA.
However, the magmas associated with extension in Tibet
are a calc-alkaline continental margin series (Coulon et
al., 1986; Arnaud et al., 1992).
Nyman et al. (1994) observed that most ~1·4 Ga
granitoids in the southwestern USA were emplaced along
or near pre-existing shear zones. The Sherman batholith
lies immediately south of the Cheyenne belt, which marks
the suture between the Archean Wyoming province and
Proterozoic arc terranes to the south. In a number of
cases, Mid-Proterozoic plutonism is synkinematic, with
reverse motion on these reactivated shear zones (e.g. the
Beer Bottle Pass pluton, Duebendorfer & Christensen,
1995; the Lawler Peak and Signal batholiths, Nyman &
Karlstrom, 1994; the Mt Evans batholith, Graubard,
1991). This observation led Nyman et al. (1994) to infer a
regional transpressive tectonic regime. However, normal
motions are also observed (Mummy Range, Moeglin &
Plymate, 1992; Mt Ethel, T. Foster & K. Chamberlain,
personal communication, 1998; Sandia, Kirby et al.,
1995). Thus, it is not clear whether these shear zones
reflect the overall stress field during pluton emplacement
or represent a late structural response following emplacement.
An important feature associated with 1·4 Ga plutonism
in the southwestern USA is a broad thermal anomaly.
K–Ar and 40Ar/39Ar muscovite and hornblende ages of
1·4–1·3 Ga have been obtained from Early Proterozoic
rocks in northern New Mexico (Karlstrom et al., 1997),
in Arizona (Van Schmus et al., 1993; Wendlandt et al.,
1996), and Colorado (Selverstone et al., 1995). In southern
Wyoming, K–Ar and Rb–Sr mineral ages of Archean
rocks were reset at 1·4–1·5 Ga (Peterman & Hildreth,
1978). These ages reflect a regional thermal event rather
than local heating related to emplacement of 1·4 Ga
granitoids, because the reset area is broader than that
affected by plutonism. Indeed, the 1·4–1·5 Ga mineral
ages in southern Wyoming extend more than 50 km
north and 200 km northwest of the nearest 1·4 Ga
intrusions, delineating an area in which temperatures
reached at least 325°C (the blocking temperature for Ar
in biotite). The regional scale of this reheating and
isotopic resetting, the widespread and voluminous A-type
magmatism that occurs throughout the southwestern and
central USA south of the Wyoming craton, and the
presence of an extensive ~1·4 Ga mafic dike swarm in
northern Colorado and southern Wyoming (Braddock &
Peterman, 1989; Chamberlain & Frost, 1995) together
suggest the heat was supplied by the mantle.
The broad thermal perturbation at 1·4 Ga in the
western USA provides strong supporting evidence for
a mantle origin of anorogenic magmatism. Reduced,
rapakivi-type granites, such as those described here from
the Sherman batholith, and their extrusive equivalents,
such as the fayalite rhyolites and icelandites of the Yellowstone–Snake River Plain province, require a source
of tholeiitic-series rocks emplaced at or near the base of
the continental crust. Large-scale mantle upwelling can
provide the heat for partial melting of tholeiitic lower
crust. Ascent along pre-existing structures will bring these
magmas to the level of emplacement with minimal crustal
contamination. In other areas, reduced granitic magmas
may interact with felsic continental crust, producing more
oxidized or peraluminous members of the A-type granite
suite.
ACKNOWLEDGEMENTS
We thank J. L. Anderson and O. T. Rämö for helpful
reviews, and editor S. Sorensen for detailed constructive
suggestions. This research was supported by NSF Grant
EAR9706237 to C. D. Frost and B. R. Frost.
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Aleinikoff, J. N., Reed, J. C., Jr & Wooden, J. L. (1993). Lead isotopic
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