Download 38_1993_Houston et al_Geo of NA_Wyoming province

The Geology of North America
Vol. C-2, Precambtian. Conterminous U.S.
The Geological Society of America, 1993
Chapter 3
The Wyomingprovince
i
R. S. Houston, editor
De-nt
of Geology and Geophysics, UniversQ of Wyoming, Laramie, Wyoming 82071
E. A. Erslev
Department of Earth Resources, Colorado State University, Fort Collins, Colorado 80523
C. D. Frost
Depariment of Geology and Geophysics, University of Wyoming, h a m i e , Wyoming 82071
IC. E. Karlstrom*
Deparbnent of Geology, Northern Arizona University, Flagstaff; Akona 86011
N. J. Page and M. L. Zientek
US. Geological Survey, 345 Middlefield R o d Menlo Park, California 94025
John C. Reed, Jr.
US. Geological Survey, MS 913, Box 25046, Denver Federal Center, Denver, Colorado 80225
G. L. Snyder, R. G. Worl, Bruce Bryant, M. W. Reynolds,* and Z. E. Peterman
US.Geological Survey, Box 25046, Denver Federal Center, Denver, Colorado 80225
INTRODUCTION
LOCATION AND BOUNDARIES
R. S. Houston and K. E. KarIstrom
a wide zone that exhibits a gradational change from Archean
dates in central Montana to Early Proterozoic dates to the
northwest, reflecting increasing influence of post-Archean thermal events.
The southwestern and southern margins of the Wyoming
province are also poorly constrained. Archean rocks have been
reported from several mges in the C o r a e n n orogenic belt
such as the Albion and Raft River Ranges (Armstrong and Hills,
1967) and possibly the Ruby Mountains of Nevada (A. W.
Snake, personal communication, 1986). T ~ Qrocks may be allochthonous relative to the central Wyoming province, but they
neverthela suggest that A r c h a basement may
w&ward under the M m o i c
&cn$ perhap as far as the
87sr/86sr = 0.706 line in western Idaho and central Nevada,
which probably marks the wetern limit of Prambrim be
meat in North America (Plate 1). Archean rocks are also found
in the Wasatch Range and Antelope Island (Hedge and others,
19831, where they were strongly overprinted at about 1700 Ma
(Crittenden and others, 1971). The most southerly ex@
Archean rocks of the Wyoming province are gneiss@of the Owiyuwrtheastern Uinta Mounw.
htsComplex in
The
mar& of
Wyoming proha
s&
to
The W Y o h g province b the r e o n in Wyoming and adJaa
states
by rocks Of
Qe ('late 2). It is an
Archean maton bordered On the east and south younger
Cambrian provinca ('late '1. Prmbdan racks are
exOf the hmide
(Late
in the
to Euly
and Outcr0PScomtiNte leSS than lo Frcent
Of
basement
'piifto, bw
meat rocks are covered by thick Phanerozoic strata, so that exp~~~~~~ of geology kom
generaUytenuous.
Of the Wyoming provinecb
mtNdng the Archean
with the late
mmpl*led
by
Oic
and thrust
the
-gingin
ha
is
the
upmgenerall~
rd excellent exposure.
The northern and northwestern margins of the Wyoming
vince are poorly constrained. Archean rocks are known as far
the
Rocky Mountains
the northwestern margin in that it appears to be a zone of overlap
1' Acmrdiog to King (lg776 Archean and
Rotermic
between h h a a ages the w& and rdUvenaf& &Iy Rotere
are intermingled at the northwest margin of thcp zOic
P
g
a to the eastt kchean rocks mop out in two
small domes in the Black Hills of South Dakota where they are
Departmentof GF'logy$
University of New
ew Mexico 87131; Reynolds, 954 National Center, Res- overlain by metasedimentary rocks that were strongly metamorphosed and deformed during the 1,700 to 1,800 Ma Black Hills
*
sawenmy
Houston, R. S., Erslev, E. A., Frost, C. D., Karlstrom, K. E., Page, N. J., Zientek, M. L., Reed, J. C., Jr., Snyder, G. L., Worl, R. G., Bryant, B., Reynolds, M. W.,
man, Z. E., 1993, The Wyoming province, in Reed, J. C., Jr., Bickford, M. E., Houston, R S., Link, P. K., Rankin, D. W., S i , P. K., and Van Schmus,
eds., Precambrian: Conterminous US.: Boulder, Colorado, Geological Society of America, The Geology of North America, v. C-2.
7
121
I
orogeny of Goldich and others (1966). Subsurface and -pical data (Udiak, 1971; King, 1976) indicate that the buried Pre
Cambrian rocks of central and westem South Dakota and North
Dakota were probably also intensely deformed and m~tamorp M during the Early Protermic.
The only e ~ p t x margin
d
of the Wyoming province is in the
Sierra Madre and Me6icine Bow Mounttiins in southeastern
Wyoming, where the boundary is a major shear m e along
which Arehean and Early Proterozoic miogeoclinal rocks on the
north are juxtaposed with late Early Proterozoic volcanogenic
racks on thg'south. No rocks older than 1,800 Ma have been
documented south ofthis shear zone @ePaolo, 1981; Nelson and
DePaoEo, 1985), which Houston and others (1979) named the
Cheyenne belt. This important feature is &ussed in detail later
in this chapter.
and Orhem
plorat.ion p r . o in~the
~
of world-elass uranium deposits
of the Wyoming province was
strom (1979) as part of
northern Medicine Bow Mountains and Sierra
taken (Karlstrom md others, 1981). The D
included mapping of parts of the U W a d
Laramie Mountains,
HISTORY OF INVESTIGATIONS
The US. Geological Survey b a n geologic mapping in the
Wyoming pro*
in the early 1900s. Notable conDn'butions
included mapping in the Sierra W e s Wind River Range and
Laramie Mountains by Spencer (1904,1916); the Medicine Bow
Mountains by Blackwelder (1926); the Laramie Mountains by
Darton and, others (1910); and the Hmtville Uplift by Smith
(1903).
By the early 1950s most of the Wyoming PreeamLnian still
remain@ m a p p e d even in reconnaissance, so that no coherent
regional view of the geology was possible. This situation was
reeected in the g&gic nap of Wyoming (Love and others,
1955), which showed the Precambrian as undivided.
Gast and others (1958) were the first to report Archean
isotopic ages for various rooks in the different uplifts in Wyolriing
and Montana. This early work was a&@ficant step in the understanding of the geochronology, although it was not possible to
evdwte the geologic significance of m a y of the dates. Regional
geologic studies began in several range of the Wyomhg province in the 1950s. The Geolqjid S m e y of Wyoming began
mapping the Precambrian in 1957, conoentfMhg in southemem
Wyoming, but atso covering subskmtia! areas in o&er Wyaming
ranges. A summary of the geology of the Mediche Bow Momtains was compbted in 1968 (Houston and others, 1968), and
that report included a general review of the Precambrian gwiogy
of the Wyoming prowince to that date. By the la@, Reed had
started a detailed study of the Teton Range (Reed,1963; Reed
and Zartman, 1973), and Bayky had begun geologic mapping in
the southern Wind River Range (Bayley and others, 1873), Seminoe Mountains (Bayley, 1968), and Black J3Ik (Bayley, 1970).
Geologic studies in the Wpming province expmded
considerably during the 19708 and 1980s because of scientific
interest in the early history of the earth, economic interest in base
metals and uranium, and the legal requirements for mapping and
evaluation of mineral resource potential of wilderness areas. The
Wind River Range, the largest Precambrian-ored uplift in t&
province.
The U.S. Geological Survey has also sponsored general geg.
logic studies and geniogic mapping in the M
e Up*'
(Snyder, 1980), mtral h a a h Moc131tahs (Snyder, 19841,
Caper Mountah (Gable and others, 1987), Bighorn Mountaim
(Barker, 1982), and Granite Mountains (Peterma$ and Hildreth, .
1978), md presented a regional Precambrian summary in Hedge
and others (1986).
Much new geologic and geochronologic information has
result& from topical studies in the 1970sand 1980s. In Montana
and northwestern Wyoming, significant contributions have been
made by Jmes fl981), Page (1977)) Reid and others (1975),
Vitdiw lrsd &ma (1979), %lev (1983), Rowan and Mueller
(1971), and W e v i c h and othm (1981). An excelllartt review
of recent geoldc and g~ockonolagicwork in tbb area is in
Mueller and Wooden (1R2). In the Bl& Hills erf %nth Dakota
and adjacent a m , t
kd e t d ~
by IUdn,kopfdReddm
(1975) and Radden (1980) I&=to rmewed ~ ~ p p l iof
l gthe
entire upW W. D. Ikwd dh
h e y dW y o d g
is currendy mapping h R e a m of the Owl Crgek
Momtak d %'yam&& wid part5 af the southern S i m Madre
and Laramie Rbuntdm are being mapped by students and staff
of the U 6 v d Of
~ Wyoming under the sponsorship of the U.S.
Geological Survey.
As of 1987, appr+tely
70 percent of the ex@
Pre
caahh
been mapped,at least in reconn-%
and mmy
arms h e
mapped in detail. Major gaps remain in s ~ b
urn a th.% aorthern Laramie Momains, Bearrm& Mountains
southern Bighorn and Owl Crwk Momaim (Plate 2); as is true
of d regions of a m p k gdogg, many ~VMSwill b e d i t .from
more detailed mapping the &&re. We are apprombhg a period in the study of the Wyoming provine a r b pioneeris%
geologic mapping is nearly complete and d M integraM
geologic, geocho3ical, and geochronological mdes can be Wdertaken to help solve the problem of origin of W e interesting
rocks.
[th
':
h
r tb
Wyoming province
ORGANIZATION
!
i
r
To interpret and clarify much new information, we discuss
the Precambrian of the Wyoming province by reviewing the
geology of individual Laramide uplifts. Archean rocks are examined first, followed by a discussion of Proterozoic rocks along the
margins of the province. The section on Archean rocks proceeds
generally from northwest to southeast and represents a semicontinuous northwest-to-southeast transect across the Wyoming
province. This @ followed by discussion of probable allochthonous Archean rocks of the southwestern Wyoming province and
other isolated Archean outcrops whose relationship to the main
outcrop areas of the Wyoming province are not well understood.
The discussion of individual uplifts emphasizes gneissic
basement and overlying rocks in order to reconstruct as much as
possible of the long and complex tectonic history. Felsic intrusions of Late Archean age are not emphasized, although they are
the most abundant Archean rocks exposed in the province (Plate
2). These rocks range in composition from quartz diorite to alkali
granite and have been variously interpreted as recrystallized volcanic rocks, intrusive rocks of unknown source, products of partial melting, and products of in situ granitization. Most recent
workers no longer accept the granitization concept proposed by
Poldervaart and coworkers (Eckelmann and Polde~aart,1957;
Harris, 1959; Larsen and others, 1966) for the eastern Beartooth
Mountains and by Oftedahl (1953) for portions of the Wind
River Range, and consider each individual pluton as a specific
genetic problem.
Granites of Late Archean age intrude the Middle or Late
Archean gneiss in all the areas discussed below. They range in age
from about 2,800 Ma to about 2,600 Ma; some are post-tectonic
batholiths and stocks that are readily mappable as discrete bodies
and yield reliable isotopic ages; others have gradational and migmatitic borders and might be confused with older felsic bodies but
for their isotopic ages. It is possible that much more of the areas
of older g n e k are affected by Late Archean magmatism than is
indicated by current mapping. Mueller and others (1984) suggest
that Late Archean granite underlies more than 90 percent of the
eastern Beartooth Mountains. Just how much of the gneissic part
of the province has been melted to form granitoid masses is not
known. Migmatite and magmatite are common in all areas, but
their origins are uncertain.
ARCHEAN ROCKS
BEARTOOTH MOUNTAINS AND SOUTHWEST
123
tary suites of both miogeoclinal and eugeoclinal affinities (Fig. 1).
The Archean basement exposures are bordered by Proterozoic
clastic sedimentary rocks of the Belt Supergroup and its correlatives (Chapter 6).
Archean rocks are exposed in the cores of Laramide foreland uplifts that have been segmented by later normal faulting.
The excellent exposures in the ranges are the result of as much as
two kilometers of vertical relief. However, the exposures are discontinuous due to flanking and infolded Paleozoic and Mesozoic
sedimentary rocks, Eocene to Recent volcanic rocks, and Neogene basin 611. The lack of continuous exposure and the uncertain
translations on Laramide and Precambrian faults have hindered
the development of regional correlations and tectonic models for
the evolution of the Precambrian basement.
With the exception of flanking areas of Proterozoic clastic
rocks, mafic dikes, and isolated Cretaceous to Recent intrusive
complexes, Rb-Sr and U-Pb age determinations indicate that the
basement of southwestern and south central Montana is entirely
Archean (Peterman, 1979). However, Giletti (1966, 1971) documented a heating event that reset K-Ar systems in the Early
Proterozoic. He originally postulated a narrow transition from
Proterozoic to Archean K-Ar ages, which has been called "Giletti's line" (Fig. 1). However, the K-Ar age transition is actually
quite diffuse. Isotopic reequilibration is also probably a function
~ r occur on both sides
of deformation, for Archean 4 0 ~ r - 3 9ages
of a Proterozoic shear zone in the southern Madison Range
(Erslev and Sutter, 1990).
In the northwestern Wyoming province, Archean rock assemblages can be divided into the following associations on the
basis of lithologic similarities: (1) heterolithic gneiss complexes
dominated by granitic to tonalitic g n e k , with interlayered units
of amphibolite, metasedimentary rocks, and ultramafic pod.,
(2) pelitic and psammitic metasedimentary suites dominated by
semi-pelitic schists and gneisses, with associated quartzite and
iron formation; (3) marble-bearing metasedimentary suites, with
extensive thicknesses of dolomitic marble, semi-pelitic schist,
green quartzite and amphrblite; (4) foliated granitic plutons containing large inclusions of older rocks; and (5) the Stillwater
Complex of the Beartooth Mountains, a layered mafic intrusion.
These associations have been used as the basis for regional correlations (Erslev, 1983), although the possibility of multiple cycles
of lithologically similar suites has yet to be rigorously explored.
The fact that rocks exposed in southwest Montana consist
largely of the first three associations while those in the Beartooth
Mountains mostly belong to the last two suggests that these areas
are parts of different geologic domains. These areas will be described separately, with possible regional correlationsdiscussed at
the end of this section.
Beartooth Mountains
The Archean basement of the Beartooth Mountains and the
m g e in southwest Montana is a poorly understood collage of
@dSpathic gneiss complexes, granitic plutons, and metasedirnen-
The Beartooth Mountains (Fig. 1) are the largest exposure of
Archean rocks in the northwestern Wyoming province. Even
though large areas of these Archean rocks have never been
R S. Houston and Others
EXPLANATION
Foliated Granitic Rocks
Mef osedhenf ary rocks;
marble shown in red
Tonolitic migmotite-gneiss
complexes with u/fromafic
pods morked with 'b"
Limif of Premmbrian exposure
-
Confocfs ond fo/jofion trends
..er*+C~
0
Myfonific Rocks
/
SCALE
50 KILOMETERS
Figure 1. Generalized map of Archean rocks in southwestern Montana and northwestern Wyoming.
Fault
mapped beyond the recam&mw level, their geab.e&tq and 1957; Gcst .ando&&s,1958). Howewer, other o r & h were 8%(1969)?iwnd Cb&a d &hers
gmhronolagy is probably the be$ c:haraeteriZed af afI A r ~ h m ~gmkd by Wbr (I%@)&
rocks in the northwestern Wyoming prowinm.
h d west of the wwmnmmt Wwatm Complus~a d t
b go& %kM
dth Bmmth Plateau block, T?leza~ksk tl& atea,as
Snowy block, whhh m&m& mu& into Wyom4ng. The Fk&i%&%h-1
by M a (19@), Butler (1JBX Skimer (19691,md
abundant xenditb of aldm material. The Norfh and &ah to
Snowy blwb contain a much U e r praportion sf rntasdknenttuy rwlns and laymed gwks complexa AIl b e e blaoks
con* similar lith010gi
the end of the Ar~hpan.
contains two of the bes~tudiedareas in the Wyamhg pbolviim
exposum adjamat, to U.S. Hi&w&y 212 (the I3wt& I%&way) in the southeastern Betrrboth MWins,and the Stillwater
Complex in the noruheas;tern flank of &e range.
Intensiveinwestigazimsin the aea around the higkw~ywere
interest was the
d tbe &tic mb.Early workws mi&
gated that the i n h & intermixtares d grapitic:and &mrl i h l ogjies were due to granitizsction (EckeImam and Pddervart,
@& g i m k d o g h l sQ&es by Wooden and
&dueller aad othm (19821bwe 5 6 m t b t the
gm$tic & a e koriginated as &wive r m b emat approximately "2.75 to 2.80 Oa (Rb&r whole rock and U-Fb &COB
~a).
h1wof diorith ~~~~ph-i'bolli&
intespreeed as older wl-
ot
n a;&% (P. A. Mndlw, p
~ t z mdm d d o n , 19841.
Skinner and o h (19m mmd'this 2.8
rnogbc pibe the
B e a w orageay. By 2.70 Ga,the post-orogenic Stillwater
,
I
.
h
I
Wyoming
amounts of multiphase cumulates and noncumulate mafic rocks.
Disseminated and massive copper-nickel sulfides are concentrated in the cumulates near the sharp basal contact and in the
hornfels, which contains the precursor sills and dikes.
The overlying Ultramafic series is about 1,070 m thick and
is composed of cumulates of olivine, bronzite, and chromite; it
averages about 1,070 m in thickness. Its upper contact is marked
by the appearance of cumulus plagioclase. It is divisible into two
zones. The lower Peridotite zone contains largescale, modally
and texturally definable stratigraphic units that show regular cyclic repetitions. Within a welldeveloped cyclic unit, olivine
cumulate containing a chromite seam grades upward into an
olivine-bronzite cumulate which grades into a bronzite cumulate.
Not all cyclic units are complete, and the number of units range
from as few as 8 to as many as 21. 'Jbe disappearance of olivine
marks the sharp contact with the overlying Bronzitite zke, which
contains a succession of laminated and size-graded bronzite cumulates. The uppermost part of this zone locally contains thin,
discontinuous layers of olivine and chromite cumulate.
The first appearance of cumulus plagioclase marks the base
of the Banded series, which makes up more than three-fourths of
the exposed complex. Although several stratigraphicsubdivisions
of the Banded series have been proposed, it can most conveniently be divided into the Lower, Middle, and Upper Banded
series largely on the basis of crystabation order of the ten to 14
zones (McCaUum and others, 1980). The crystallization order
throughout much of the Middle Banded series is olivineplagioclase-augite-bronzite. This contrasts sharply with the crystallization order of olivine-bronzite-plagioclase-augite, which is
typical of the Ultramafic, Lower Banded, and Upper Banded
series. The Lower Banded series, from bottom to top, consists of
units of plagioclasebronzite and plagioclase-bronzite-augite
cumulates overlain by a complex sequence of olivine-bearing
cumulatesthat contain the platinum-bearing J-M reef. The reef is
overlain by units of plagioclase-bronzite, plagioclase-bronziteaugite, and more olivine-bearing cumulates. A thick plagioclase
cumulate defines the base of the Middle Banded series and it is
overlain by a sequence of olivine-bearing and plagioclase cumulates. The upper part is another thick plagioclase cumulate. Overlying this is another sequence of olivine-bearing cumulates, which
form the base of the Upper Banded series that is composed of
cumulates of bronzite, augite and plagioclase followed by cumulates with inverted pigeonite in place of bronzite.
Rocks adjacent to the Stillwater Complex have been extensively studied because of economic interest in the banded iron
formations in the contact aureole. Butler (1966) mapped an area
south of the complex that consists of granitic gneiss and migmatite, biotite schist, metasedimentary homfels, and unfoliated
Mouat Quartz Monzonite. The tightly folded amphibolite-facies
gneisses and schists are similar to those in the southern part of the
Beartooth Plateau block. Isotopic ages range from 2.28 Ga K-Ar
forms the lowest 100 m of the complex and and Rb-Sr determinations (Gast and others, 1958) to 3.1 to 3.3
rally continuous, but with local heterogeneities. It is com- Ga U-Pb determinations on zircons from the biotite schists (Capredominantly of bronzite cumulates but contains minor tanzaro and Kulp, 1964; Nunes and Tilton, 1971). The hornfels
Complex was intruded at much higher structural levels, suggesting that the Beartooth orogeny was a relatively short-lived period
of granitic intrusion, deformation, and uplift. K-Ar ages ranging
from 2.28 to 2.53 Ga (Gast and others, 1958) suggest either a
prolonged cooling episode or later tectonic events following the
intrusion of the Stillwater Complex.
Xenoliths of metasedimentary units, ultramafic rocks, and
basaltic amphibolite in the Late Archean granitic complex appear
to be relicts of an old polymetamorphic terrane. In the center of
the complex, Tim (1982) obse~edan early garnet-cordieritebiotite-andalusite assemblage overprinted by contact metamorphic assemblages associated with a series of granitic plutons.
Farther east, Skinner (1969) reported multiple metamorphic assemblages ranging from granulite facies (spinel-olivine-orthe
pyroxene-hornblende) to amphibolite facies (anthophyllite-hornblende-chlorite) to yet lower grade serpentinite. Henry and others
(1982) analyzed pyroxene-bearingsupracrustal assemblages from
the Quad Creek area that indicate granulitefacies conditions of
approximately 780°C and 6 kb. Rb-Sr wholerock analyses of
these rocks yielded an age of 3,390 55 Ma (Sr initial ratio
0.6991), which they interpreted as the age of granulite-facies
metamorphism.
Stillwater Complex (N. J. Page, M. L. Zientek, and E. A.
Erslev) The Stillwater Complex is a Late Archean stratiform
mafic and ultramafic intrusion that crops out for about 48 km
along the northern border of the Beartooth Mountains (Fig. 1).
Igneous layering in the complex strikes northwest and dips
steeply, so a section across approximately 5,500 m of the layered
intrusion is exposed. Gravity and magnetic anomalies suggest that
the complex extends beneath Phanerozoic sedimentary rocks in
an area of about 4,400 km2 to the northeast (Bonini, 1982).
Sulfiderich rocks were discovered in the Stillwater River valley
in 1883. Since then, the complex has been the focus of many
scientific studies and extensive exploration for Cr, Ni, Cu, Pt, and
Pd. A summary of recent studies is provided by Czamanske and
Zientek (1985), and much of the following description is based
on papers in that volume.
The complex was emplaced in Middle or Late Archean
clastic rocks including pelites, quartzite, and iron formation. The
wallrocks were complexly folded prior to 2.70 Ga, when the
basaltic magmas of the Stillwater Complex were intruded, forming a metamorphic aureole. During early stages of cooling, minor
tectonism caused irregularities in the basal contact and formed
basin-like structures within which the complex crystallized.
Crystallization of the basaltic magmas produced a layered
sequence of ultramafic and mafic cumulates consisting mainly of
combinations of olivine, bronzite, plagioclase, and augite. These
can be divided into three major series: the Basal, UItramafic, and
/ Banded series. The precursor silk and dikes are largely diabase
end mafic norite, and can be divided into at least three petrologi-
*
I
PZand
c o d of a layered sequence of folded pelitic to panmitic
hornEels, blue quamite, metamorphosed dhmictite, agid izon
formath. Cfossbedding cut-and-fiU stmctws?and enrichment
in elmeots of ultramafic affinity suggest that the homfeIs are not
correWvewith the regionally metamorphosed mbists to the west,
which do not pmerve sedimentary structues (Page, 1977). Thus,
the protolitb of these homfels may be dochtho~ous
and
Geissnian, 1984) or deposited after regional metamorpperhaps in a basin formed during a rifting event that eventually
resulted in the emplacement of the Stillwater Complex.
The Stillwater Complex and the s m u n a g mtanorphic
complex are cut by stocks of M o d Quartz Momnite and bsser
mounts of quartz diorite. The quartz momni& is an unfobred
biotitekming rock with variable grain size. Nunes and Tilton
(1971) determined 2.75 to 2.70 Ga U-Pb ages on zircons from
both the Stillwater Complex and the Mouat Qua& M m n i t e .
More m a t NdSm geochronology by DePaolo and Wasserburg
(1979) yielded a 2,701 f 8 Ma age for the Stillwater Complex,
which may be regarded as a a d u m age for the Mowt
M o d % . C l d y , these two events must have been nearly synchronous. Thw intrusions may mywmnt a bimodal tholeiite
granite suite common in Phanerozoic rift envirdnmeuts.
Presedy, the Stillwater Complex is juxtaposed against the
Beartooth and Nurth Snowy blocks (Fig.1). At 2,700 Ma,quartz
m o d t e plutons were intruded into the complex and into the
m e between the Beartooth block and Stillwater Complex, Jib
tween 1,800 and 1,600 Ma, the area was involved in a weak lowgrade regional metamorphic event and a penetrative foliation
developed in the caqdex.
Elliott and others (1983). Reid and others (1975) and Mogk
(1983) did detailed geologbl and geochemical mdies in the
Pine Creek m a k dm mEral part d the block. The North
Snowy bWk 6onsLts of northeast-strikhg units of trondhjemitic
gneiss, interdated amphibolite and q d t e , mylonitic schist,
isdinally folded supracrustal rocks, heterogeneous paragneb
and orthogneiss, and granitic to granodioritiie aqen gn&. Reid
and others (1975) originally interpreted the Pine Cr& area as
one large refolded nappe with a core of marble, quartzite, and
arnph~hlite.However, Mogk (1983) has shown that the gneissic
units are not symmetric about the memedimentary rwks in the
core of the fold, and sqgested that the structure is patt of a series
of east-vergiag thrust she& that form a duplex of ciys%llhe
rocks.
Reid and others (1975) proposed a detailed chronalogy on
the basis of U-Pb, RWr, and K-Ar ages determined by commercial laboratories. More rigorous R W r isochron and Sm-Nd
model ages an the trondhjmitic and heterogeneousg n e k indicate that these rocks may be Early Archem (Mogk, 1983). The
Mount Cowen Augen Gneiss in the m t e r of the block yielded a
seven-point Rb-Sr isochron d 2,730 f 70 Ma (initial Sr ratio
0.7026 0.0014), similar to the age of plumns in the lhrtooth
Plateau block.
*
Ofhers
The Rb-SI and K-Ar ages of biatites' from
rocks reported by Reid and others (1
Ma This indicaks that the Early
affected rocks comidenrbiy east of
fined by Giletti (1971) md &own in Figure 1.
Conditions of
hic equilibration range from fie
and a3mphEwIite Eacies ia
boundary btww@
strained gneisw a
d mylrmtic schists to upper a m p b b i
in m~~
rocks and heterogeneous~ekesto the
(Mogk, I!%&). Moderate pressure c
o am indicated
~
by~
kymite owz~grownby fibrolitic sillin&&. AssnCiated stsu~~urw
include at leqt two stages of duotile frhihg amound axes &at
pmeatly trend northeast, followed by the development of
northwest-dipping myZonitea ernd by upright open folding of adjacent units mund sub-tal
axm.
The complex array of K-Ar ages d metamorphic e
q
a
b t b n temperatures, and the common mylonitic textures s u g ~
that 1-t
Archm and homozoic events affected the North
Wwy block mote: tkaz14nteastern ~ 0 0 t Mountains.
h
Reid .
and others (8975) rewgaked five deformatios on the basis of
structural ge&n&q a d h b d geochmnobgy. Erslev (1982)
suggested,that 1-r m n t s of mylonithiion, greenschist- to
amphibolitefacies metamorphism and open folding were a r
s .
spouse to motion on a regional zone of duetile shearing. Mogk (1983)
that the different metamorphic domaim were
juxtqmd durisg the ArIw along a sekes of &rust zones thett
are now r e p m n t d by h b w c ~~tltwti
Mope
. isotopic and
stmiwal inf~nogationis needed to dekmine the age and nature
of tectonic events in the North Snowy block.
I
,
I
I
I
block where biotite schists are interlapred with iron formations
that locally contain eoonomic gold deposits. Elsewhere, the rocks
cooist largely of an assortment of poorly studied g n e k and
wsme National Park
qwrtzi@ with &or
and silicatefacies badad
The layering and fokand dip w t . Graded
cates that approximat& &If of the beds werecoverNrned in an early ljhase of isbk&agSubsequent
.
deformation has f~fW
the racks twiee, with the later stage of more open
folding occunring about northeast-plunging fold axa and
northwest-dipping axial planes.
The amplhlite-hies, andalus&bearia& biotite-garnet
schists are tr&rmed
to sillimanitshring migmatites wound
granitic plutom. 'Pfiese intmives range &om twbmica granites
and p d o r i t w with sharp contacts to earlier, crudely foliated,
homblmde quartz didtes. Various isatnpic age debminatiom
on these intmsiolls by Brookks (1968), Montgomery (1982), and
Montgomery and Lytwyn (1984b) all cluster around 2.7 Ga
I
~
I
Wyoming
I
2
of high-grade gneiss have been uplifted on northwest-striking
faults. The gneissic and metasedimentary rocks of the southern
Madison and Gravelly Ranges were once part of a single Laramide thrust block that has been split by north-striking normal
faults. Isolated exposures of similar lithologies occur in the Gallatin Range near the northwestern corner of Yellowstone National Park (Witkind, 1969) and in the central part of the
southern Madison Range (Hadley, 1969a).
Spencer and Kozak (1975) describe the gneisses in the
northern Madison Range as plagioclase-quartz gneiss and
microcbplagioclasequartz gneiss with abundant migmatitic
textures. Ultramafic pods are common, locally with corundumbearing assemblages produced by reaction with siliceous wall
rock (Bakken, 1980). Minor areas of marble, muscovite schist,
and quartzite locally contain either kyanite or sillimanite. In the
western part of the range, McThenia (1960) reported ionalitic
migmatitic gneiss containing variable amounts of biotite, hornblende, and garnet interlayered with banded amphibolite and
quartzite containing green muscovite. D. W. Mogk (personal
communication, 1984) has found relict pods of granulite-facies
gneisses in this area. The Archean rocks of the northern Madison
Range have layer-parallel foliation with a uniform strike to the
northeast and variable dips. Spencer and Kozak (1975) interpreted the dip reversals as a series of northeast-trending synforms
and antilorms. These structures refold an earlier fabric containing
transposed isoclinal folds, amphibolite boudins, and folded
pegmatites.
Giletti (1971) reported whole-rock Rb-Sr ages of approximately 2.9 Ga for g n e k in the center of the range. James and
Hedge (1980) included rocks from the northern Madison Range
with some collected from the Tobacco Root and Ruby Ranges in
a composite RbSr isochron that gave an age of 2,730 + 85 Ma.
They concluded that all of the rocks are of essentially the same
age. However, it should be noted that the samples from the
northern Madison Range gave a mean Rb-Sr model age of 3.01
Ga (Sr initial of .701) as compared to a mean model age of 2.76
Ga for the other samples (omitting anomalous 5.1 and 4.7 Ga
ages). Thus, the analyzed rocks from the northern Madison
Range may be significantly older than those from the Tobacco
Root and Ruby Ranges.
The southern Madison Range consists of a central gneiss
complex flanked by marble-bearing metasedimentary sequences
to the north and south (Erslev, 1983; Hadley, 1969a). The gneiss
complex consists of interlayered heterogeneous migmatites,
tonalitic gneiss, amphibolite, and minor metasedimentary lithologies. Erslev (1983) suggested that the gneisses were derived from
a sequence of volcanic rocks and intercalated sediments. The
gneisses wrap around a dome of granitic to dioritic augen gneiss.
Sills and irregular lenses of poorly foliated granite gneiss crosscut
the earlier migmatite complex. A major zone of ductile shearing,
Madison and Gravelly Ranges
the Madison mylonite zone, (Plate 2; Fig. 1) occurs along the
The next major exposures of Archean basement west of the southern contact between the gneiss complex and the metasedirtooth Mountains are in the cores of the Madison and Grav- mentary rocks (Erslev, 1982).
Ranges (Fig. 1). In the northern Madison Range, two blocks
At least three metamorphic equilibrations have been recog-
Montgomery and Lytwyn (1984a) showed that this is also the
approximate age of amphibolite-facies metamorphism that was
overprinted by a 1.8 Ga thermal event that increased in intensity
to the northwest. The later event may be related to ductile shearing along the northwest margin of the North Snowy block
(Erslev, 1982).
North of Yellowstone National Park, geologic studies have
attempted to delineate and understand the origin of the gold- and
arsenopyrite-bearing iron formation that was mined from 1882 to
1947 (Seager, 1944, Hallager, 1980). Adjacent metasedimentary
lithologies are identical to those in the park. Hallager (1980) reported matrix-supported conglomerate with clasts of quartzite
and biotite schist, suggesting earlier Archean crust in the source
region. On the basis of graded bedding, petrography, and mineral
chemistry, he documented an early phase of isoclinal folding,
flattening, foliation development and lower-grade metamorphism,
followed by folding about upright northwest-striking axial surfaces. Both of these events have been overprinted by crenulations
and large-scale folds whose axial planes strike northeast and,
to the northwest. Compositional data indicate a lack of volcanic
units and the association of gold with enrichments of Mg, Cr, and
Ni in the sediments (Hallager, 1980).
The northwestern corner of the block consists of a separate
exposure of heterogeneous tonalitic to granitic gneiss, amphibolite, quartzite, pelitic schist, and mylonitic equivalents. Foliations
strike northeast and dip northwest (Mliott and others, 1983),
parallel to zones of mylonitization (Erslev, 1982). This major
zone of shearing coincides with the northwestem boundary of the
Beartooth block due to the reactivation of shear planes during
Tertiary normal faulting. The ductile shear zone contains epidote
amphibolite-facies phyllonitic schists, amphiiolites with actinolite-epidote rims on hornblende augen, green quartzite, and porphyroclastic mylonitic gneiss. Amphibolite-facies schists containing andalusite north of the shear zone near the westemmost
comer of the South Snowy block are highly folded, with axial
planes parallel to the shear zone. South of the major shear zone,
zones of protomylonite cut tonalitic to granitic orthogneiss and
sillimanite-bearing pelitic gneiss with a complicated isotopic history possibly extending back to 3.6 Ga (Guy and Sinha, 1985).
Preliminary fabric analysis in this area indicates a combination of dip-slip and strike-slip motion. In some places, fabrics
in amphibolite-facies rocks show a subhorizontal stretching lineation consistent with a strike-slip motion. These rocks are cut by
mylonites of the upper greenschist facies with asymmetrical feldspar clasts indicating dipslip motion of a normal sense. The age
of shearing is uncertain, but it appears to be correlative with Late
Proterozoic isotopic disturbances throughout the Snowy blocks
and southwest Montana.
128
R
S. Houston1 and Others
nized in the central gneiss complex. Megacrystic bronzite coexisting with hercynite spinel in ultmm&c pods, and garnet-bearing
two-pyroxene assemblage in mafic granulites represent relict
granulite-facies assemblages. However, most rocks have been retrograded to upper amphibolite-facies assemblages. Epidote
amphiilite-facies assemblages occur adjacent to the Madison
myloniie zone and other northeast-strkhg shear zones.
U-Pb zircon and RbSr wholerock analyses of tonalitic
gneisses from areas of minimal later deformation and retrogradk
metamorphism yield Middle Archean ages (P. A. Mueller, personal communication, 1984). 40Ar-39Arspectra of hornblende
from one of these samples indicate a 2.5 Ga cooling age.
South and north of the gneiss exposures in the southern
Madison Range a supracrustal sequences of dolomitic marble,
biotitsstaurolite-garnet schist, quartzite, and amphiilite is exposed. The sequence in the southern part of the range is continuous with units on strike to the south and southwest (Witkind,
1972,1976; Sonderegger and others, 1983) and is correlated on
the basis of lithologic sequence with the Cherry Creek Metamorphic Suite in its type area in the Gravelly Range.
The Cherry Creek Metamorphic Suite in the southern end of
the Madison Range shows several clear stages of deformation and
metamorphism. The earliest stage of middle to upper amphiilkfacies metamorphism (60OQC,>6 kb) is associated with recumbent folding and northeast-vergent thrusting (Erslev, 1983).
These rocks cooled below the argon retention isotherms for
hornblende by 2.7 Ga. Intrusion of granodiorite sills, dated at
2.59 Ga by the U-Pb zircon method (P. M d e r , persooal communication, 1987) locally reset the 40Ar-39Ar sysbms. The
second stage of epidote amphibolite-facies metamorphism was
centered around the Madison mylonite zone where shearing
and retrograde metamorphism coincided (Erslev, 1982). Within
the ductile shear zone, a heterolithic, highly strained metaconglomerate separates mylonitic metawhentary rocks from
myloniric g n e k . This horizon may represent a nonconformity
where the Cherry Creek sediments were deposited on precherry
Creek granites.
The Madigon myionite zone is a 3-km-thick ductile shear
zone dipping moderately to the northwest. Lithologic units in the
zone are highly strained but not dismembered. Reverse move
ment is indicated by rotated foliation in adjacs&trocks, down-dip
stretching, and asymmetrical shear fabrics. Sptemnic garnetchloritoid-biotite assemblages and epidote-baring amphrblites
indicate shearing in the lower amphibolite facia. met-biotite
geothermometry indicates temperatures of about 500°C. Adjacent rocks underwent major argon loss, with 40Ar-39~r reset
to approximately 1.9 Ga
In the Gravelly Range, north- to northeast-striking units of
metasedimentary rocks alternate with domains of granite gneiss.
No major units of tonalitic gneiss or migmatite have been r e
ported. The metasedimentary
in the southern part of the
range consist of feldspathic biotite-staurolite schist, quartzmuscovite schist, amphibolite, and grunerite iron formation. The
sequence is cut by metagabbro dikes and by sills associated with a
large plutan of granite gneiss (Wier, 1965; WB
The metasedhentary rocks are m o d d y
cross-bedding and a simple stratigraphic sequence.
amphibolite-facies assemblages in the main body of biotib
staurolite schist contain andalusite. These rooks gs&e inm
sinimanitegarnet-staurolite-biotiteschistsadjacent to the gneii@;
&te intrusion that bounds the schists to the north, Green8cbtand epidote amphiilite-facies assemblages commonly overprint
the earlier equilibrium assemblages.
North of the central granite gneiss is the type section of the
Cherry Creek Metamorphic Suite as orighdy defined by Peak ,
(1896). Extensive mapping in this area by Peale (1896), Heinrich
'
and Rabbitt (1960), Hadley (19b) and ~
o (1976) ~
has revealed a well-stratified sequence of biotite schist, dolomite
marble, quartzite, hornblende gneiss, kydtsstaurolite schist,
iron formation, phyllite, and anthophyllite gneiss. Tbe northern
and southem boundaries of the type section of the Cherry Creek
Metamorphic Suite are marked by sharp metamorphic transi~
Milbolland (1976) suggested that
tions and m y 1 0 textufes.
these contacts may have been the sites of kctonioemplacement of
solid yet hot p d t i c sheets; which t h e d y and dynamically
metamorphosed the sedimentary units. Recent studies have
shown that the prograde amphibofite-facies assemblages have
reequilibrated to greenschist-facies~ssernbkigesduring southeastvergent ductile shearing at the base of fold n a p ,
In the northem part of the Gravelly Range, Heinrich and
.
Rabbltt (1960) and Iiadley (1%9a) mapped a thick unit of
1
granite gneiss hounded to the north by marble-metasedi,
mentrtry rock. These consist of calc-siliate marble, pelitic schist,
kymite-wet schist, and anthophyIlite gneiss. They appear to be
an incomplete higher-grape part of the Cherry Creek Metamorphic Suite. Biotitegarnet geothemometryindicates temperatures of approximately 600°C (MiUhoIlaad, 1976). Abundant
kyanite, mmmonly rimmedby oillimmiits, suggests high pressures
during progrde metamorphism followed by lower pressure
1
equilibration, which may correlate with the low-grade retrograde
I
metamorphism and mylonitization to the south.
1
Ruby und Gmnhorn Rcurgm
The Greenhorn Range spans most of the gap between the
Ruby Range to the west stnd the Tobacco Root Mountains to the
north (Fig.1). The low-lying expures of Archean rocks in the
G r e ~ n h mRange have been mapped by Hadley (1969b), Tilford
(1978) and Berg (1979). These maps reveal a complex array of
metsedimentary rocks and quartzo-feldspa&ic gneisses with diverse orientations of foliations and fold axes.
Berg (1979) described a variety of rock types of generally
higher grade than those in the Gravelly Range. The most abunc
of fonalitic to granitic
dsnt units are q ~ f e l d s p a t h i gneiss
c o m e o n and an amphiilite assemblage containing equigranu1a amph~"bolite,quamite, anthophyllite gneiss, and pyroxene granulite. Meamorplmed ultramafic bodies contain olivine
enstatite assemblages partially retrograded to serpentine, talc, and
I
1
Wyoming
,
,
t
'
tremolite. Concordant layers of calc-silicate marble with minor
quartzite occur in both the quartzo-feldspathic gneisses and the
amphibolite assemblage. These rocks probably represent several
lithostratigraphic suites, but inadequate exposures and complex
relationships have obscured the geologic history. The discontinuous and chaotic array of lithologies and structures could have
been generated in a deformed plutonic terrane with roof pendants, giant xenoliths, and country-rock screens of older
lithologies.
The trend of increasing metamorphic grade to the west continues into the granulite and upper amphibolite-facies rocks of
the Ruby and Blacktail Ranges. Extensive mapping in the Ruby
Range has been summarized by Karasevich and others (1981).
The Blacktail Range exposure was mapped and briefly discussed
by Heinrich (1960). The stratigraphy and structure of the Blacktail Range parallels that in the southern Ruby Range, where the
layering and schistosity of the rocks strike northeast.
Heinrich (1960) subdivided the northwest-dipping units of
the southern part of the Ruby Range into three lithostratigraphic
domains: an eastern strip of pre-Cherry Creek rocks, a central
strip of Dillon Gneiss, and a western strip of Cherry Creek lithologies. This subdivision has been followed by later workers despite
considerable disagreement as to the relative age and genesis of the
Heinrich (1960) descriied the pre-Cherry Creek rocks in
the southern exposures as a sequence of coarsegrained banded
gneisses of diverse compositions, and minor biotite schist and
sillimanitegarnet gneiss. Mineral assemblages in ultramafic pods
indicate equilibration at approximately 7 10°C and 5 kb (Desmaria, 1981). The structural position of the pre-Cherry Creek
rocks, combined with their structural complexity, led Heinrich to
suggest that this unit is considerably older than the overlying
Dillon Gneiss and Cherry Creek Metamorphic Suite. Karasevich
and others (1981), however, found no evidence of a separate
deformation or metamorphism in the central Ruby Range preceding the deposition of the Cherry Creek sequence. The fact that
field mappers have been unable to subdivide the pre-Cherry
Creek rocks suggests more structural complexity relative to the
easily mapped Dillon and Cherry Creek lithologies.
The Dillon Gneiss forms the backbone of the Ruby Range.
ts of amphiilite-facies gneiss of graHeinrich, 1960). Leucocratic units are nonornblende-bearing units are commonly con(Garihan and Williams, 1976). James and
d Rb-Sr model ages on Dillon Gneiss rangm 2.66 to 2.82 Ga. Heinrich (1960) proposed a synkineintrusive origin for the granitic sheets. However, Garihan
kuma (1974) noted intercalated layers of dolomitic marble
ng as 6 km separating units of granite gneiss. They suggested
the protolith of the Dillon Gneiss was an arkose. In view of
ility of an arkose-limestone sedimentary
Karasevich and others (1981) proposed an "illire(?)tone or siltstone" assemblage as a protolith. A major
cation of protoliths arises because all
potassium feldspar-bearing gneisses have been identified in the
field as Dillon Granite Gneiss. Units mapped as Dillon Gneiss
include thin layers of sillimanite-bearing metapelite as well as
irregular pods of gneissic granite with apparently magmatic
segregations of tourmaline. There may also be pre-Cherry Creek
rocks of granitic composition that have been grouped with the
Dillon Gneiss.
Cherry Creek lithologies include calc-silicate marble, calcsilicate rock, aluminous pelitic gneiss, biotitegamet gneiss,
quartzite, anthophyllite gneiss, layered hornblende gneiss, iron
formation, and large irregular bodies of equigranular amphibolite.
The homogeneous amphibolite occurs as dikes and as large sills
with local ultramafic zones. Geochemically, the homogeneous
amphibolite closely resembles hypersthene-normative continental
tholeiites with little or no sedimentary contamination (Husch and
Hennigan, 1984). These rocks are highly folded, forming largescale domes and basins in the southern part of the range.
Karasevich and others (1981) postulated an early stage of
isoclinal folding followed by as many as three generations of
more open folds, possibly generated during progressive deformation accompanying the latestage emplacement of nappes.
Multiple phases of intense metamorphism pose a major
problem in determining the stratigraphic and structural history of
this area. Prograde equilibration occurred above the second sillimanite isograd throughout the range. Geothermometry on iron
formation assemblages indicates temperatures from 675OC to
745OC (Immega and Klein, 1976). Relict kyanite in these
sillimanitebearing rocks suggests pressures from 6 to 8 kb, quite
high for amphibolite-facies rocks in the Archean. The regional
Early Proterozoic K-Ar resetting event (Giletti, 1966) has been
correlated with the formation of cordierite coronas on garnet and
sillimanite in the central and northern Ruby Range (Dahl, 1979).
Various geothermometers and geobarometers indicate conditions
of equilibration in this later event at approximately 550' and
4 kb.
TobaccoRoot and H i g h h d Ranges
Feldspathic gneisses in the northeastern Highland Range are
deformed by east-west-trending fold systems, which probably
correlate with similar structural trends in the northern Tobacco
Root Range. In the southwestern and central parts of the Highland Range, late stage folds plunge to the southwest (Duncan,
1976; Gordon, 1979) and the gneisses include minor amounts of
marble, sillirnanite-garnet gneiss, and iron formation. Other
upper amphibolitefacies lithologies include metamorphosed
ultramafic lenses, amphibolite, and hornblende-plagioclasegneiss.
The Archean geology of the Tobacco Root Mountains is
discussed by Vitaliano and others (1979), whose complete map of
the range and explanatory text represent the integration of nearly
20 years of research. The following discussion of structure and
lithology comes largely from this excellent summary.
At least two sets of ductile folds are recognized: an early set
of north-plunging isoclines, and later, more open folds with
.
130
-
I
1
R S,Houston and Others
variable axes. The change in orientation of the axes of the
younger folds firom north-northeast in the south to east-southeast
in the north was attributed by Vitaliano and others (1979) to
stretching during the emplacement of the Laramide batholith in
the center of the range, However, James (1981) documented
north-northwest-trendingcrossfolds in the Copper Mountain r e
gion to the south, which may indicate that the diverse orientations
are due to superimposed fold sets, not high-level basement
ductility.
The most widespread rocks in the Tobacco Root Mountains
are granitic to tonalitic gneiss and.migmatite. These gneisses are
commonly interlayered with amphibolite, hornblende gneiss,
quartzite, and sillimanite =hist. Other lithologiesinclude laterally
continuous units of hornblende gneiss (locally pyroxene-bearing),
anthophyllitegedrite gneiss, green quartzite, and. aluminous
schists with kyanite, sillimanite, or both. Caldcate-bearing
marble and garnet gneiss units define several broad northnorthwest-plunging antiforms and synforms in the southern and
central part of the range. These units are rare in the northern and
eastern p& of the range, where more quartmfeldspathic and
hornblende gneisses with ultramafic pods predominate. Wholerock RbSr data indicate a metamorphic age between 2.61 and
2.75 Ga (Mueller and Cordua, 1976; James and Hedge, 1880).
Several geothermometers and geobarometers indicate that
the iron formations were metamorphosed at approximately
700°C and 5 kb (Imega and Klein, 1976). Vitaliano and others
(1979) suggested that these rocks underwent two stages of
metamorphism. The earliest event, at granulite to uppermost amphibolite facies, produced pyroxenebearing gneiss and iron
formations. A widespread s m d event caused reequili'bration in
the almandine amphibolite facies and formed kyanite-muscovite
schists and amphlile overgrowths on pyroxenes. Vitaliano and
others (1979) suggested that the second event lacally reached
granulite facies, with pyroxene rims developing on amphiboles
and garnets. Alternatively, they suggest that these two events
could a c W y be a single prolonged equili7,ration event.
Striking similarities in lithology and sequence among the
metasedimentary rocks ia the ranges in southwest Montana and
northwestern Wyoming suggest stratigraphic correlations. However, variations in grade of metamorphism and intensity of deformation, together with discontinuous exposure and lack of
geochemical investigations, make correlations controversial.
Peale (1896) divided the preBelt rocks of the northern Gravelly Range into the "Proterozoic Cherry Creek Series" and
"Archean Gneiss." Winchell (1914) and some later workers
iorrelated marble-bearing metasedimentary sequences in the Tobacco Root Mountains and Ruby Range with the Cherry Creek
Series. Adjacent gneissic units in the northern Tobacco Root
Mountains were interpreted to be an older sequence named the
Pony Series by Tansley and others (1933).
Many recent workers have considered the relative age and
correlations of the metsedimentary and gneissicsuites to be inde
terminate because primary features are lacking. Vitaliaao a d
others (1979) concluded that there are no significant metamorphic or structural ditkences between the Pony and Cherry Creek
and that "the designation of Precambrian rocks in any one area
Cherry Creek Group or Pony Series may be unsound." D&dties with lithologic correlations were recognized by McThenia
(1960), who noted that all of the major rock types in the type
area of the Pony Series also occur in the Cherry Creek. However,
the converse is not true.
Most of the stratigraphic debate has centered in the T o h w
Root and Ruby Rangm where the grade of metamorphism and
intensity of ddmnation are the highest in the region. In the
lower^ rocIra in the southern Madisan and northern Gravelly Ranges, units in the Cherry Creek Metamorphic Suite can be
correlated both by lithology and strawaphic sequence (Fig.2).
In the southern Madison Range, graded beds iildicate that tops
are consistently to the south, with localked thrust faults causing
minor stratigraphic repetitions. The northern contacts of the
Cheny Creek units in the northern Gravelly and southern Madison Ranges are at the same st~atigraphic
position. On the basis of
relative position indicated by stratigraphic tops and the presence
quartzite-dominated metaconglomerates, Erslev
of -S
(1983) sqgeseed thatthe contact may be an unconformity where
Cherry Creek wdbents were deposited on pre-Chary Creek
basemat, This interpretation is supported by evidence of more
compb deformation in the prdherry Creek rocks, by the
occurrence of granulites, and by Middle to Early Archean ages
found for some rocks north of the contact.
Correlation of the marblebearhg sequences in the Ruby
and Tobacco Root Ranges with the Cherry Creek Metamorphic
Suite is reasonable, considering the r h t y of thick Archean carbonate units.
The Cherry Creek lithologies in the Gravelly and southern
Madison Ranees are similar, but the sequence in the Gravelly
Range costains thin beds of iron-rich, cxosbedded quartzite,
whkb are lactabg M e r east. In the Ruby Range, thick units of
qtwtm-fddspaxhic Dillon CSneigs are interlayered with marble. If
the gneiss is &y
of ~e~
origin, as s~~
by recent
workers, then these u8i@have no equivdents to the east. This
Iithol~c
could have been generated in a fan-delta
en-nment,
where highly diseonhuous units of arkose and
limestone are common. The Ruby Range may have been the
locus of nearshore deposition of alternating and overlapping
arkosic and limestone beds, which laterally pinch out due to shifts
in distributary channels. This postukted fan-deltasequence could
have graded vertically md laterally to the east into carbonate
shelf deposits with intercalated quartz arenite and chert.
All ofthese lithologic observations, as well as geochemical
indicators discussed by James and Hedge (1980) and Gibbs and
others (1986), suggest stable shelf deposition for the Cherry
Creek ~ e t a m o r ~ bSuite.
i c The lithologic changes in the Cherry
Creek Metamorphic Suite suggest a source to the west or northwest. The rocks in the Ruby Range are interpreted as a mastal
fan delta comp1ex, and those in the Gravelly and southern Madi-
i
I
1
1
I
R S Houston Md Others
I
*
BIGHORN MOUNTAINS
2. E. Peterman and R.S. Houston
The Bighorn Mountains contain two major Precambrian
lithologic units (Heimlich, 1969). The southern half consists
mostly of quartzu-feldspathic gneiss; the northern half is largely
tonalite and granite. Minor rock types include amphibolite,
hornblende-biotite schist, muscovite schist, quartzite, gametEer011s gneiss, marble, calc-silicate rock, and metamorphosed iron
formation. Heimlich (1969) suggested that the gneisses and granitic rocks were derived from a sedimentary protolith by recrystallization and metasomatism. However, detailed studies in the
Lake Helen area in the southern Bighorn Mountains (Barker and
others, 1979; Barker, 1982) indicate that they are derived from
igneous rocks that have undergone multiple episodes of deformation and metamorphism. Barker (1982) identifled four major
units in the Lake Helen area The oldest unit, representing the h t
plutono-thermal event in the region (designated E-1 by Barker
and others, 1979), consists of foliated and banded trondhjemitic
gneiss with minor lenses and interlayers of amphlilite. The
younger event @2) is represented by a trondhjemitic pluton and
a granitic intrusive complex onsisting of equal amounts of granite and granodiorite with minor biotite-hornblende quartz diorite
and biotite tonalite. An agmatite of mappable proportions contains fragments of tonalite, trondhjemite, and minor amphlilite
in a matrix of granodiorite and granite.
Isotopic dating establishes a protracted history for the Pre
cmbrian rocks in the Bighorn Mountains. K-Ar biotite ages
show a correlation with rock type and position, the latter being
the key factor. Heimlich and Armstrong (1972) report ages of
2,730 Ma for tonalites and granites in the northern haif of the
range and 2,500 Ma for g n e k in the southern half. Peterman
(1979) suggested that the difference in biotite ages between the
northern and southern parts of the range is due to differential
uplift in the Late Archean. The age discontinuity is probably
coincident with a major northeast-trending shear zone, one of
many east- and northeast-trending faults and lineaments. The
eastern end of the age discontinuity is approximately coincident
with a tear fault at the north side of the Laramide Piney thrust
along the east side of the range. The Late Archean biotite ages in
both terranes attest to crustal stabilktion at this time. Rocks
presently exposed have not undergone any penetrative deformation or heating above 300°C since the Late Archean.
Rock-forming events extend back to the earliest Late Arohean or the latest Middle Archean. In the Lake Helen area of the
southern gneiss terrane, the two major events involving *atism, deformation, and metamorphism are well dakd (Arth and
others, 1980). Event E-1 is dated at 3,007 68 Ma by the
wholerock Rb-Sr method and at 2,947 f 100 Ma by the U-Pb
zircon method. The younger event (E2) is dated by the whole
rock Rb-Sr method at 2,801 f 62 Ma. The low initial 87sr/S6sr
ratios for the E-1 gneisses preclude the possibility of any significant crustal history prior to 3,000 Ma. Stueber and Heimlich
*
(1977) reported a wholerock Rb-Sr isochron age of 2,790 &O
Ma bawl aa a regional collection of samples from both the
northern and southern Bighorn Mountains. This age indicates that
the E-2 event was regional in scope.
U-Pb data for monazite aad zircon from a granite at the
northern end of the range give an age of 2,850 25 Ma (Hehlich and Banks, 1968). Banks and Heimlich (1976) report U-PI,
zircon ages between 2,840 and 2,865 Ma for granitic rocks in the
northern Bighorns and ages between 2,890 and 2 9 5 Ma, with
one age of 2,710 Ma, for gneiws in the southern Bighorns.
Both terranes in the Bighorns are cut by diabase dikes.
Stueber and others (1976) report Rb-Sr wholerock isachron ages
of2,?B6f58Ma~2,193i100Mafortwosetsofdikesinthe
northern &$horns. An internal Rb-Sr isochron for one of the
younger dikes is 2,154 f 35 Ma, which is consistent with the
wholerock isochron. Included in these suites were samples for
which K-Ar ages are as low as 1,880 Ma, suggesting that the
K-Ar wholerock ages are umehbIe.
The age of the Late Archean diabases is cl& to the average
K-Ar biotite age of 2,730 Ma for the ngrthern terrane, suggesting
that emplacement of the dikes c0inciBed closely with uplift and
o l i n g in the northern Biiorn Mountains.
+
*
'
.
'
OWL CREEK MOUNTAINS
R.S.H o d o n and K. E. KmM-orn
The Owl C r d Mountains include an eastern segment that
is largely gnekq a central m n t that consists of gneissic base
ment, suprmmtal rocks, and cross-cuttinggranite; and a western
segment that is mostly Late Archean granite with some s.upracmtal rocks (Plate 2). The gneissic rocks of the eastern area may be
part of an Early and Middle Archean core of the Wyoming
proyince.
The major rock type in the eastern Owl Creek Mountains
(Web, 1975) is layered quartz-feldspar gneiss. Amphibolite and
biotite gneiss are interlayered with the quartz-feldspar gneiss in
concordant layers centimetersto a few meters Ehick. The laye*
is apparently a t
r
a
mcornpxitiond layering, not bedding.
Two sets of fol& am be r e o o in~ the gneiss. F1 folds
are i s o c b l and reembent and, although they are refolded by
F2, they plunge gently east. F2 foMs are relatively open and
pludge shallowly east-southeast. F1 recumbent fold systems can
be recognized in other parts of the eastern Owl Creek Mountah,
indicating that a large area was affectedby this style of folding.
Recumbent folds were cut by diabase dikes prior to the last
folding. There is no evidence of metamorphism higher than amphibolite faeies during either of these deformations.
The two deformational events in the eastern Owl Creek
Mountains may correkte with the events recognized by Barker
(1976) in the Bighorn Mountains. Wells (1975) did not recognize
a post-tectonic mfic dike swarm in the Owl Creek Mountains,
but Houston has recently found that the post-tectonic dikes are
present, although uncommon. No geochronological studies have
*
I
I
f
I
133
Wyomingprovince
been undertaken in the eastern segment of the Owl Creek quartz-feldspar gneiss, fuchsitic quartzite, and metapelite. Part of
the succession in the Wind River Canyon has been interpreted as
Mountains.
The central segment of the Owl Creek Mountains consist of an isoclinally folded bimodal volcanic suite (Mueller and others,
an east-west-trending belt of metasedimentary and metavolcanic 1981). If the northern part of the central Owl Creek supracrustal
rocks, intruded on the north and south by post-tectonic granite sequence is isoclinally folded, the stratigraphy suggested by Gli(Fig. 3). The supracrustal rocks lie with structural discordanceon ozzi (1967) may be incorrect.
a basement consisting of interlayered biotite augen gneiss, granite
Dacite from the bimodal volcanic rocks of the Wind River
gneiss, amphibolite, and minor biotite schist, chlorite schist, and Canyon has been dated by the U-Pb zircon method at 2,905 25
biotite-muscovite schist (Bayley and others, 1973). The structural Ma. A Rb/Sr whole-rock isochron for dacite, basaltic andesite,
history of the basement is undetermined, so comparison with the and tholeiitic basalt indicates an age of 2,755 i 96 Ma, which is
gneisses in the eastern Owl Creek and Bighorn Mountains c a ~ o t presumably the time of metamorphism (Mueller and others,
be made.
1981). High-uranium granite of the post-tectonic granite complex
Foliation and rare bedding in metasedimentary and meta- has been dated at 2,730 35 Ma (U-Pb zircon) and 2,704 k 30
volcanic units strike east-west, parallel to the trend of the supra- Ma (Rb-Sr whole rock) by Stuckless and others (1985).
crustal belt, and dip steeply south. According to Gliozzi (1967)
.
and Bayley and others (1973), topping criteria suggest that the TETON AND GROS VENTRE RANGES
supracrustal succession is overturned. Hausel and others (1985)
J. C.Reed, Jr. and R. S. Houston
are uncertain of stratigraphic top and believe additional studies
are necessary to con6rm topping direction.
The Teton and Gros Ventre Ranges are segments of a
If the stratigraphy suggested by Gliozzi (1967) is correct, the northwest-trending Laramide uplift that was sundered in the late
basement gneiss is overlain by a succession composed of more Pliocene or Pleistocene time by north-south-trending normal
than 60 percent metasedimentary rocks, chiefly quartzite, iron faults. Precambrian rocks are exposed in the core of the Teton
formation, metapelite, para-amphibolite, and chert. In addition to Range (Fig. 4) and in smaller areas to the southeast in the Gros
the metasedimentary rocks, amphibolite and quartz-feldspar Ventre Range. Although the outcrop areas are limited, the local
gneiss of possible volcanic origin are present (Hausel and others, topographic relief in the Teton Range is more than 2 km,afford1985). The presence of chlorite-muscovite schist, cordierite- ing spectacular exposures (Fig. 5). The Precambrian rocks of the
muscovite schist, and andalusite-muscovite schist suggest lower Teton Range have been mapped and described by Reed and
amphiilite-facies metamorphism.
Zartman (1973) who also made a preliminary geochronological
The uppermost part of the supracrustalsequence is chiefly of study. Precambrian rocks in the Gros Ventre Range have been
para-amphibolite and orthoamphibolite, with minor interlayered mapped and described by Simons and others (1981).
*
*
I
EXPLANATION
Gronific rocks:
/eucogranife,
monzogronite,
and pegmatite,
.
Metamorphic rocks:
quartz biotite schist,
para -amphibo/ite,
Orfhoamphibo/ife,
quartiite and
iron- form0 tion.
i
a
Mine
4 mi.
1
I
SCALE
0
I
-4
'
Synform
5 KILOMETERS
d
Figure 3. Generalized geologic map of the central Owl Creek Mountains and Wind River Canyon (after
Hausel and others, 1985).
R S. Houston and Others
EXPLANATION
u
Quaternary and
Tertiary deposits
...
Mesozoic and
Paleozoic rocks
-
Diabase dikes
Mount Owen
Quartz Monzonife
Rendezvous Metagobbro
g
m
Augen Gneiss
Webb Canyon Gneiss
Layered Gneiss
49'
Strike and dip
of fo/iotion
Reverse fauN
(Teeth on upthrown side)
Nbrmo/ fau/t
(BUN on downthrown side)
-
0
SCALE
10 KILOMETERS
Figure 4. Geologic map of Precambrian rocks of the Teton Range (after Reed and Zartman, 1973).
The oldest rocks in these ranges are conspicuously layered workers), and a few pods of oxide-facies iron formation. No
and complexly deformed biotite gneiss, amphibole gneiss, am- quartzite or marble layers have been found, and pelitic rocks are
phibolite, and migmatite shown as layered gneiss in Figure 4. rare. No primary textures or structures have been recognized in
Pods and small tabular bodies of serpentinized dunite and perido- any of the gneisses, but the general range in compositions suggests
tite and associated uralitized gabbro are common along discon- that they were largely derived from mafic to intermediate voltinuous horizons in the gneiss sequence. In the Teton Range the canic and volcaniclastic rocks. Mineral assemblages are generally
layered gneiss sequence locally contains lenses and layers of those of the upper amphibolite facies, but Miller and others
quartz-plagioclase gneiss (mistaken for quartzite by some early (1986) report local assemblages with garnet, kyanite, and two
Wyomingprovince
I
'
R
Figure 5. Oblique aerial view of the mt face of Mount M o r 6 i c t h e T e t o n ~ e . T o ~ ~ C G h k f from foreground to summit is more than a kilometer. Country rock is predominantly migmatitic biotite
gneiss, light colored dikes, most of which dip gently north (right), are Mount Owen Quartz Monzonite
and related pegmatite. Prominent dark vertical band is a diabase dike 30 to 50 m thick. Light gray cap
on summit is Cambrian Flathead Quartzite.
ia a makd gay g & $ ~ ~ .
at -1 fmegmrdfeids
snrd Em\rwr par&bJ to l a y e tlw bgmt o b s a d have limb9 8
fsw !am~'s
bat hw.mu &re p C a w pram$. me'
h s l b d bl&, whBe adjamat byen are
..,
. -. -:
7
-~-&7.
L .
A
.
,
q.7
6s..
e 1
-.
' r y :
:-
..m
'-
R
136
+Y5
-.I
*
a
.
7
S. Hourton and 0
-
isoclinal limbs.Thus, while the layers probably reflect original
compositionaldifferences,their original sequence may have been
largely obliterated by layer-parallel ductile deformation.
Superimposed on the isoches are more open folds with
diverse axial orientations. Foliation in the layered g n e b is
generally parallel to layering, and mineral lineations are patallel
to the axes of the younger folds. This suggests that ampBi$olitegrade regional metamorphism was synchronous with formation
of the younger folds. Parall*
between fdliation and lineations
in the Webb Canyon Gn& and s i m k struclt.wesin the encIosing rocks indicate that the Webb Canyon was metamorphosed at
the same time as the layered &nehs. The RendezvousMetagtbbro does not display conspicuous heation, but crude foliation is
parallel to that in the nearby- gneiss,
and boudins and Mormed
layers i$ metagabbro in the gneh indicate emp1~~ement
prior to
or during deformation.A comwsite whole-rock Rb-Sr isoelvon
on sam~les&om the Webb
Gneiss, the Rendezvous
Metagabbro, and one sample of plagioclase gneiss yields an age
of 2,8 15 150 Ma (Reed and Zartmsuq 1973). They interpret
this as the approximate date of the amphibolite-grade regional
metamqhkm.
An irregular pluton of post-tectonic biotite momqranite is
widely exposed in the central past of the Teton Range. This rock,
the MOW Owen Qwrtz Momonite, locally displays faint flowfoliation, but no tatonle foliation. Dikes,pods, and irregular
bodies of p e m t e a few cedneters to several tens of meters
thick are ccunmon throughout much of the main part of the
ptuton; in many areasthey make up as much as one quarter of the
volume of the rock. Contacts of the Mount Owen Qmtz Monmnite are highly irregular and dBcult to depict on a map,
Blocky inclusions of wall rocks a few meters to tens of meters
moss are common throughout the pluton. As the margins are
approached, the inclusions become more and more abundant,
and there is a gradual tramition from momgranite with abundant inclusions into wall rocks mntainhg myriad c r m t t i n g
dikes of mowqranite and m t i t e . Reed and Zartmrtn (1973)
obtained aa Rb-Sr whuEe-rookisochron age of 2,440 75 Ma on
the Mokt Owen, but the age is suspect because of the anomalously high initial 8 7 ~ r ~ wratio
$ i of 0.732. Minetd isocbrons on
plagioclase and microcline separates from two samples of the
Mount Owen give agas of about 1$00 Ma, w M Reed and
Zartman (1973) suggest dates local strontium reqdiiration during a thermal event that affected many parts of the Wyoming
province.
Massive to faintly foliated biotite momgranitG is also exposed in the Gros Ventre Range. The rock contains abundant
inclusions of migmaW biotite gneiss, faliated graniticgneiss, and
layered biotite and hornblende gneiss,and locally it displays conspicuous megacrysts of potassium feldspar. The momgranite of
the Gros Ventre Range has not been dated, but it resembles the
Mount Owen Quartz Monzonite of the Teton Range and may be
equivalent to it.
Slightly metamorphosed but undeformed diabase dikes
crosscut all other Precambrian rocks in the Teton Range. The
*
., ' -
dikes are mar v;eEtd and trend approximately east to west. They
range in thi~k8essfrom less than a meter to mre than 30 xm&q
andmm6m betracedformorethan 1 0 h Reetaad(1973) did not obtain an unequivocal date for emplacement &
the dikes, but found Ehat the msimum K-Ar age on bi*
frm
wall m b nwr the dikes was a b u t 1,450 Ma. They suggest &t
this age is either a reflection of heating during e m p b e n t of &e
dikesdildildildildildildildildildildi
af d m t b g during a sttbequent thermal event. They
therefore s-t
that it is a k.iimum age of dike emplacetnena
W a r Rb-8-t and K-Ar ages of biotite reported from the Teton
hinge and adjacent areaa suggest a distinct thermal event during
the interval 1.3 to 1.5 Ga.
WIND RIVER RANGE
w.rl rudR
B.ILSton
Archean rocks in the northem half of the Wind River Range
consist largely of poorly dated g n e k and granitic rocks, those
in the smthern part are largely supracrustal (Plate2; Fig. 6).
h
hin the n-orthem half of the range (Agn in Fig. 6)
have yielded U-Pb zircon ages as old as 3,358 & 30 Ma and may
be part of the Early and Middle Archem wre of the Wyoming
craton. Late Arcbean felsic intrusives and orthogn&sm (Wg and
Wo in Fig. 6) divide the gmiss complex into m r a l segments
whose relative agm we unknown. Border phases of the orthogqeiss in the northvestern part of the range have U-Pb zircon
age of 2,699 i 7 and 2,671 i 5 Ma (J. N. Aleinlkoff, written
'comudm*
1984in Worl and others, 1986). The mjor rock
type in the gneiss complex is quartz-feldspar-biotite~~,which
contains layers and variously shaped masm of gmphibalite, biotite schist, pelitic gneiss, garnet gneiss, pyroxene gneiss, metadiabase, iron formation, and ultramatic rodr (Granger and others,
1971; Worl and others, 1984). Migmatitic gneiss in the northwestern part of the Wind River Rang (Fig. 7) records a comp1.e~
Wry, including at least two major events (Worl, 1968). Two
gener~t$onsof dgmatiw are recognized, an o b r of regiond
e m t , and a ysugger rc10td to a Late Archean granitejust to the
east. Struehwd d p b Ma@ aa early stage of mumbent
folding and d e v e l m t of ax& p h e trampmition foliation,
followed by fahhg sf the FoWon, fhs around m 1 y horizontal
axes and then around nearly vertieal axes. Diabase.dikes were
intruded
the fdiation prior to the last folding. In one area
the ma& part of the older migmatite is largely iron formation
that shows evidence of two episodes of metamorphism, an early
granulite-Eaciesevent and a late amphibolite-fuiesevent (Worl,
1968):
Two of the major plutans in the central part of the raage are
the Louis Lake Batholith and the Bears Ears plubn. The farmer
consists diefly d granodioriEe, but includes quartz diorite and
quartz
the latter is chiefly @te. The age of the
Louis L a b baolith is mtimatd as 2,630 20 Ma on the basis
of U-Pb zkcoa, R W whole-rock, and other methods. The best
estimate of the age of the Bears Ears pluton is 2,545 i 30 Ma by
the same methods (Stuckless and others, 1985).
*
Wyomingprovince
--
In and Others
other rmprammd rocks, but is probably the youngest. It is mm. posed of metagraywadre, meatuff, metacmaglomerate, and man$&&. The graywwke was probably deposited by turbi&~
currents in an m&bk basin or trough.
The s u p r m a rocks of the South P w area are intruded
by a vari&y of igomus rocks hdudhg,from oldest to
pexidotite, digbase dikes and sills, diorite dikes.and sills, ha&cite and ton&b b a n d s d Stocks,
dike,
ite of the L d s M e Batholith, and dhbase dikes. All igneam
units oMer t b n the Louis Lake Batholith were affected by re
g i d m-ism.
The m p ~ succmsion
d
at South Pass has been sub
jectwI to at l m t two episodes of deformation. The earliest prod& a northmt-strikhg synclinorium with both upright and
recumbent folds. Some mafic intrusive bodies were folded during
this event, and others cut obliquely across the folds. The second
deformation produced a series of antifoms and synfonns that are
best develop& in the western part of the South Pass area. The
second ddomtion is interpreted a9 syndx~nouswith e m p b
e part of the South
m a t of the Lark Lake Batholith. In the w
Pass area and in Bhe area to the west, the sup~acrustalrocks are
and presumably have a
more highly cbhmed and
history. Worl(1%9) md Hedge (1963)
more mnz~~l~structwra1
bosh r&
the two epJsoda of folding suggest& by B-ayley
and dbms (1973), but belieye that the western area was de
f o d a third time to develop faults and similar folds related to
the development of a steeply dipping n o r t h w a t foliation.
~
Worl and others (1984) have demonstratsd that supracrusta1 rocks extend for 16 km or more northwest of the South Pass
area along the wmtem margin of the range. Th;ms u p m E a l
rocks (shown as Agn in Fig. 6) are of higher metamorphic grade
than those of the South Pass ar* are extensively migmatimd,
and are sepaated from the South Pass succession bv later intmsive m&s.-~heirrelationship to the South pass suPr&wtal m k s
istlnkmwn.
Figure 7. Migmtitic rocks in the Wind Rivw &age. A. Mipatitic
Thew of the supr-tal
r& in the South Pass area has
gneiss Wt $ray) amtaking boudins of &Bob
(arkgray) cut by
been
e&abLLed.
hterman
(1.982)
d%&& a RblSr whole
not
granite dikes pmbab1y related to the Bears E m plubn. B. Passive flow
gneiss. Much of 6 6 @&sic rwk in the mge rock imcbcm metagaywacke and hterloyerd metavolcanic
folds in layered *titic
is similar to this in that it has an i p p ~ o u s - a p( ~g e m h ~ d y racks that a
rn age af 2,800 i 100 Ma. This date is
mobile) leuamme and a metamorphic-appearing @mhe&dy im- imprecise b u m of the lirniteU range of Rb/Sr ratios, but within
mobile) melasome.
the large d t y , it probably reflects the date of amphibolitegrade metamwphism.
basal unit is overlain by a thin sblf sequence, the Goldman
Meadows Formation, which is in turn overlain by 3,000 m of GUNlTE, FERRIS, AND SEMINOE MOUNTAINS
volcanogenic rocks of the Roundtop Mountain Chenstone and
Miners Delight Formation. The Goldman Meadows Formation
Precambrian rocks of the Granite Mountains uplift are exconsists of quartzite, iron formation, and pelitk &kt. The schist
is interlayered with iron formation and has gradational contacts posed as exhumed peaks surrounded by Middle and Late Cenowith q d t e . The schist is chiefly quarti-raia-~ndalusitescW zoic sedimentary rocks and constitute the Granite, Ferris, and
but some par@are rich in chlorite, garnet, amphibole, and biotite, %minoe MouatainS (Rate 2).
The northwest Granite Mountains am&%of g n e k that
indicating lower amphibolite-grade metamorphism. The Roundplutom of &te
and deformed
top Mountain Greenstone is a metavolanic unit that consists of are invaded by --tectonic
greenstone, greenschist, vei& and pillowed metabdt, and cliabase clilms. The gneisses are intahymed with metasedhenmetatuff. The Miners Delight Formation is in fault eontact with tary rocks and are so deformed that age relationshipsare obscure.
-
~~
$:=*-
139
Wyomingprovince
Gradotionaf or disconfomobIe
contact
Iron-format(on
schist, quartzfts
(Gofdman Meodows Formation)
-
/
SCALE
Contact
0
/LO;-
k
+I
Gabbro dikes
Fault
Sherer (1969) considered a biotite gneiss to be the oldest unit, and
Peterman and Hildreth (1978) include the biotite gneiss of Sherer
in their oldest rock unit, which consists of tonalitic to granitic
biotite gneiss containing subordinate interlayers of mica schist,
amphibolite, epidote gneiss, and augen gneiss. Although both
Sherer (1969) and Peterman and Hildreth (1978) establish tentative stratigraphic columns with supracrustal rocks apparently
younger than the gneiss, they indicate that age relationships are
uncertain.
Peterman and Hildreth (1978) obtained a whole-rock Rb-Sr
age of 2,860 k 80 Ma for the biotite gneiss. They interpret this as
the time of metamorphism. The isochron plot can be interpreted
to suggest two rock suites: a layered, metavolcanic, metasdimentary rock suite with an initial ratio of 0.7048 0.0012 and a suite
of amphibolite andgranite gneiss with an initial ratio of 0.7017.
They suggest the massive gneisses may have had a different origin
than the layered gneisses and that original differences in isotopic
composition may have been partially preserved during metamorvalue of 0.7048
phism. The layered gneiss has an initial 87Sr/86~r
0.0012. This high ratio suggests that the protoliths of the
layered gneiss had a significant crustal history prior to metamorphism. Peterman and Hildreth (1978) consider two possibilities:
(1) derivation from a 3,200 to 3,300 Ma volcanic pile and/or
sedimentary sequence with mantle-like 8 7 ~ r / 8 6ratios
~ r that
*
*
10 KILOMETERS
evolved to the 0 b ~ e ~ value
e d before the 2,860 Ma metamorphic
event, or (2) derivation from a sedimentary pile deposited shortly
before 2,860 Ma, but derived from older materials.
In the western Granite Mountains, Peterman and Hildreth
(1978) d e s c n i a hegrained epidote gneiss that may have
been derived from silicic volcanic rocks, a medium-grained amphibolite that may have been derived from mafic volcanic rocks,
and a fine-grained layered biotite gneiss that they believe may
have been graywacke. In the Barlow Gap area of the Granite
Mountains, Houston (1973) and Bickford (1977) mapped a body
of supracrustal rocks surrounded by younger granite. These supracrustal rocks consist of interlayered amphibolite, hornblende
gneiss, felsic gneiss (probably graywacke), calc-schist, quartzite,
and oxide- and silicate-facies iron formation. Carey (1959) and
Pekarek (1974) mapped a large area of hornblende schist and
mica schist with local beds of iron formation north of Barlow
Gap. Not enough information is available to estimate the relative
proportions of rocks of metavolcanic and metasedimentary rocks
in these units.
Foliation in gneisses and supracrustal rocks strikes northeast
and dips south in the western Granite Mountains and changes to
northwest with a south dip in the central Granite Mountains. This
arcuate form is like that of metavolcanic rocks of the Seminoe
Mountains to' the south.
140
R S Houston and Others
The Ferris Mountains form the southem margin of the
Granite Mountains uplift. Precambrian rocks are exposed in the
cores of two elongate faulted anticlines. They are known also
from oil test holes south and southwest of the Ferris Mountains.
Precambrian rocks of the Ferris Mountains are similar in composition and age to those of the Granite Mountains, and consist
dominantly of granodiorite and granite (about 2,550 Ma) that
contain inclusions of metavolcanic and probable metasedimentary rock (about 2,850 Ma). Basalt and diabase dikes are abundant. Rocks in the subsurface to the south are mainly
metasedimentary.
The granodiorite is a medium-grained locally porphyritic
hornblende-biotite granodiorite containing phenocrysts of alkali
feldspar 1 to 4 cm long. The fresh rock is medium light gray to
pinkish light gray, but is locally pale red with pervasive hematitic
alteration. A weak to locally strongly developed gneissic fabric
generally trends east-northeast or west-northwest and dips 70° or
more north. Elongate inclusions rich in hornblende and biotite
are locally common. Fine to coarsegrained muscovite pegmatite~cut the granodiorite; aplite is present locally. The granodiorite
contains xenoliths of quartz diorite, tonalite, and felsic gneiss as
much as a kilometer in diameter. Bodies of amphl'bolite containing highly contorted pink quartzo-feldspathicpegmatite are present along the southern margin of the granodiorite body. Such
amphl'bolite lenses may originally have been basalt. Lenses of
epidote- and hematiterich quartz rock, quartzite, and epidote
chlorite schist are also caught in the granodiorite. These inclusions range from 0.5 m to 60 m thick and are as much as 300 m
long. Contacts of the granodiorite with porphyritic granite are
gradational or interfingering over a few meters to tens of meters;
the granodiorite seems to predate the porphyritic granite, but
perhaps by only a short time.
Abundant ma!% dikes cut the granitoid rocks. One suite
consists of aphanitic greenstone composed primarily of chlorite,
hornblende, lesser amounts of plagioclase, and minor pyroxene.
These dikes range in thickness from a feather edge to about 8 m,
but are generally less than 5 m thick. Dikes trend about N20° to
N30°W in much of the northeastern blocks, north to N15OE in
the central part of the area, and north to N30°W in the southwestern block.
The Seminoe Mountains (Plate 2) have been studied by
Bayley (1968), Dixon (1982), Klein (1982), and several earlier
workers. A sequence of metavolcanic and metasedimentary rocks
resembling a classic greenstone belt is exposed in the Bradley
Peak area. The sequence is intruded by granodiorite plutons, but
no basement is exposed. The su~racrustalsequence is at least
5,000 m thick; volcanic rocks make up 3,500 m of the sequence.
The lower 3,000 m is high-Mg thoIeiite interpreted as he
grained volcaniclastics and massive flows. The tholeiitic volcanic
rocks contain thin layers of siliceous and aluminous metasedimentary rocks in their upper part. Klein (1982) has recognized
komatiitic ultramafic rocks in flow units separated by thin layers
of iron formation. The upper part of the succession includes thick
beds of oxidefacies iron formation, fine-grained siliceous and
aluminous metasedimentary rocks, and minor calealkaline metavolcanic rocks. The greenstone belt succession is deformed and
metamorphosed to amphibolite facies, intruded by tholeiitic gabbro and granodiorite, and cut by dikes and sills of diabasic
gabbro.
S. S. Goldich (personal communication, 1985) reported a
U-Pb zircon age of 2,718 18 Ma for rhyodacite interbedded
with the metasedimentary rocks. Klein (1982) suggests that the
age of the granodiorite pluton, which intrudes the Seminoe succession, is approximately 2,750 Ma.
*
MEDICINE BOW MOUNTAINS
R. S. Houston and K.E. Karktrorn
A unit of quartzo-feldspathic gneiss in thesMedicine Bow
Mountains is interpreted by Houston and others (1968) as base
ment to supracrustal rocks of Late Archean and Early Proterozoic
age (Plate 2, Fig.9). In some parts of the area the gneiss complex
is either in fault contact with supracrustal rocks or is separated
from them by intrusions, so that the relative age of the two
successions cannot be established.
The quartzo-feldspathic gneiss complex cobists of layered
gray biotite gneiss with subordinate interlayers of hornblende
gneiss and amphibolite, and rare layers of quartzite, marble, ultramafic rocks, paraconglomerate, and pelitic schist. Weakly
layered pink quartz-feldspar gneiss locally cross-cuts layering in
the biotite gneiss, but is deformed with it and is generally concordant. The entire gneiss complex is intruded by at least two sets
of diabase dikes, one that is folded with the gneiss and another
that cuts the gneissic foliation nearly perpendicular to strike. If
some amphibolite layers in the biotite gneiss are intrusive, as field
relations suggest, there may have been three sets of mafic dikes.
Inasmuch as most of the dikes and sills that cut supracrustal rocks
are nonfoliated, it is possible that the "lmwment" gneiss succession has a longer structural history than the supracrustal rocks.
This provides indirect evidence that the gneisses are older than
the supracrustal rocks.
Hilb and others (1968) dated samples of the biotite gneiss of
the western Medicine Bow Mountains by the Rb-Sr wholerock
method at 2,500 50 Ma, a date that they interpreted as a
metamorphic age. This confirms an Archean age for the
protoliths.
The Phantom Lake Metamorphic Suite of the northern
Medicine Bow Mountains (Fig. 9) is a Late Archean succession
containing about 60 percent volcanogenic rocks and 40 percent
sedimentary rocks. The rocks are complexly deformed and are
metamorphosed to amphibolite facies. Outcrop patterns are
characterized by abrupt lithologic changes, either facie5 changes
or tectonic slivers-hence use of the term "Metamorphic Suite"
rather than "Group" (North American Commission on Stratigraphic Nomenclature, 1983). The Phantom Lake is divided into
five lithodemic units (Fig. 10).
The lower contact of the Phantom Lake Metamorphic Suite
*
~
Wyoming province
1 0 6 O 30'
106'
I
I
MEDICINE BOW MOUNTAINS
EXPLANATION
ROCKS SOUTH OF
CHEYENNE BELT
NORTH OF CHEYENNE BELT
- . # m O C K S
1 ,:F-
L
,& .i
1-L
C
Granite
Sherman Granite
/-I400 Ma)
Mafic Sills and Dikes
SIERRA MADRE
/-
MED~CINEBOW
MOUNTAINS
+
Gronodiorite
Granite
1 7 0 0 Ma)
..=--
-,,-Shear Zone
-+.a
4
Thrust FouN
v
and dikes
Vo/conogenic Gneiss
l- 1800 MOI
\
/\"&
Phantom. Lake
Metomorph~eSurh
Anticl(ne
Milcon Mountain
Melovolconics
ond
Overfond Creek
. Gneiss
--*k---
Syncli~e
'
T o
Attitude of Bedding
-- ---- Gneiss
Figure 9. Generalized geologic map of the Precambrian rocks of the Sierra Madre and Medicine Bow
Mountains (after Karlstrom and others, 1983).
R S.H o w n and Othem
142
MEDICINE BOW
MOUNTAINS
t
ic.
I.
.'-
Fiwe 10, S&awPhYof Archean and koerozoic
me-enw
rocks in the Sierra Madre and Medicine Bow Mountains.
-
.
.
, ,
.
.
-,-
7
.
v.
is poorly ex@, h the northern Medicine Bow Mountains, the
Phantom L&@ ~ t a m o r p h i Suite.structually
c
overlia the h r land Creek Gneiss, whicb coQsists of hornblende and bi&
gneiss, prcrbably of v o ~ o@h.
~ c h t h units axe crosscut by
granitic &k and dikes of dinown age &at rnay welate with
the. 2,425 hh Baggot Rocks Granite of the westem Medicine
Bow Mountains (Hills and others, 196% Remo, 1383) or the
2,665 to 2,683 Ma orthogneiss of the Sierra Madre (Prmo9
1983). The s
~ ~of the
c Overland
e
Creek Gneiss is obscwe.
It may be a remnant of an extensive greenstone surx;essioa, or it
may represent volcanism in the earlimt stages of deposition of the
Phantom Lake Metamorphic
The q ~ - E e l d s p a ~gneiss
c in the wester11 Medicine
Bow Moun&ins is nowhere in contact with the Phantom Lake
M e t a r n e e Suite. m e fobwing summary of the stratigraphy
of d e W:~MWBM e Suite of the Medicine Bow Mountains (Fig.
10) is from Ktulstrom and others (1981).
The oldest unit in the Phantom Lake Suite, the Stud Creek
Vol&l~stiics, contains amphile schist and pelitic schist believed to represent a period of dominantly subaerial voleanism.
The d t also cothin, slightly radioactive quartzites and
oot@jlo-t;es
in its upper part, which may represent W fluvd
Ifepodtian ~ o ~ t aad
o nKarlstrom, 1979).
The over1ybg Rock Mountain (hglomwate consists of
arkosic polymictic paracosglomaate, e , a b i h g stretched c k t s
of quartz, quamite, rrndschhtose rock, iaterbedded with p u l a r
qmrtdte. Thew mi& are interprebxl to have been &posited in
alluvial ftbns and &ated
braided streams because they contain
~ t i n u o u zones
s
of poorly sorted, matrix-supported, arkosic
paraco@omera@ which are interpreted to represent proximal
debris flow wmchted with the alluvial fans and sveam deposits.
The paswnglomrate is slightly radioactive and contains interbedded, mawe, quartz-pbbie conglomerates that are l d y
moderately ~ ~ vRadioactive
e . mine& indude monazite,
thoriadie, and zirmn, which probably represent heavy minerals
concentrated by EZuvialprmwes,
The RBck MQ*
Conglomerate grades u p m d into the
Bow Qawtzk, W h is compased of fiine+yained arkosic and
subarkoak quartzite. U&m hegrain h,
l a r p s d e planar
ao&eds, osdktiod ripple inarb, and a bimdal-bipolar pal m m t distribufion mggmt deposition in a shallow maiine
environ-t.
The Mckest with the P B B LLake
~ ~MehmorphicSuite is
the Colberg MeEavolc8&s, a s q m c e of metabasalt, m&c tuff,
and wlCani&& ro& A b p-t
is a pamanglomerate containing cobblesand M 6 t s 6f granite, q d 4 and basalt in an
a m p h i b l ~ u m t zmatrix. The pataconglomerate contains 1 4
interbedsd qmrtz-pebble conglomeratethat are slightly radioacr
tive. The W i t i o n a l history of the CoIberg Metavolcanics is not
well undersmxt The unit is interpreted to be partly marine because of the presence of piIlow basalt and overlying and underlying m a d e quartzites. The paracoqlomerates may have been
deposited in akivial or submarine channels and fans adjacent to
Mt scarps bounding volcanic h@dands. However, the low pro-
.
143
Wyoming province
r
portion of pillow basalt, the wide variety of lithologies, and the
rapid facies changes in the unit seem more compatible with subaerial volcanism and volcaniclastic deposition. If so, the Colberg
Metavolcanics may contain complexly interbedded subaerial and
submarine rocks.
The Conical Peak Quartzite is the youngest unit in the Phantom Lake Metamorphic Suite. It contains white, fine-grained,
micaceous subarkose similar to the Bow Quartzite except that the
Conical Peak Quartzite is richer in plagioclase. It seems likely
that it was derived from reworking of the Bow Quartzite with the
addition of detrital plagioclase from the Colberg Metavolcanics.
The unit is considered marine on the basis of fine grain size,
large-scale planar crossbeds, and a bimodal-bipolar paleocurrent
distribution.
Rocks of the Phantom Lake Metamorphic Suite are folded
Fnto tight to isoclinal folds with axial traces ranging from eastwest through northeast to north-south. These folds were formed
prior to deposition of the Early Proterozoic Deep Lake Group, as
shown by local angular unconformity between the two successions and by major differences in structural style and degree of
deformation. Folding of the Phantom Lake Metamorphic Suite
-. may have taken place during the 2,500 Ma tectonic and thermal
--. episodes that coincided with intrusion of the Baggot Rocks and
related granites of the Medicine Bow Mountains and Sierra
Madre.
The age of the Phantom Lake Metamorphic Suite is best
established in the Sierra Madre and will be discussed below.
/'
sIEm w
m
R. S. Houston and K.E. Karlstrom
The quartzo-feldspathic gneiss in the Sierra Madre (Houston
and Ebbett, 1977) is probably a continuation of the similar gneiss
in the western Medicine Bow Mountains, although outcrop is
interrupted by Tertiary cover (Fig. 9). In the Sierra Madre the
gneiss is generally in fault contact with supracrustal rocks of Late
Archean to Early Proterozoic age, but locally the contact with
Late Archean supracrustal rocks is marked by a quartz-pebble
conglomerate, the Deep Gulch Conglomerate. The presence of
the conglomerate suggests that the contact is an unconformity
(Karlstrom and others, 1983). A quartz-mica schist, which occurs
locally along the contact, may represent a metamorphosed
regolith.
Several attempts have been made to date the gneiss in the
Sierra Madre (Hedge in Karlstrom and others, 1981). Early RbSr whole-rock ages were Archean, but with large uncertainties
(Hills and Houston, 1979). Premo (1983) dated two intrusives
using the U-Pb zircon method. One of these, the Spring Lake
Granodiorite, intrudes both "basement" gneiss and the lower part
of the Phantom Lake Metamorphic Suite in the north-central
I Sierra Madre (Karlstrom and others, 1981). The granodiorite
contains a foliation that is parallel to foliation in its wall rocks.
The age of 2,710 rt 12 Ma for the granodiorite is a minimum for
the enclosing gneiss and the lower Phantom Lake Metamorphic
i
Suite. A pink quartz-feldspar orthogneiss at two localities in the
northern Sierra Madre yielded ages of 2,665 28 Ma and 2,683
5 Ma, but its field relations are uncertain.
There are two Archean supracrustal successions in the Sierra
Madre: the Vulcan Mountain Metavolcanics, and the Phantom
Lake Metamorphic Suite. The Vulcan Mountain Metavolcanics
comprise a succession of severely deformed metavolcanic and
metasedimentary rocks that apparently underlies the Phantom
Lake Metamorphic Suite (Fig. 9); however, the relationship is
uncertain. The Vulcan Mountain Metavolcanics (Fig. 11) include
pillow lava, tuff, quartzite, marble, conglomerate, and pelitic
schist that may correlate with the Overland Creek Gneiss of the
Medicine Bow Mountains (Fig. 10) or that may simply be a lower
part of the Phantom Lake Metamorphic Suite.
The Phantom Lake Metamorphic Suite in the Sierra Madre
contains the same lithologies as in the Medicine Bow Mountains
(Fig. lo), but the succession is even less certain because deformation has destroyed most topping criteria.
The oldest unit in the Phantom Lake Metamorphic Suite in
the Sierra Madre is the Jack Creek Quartzite, which includes the
basal Deep Gulch Conglomerate Member in the northwest. The
Deep Gulch Conglomerate Member lies unconformably on
basement gneiss, and contains arkosic quartzite with beds of radioactive quartz-pebble conglomerate. It is interpreted as fluvial
(Karlstrom and others, 1983). The Deep Gulch is overlain by
arkosic and micaceous quartzite with phyllite lenses, carbonate,
and phyllite. The upper part of the Jack Creek Quartzite consists
of micaceous 'Suartzite, arkosic quartzite, and paraconglomerate.
Planar crossbedding is common in the quartzite, and wedgeshaped and herringbone cross-stratification has been identified. In
outcrop of the central and eastern Sierra Madre, the Deep Gulch
Conglomerate Member is missing and the Jack Creek consists of
fine-grained sericitic and calcareous quartzite. Except for the
Deep Gulch Conglomerate Member, the Jack Creek is interpreted as marine.
The Jack Creek Quartzite is overlain by the Silver Lake
Metavolcanics. Although mafic volcanic rocks (probably basalt
flows) are present, they are much less abundant than in the Medicine Bow Mountains. The most abundant rocks in the Silver Lake
are metatuff, metagraywacke, biotite schist (probably tuff), paraconglomerate, quartzite, and carbonate. The rocks are generally
finer grained than the Colberg Metavolcanics, and include fewer
flows, suggesting deposition farther from the source.
The overlying Bridger Peak Quartzite consists chiefly of
fine-grained quartzite and phyllite, but its depositional environment is unknown.
The metasedimentary rocks of both the Vulcan Mountain
Metavolcanics and Phantom Lake Metamorphic Suite are involved in overturned folds or nappes that have east-striking and
shallowly south-dipping axial planes (Fig.
9). Geometry of minor
folds suggests that early recumbent folds were subsequently refolded. Later deformation produced upright macroscopic folds
with west-plunging fold axes. These folds are best developed in
the western Sierra Madre.
*
*
R S. H ~ m & and
n Others
1 G.L
WUMH MOlJNTA@JS AND CASPER M O m m
II
Snyder
The Laramie Moyntah are a range af low mountains and
hills carved from a block of Precambrian rocks u@&d in Lste
,
B
Figure 11. Primary structures in Late Archean rocks in the central Sierra
Madre. A. Pillows in the Vulcan Mountain Metavolcanics. B.Paraconglomerate in the Silver Lake Metavolcanies.
The gneissic basement, Vuloan Mountain Metavo1mics,
and the Jack Creek Quartzite of the Phantom M e Metamorphic
Suite are all intruded by the 2,710 Ah Spring Lake Granodiarite.
Because the g r a n h r i t e is deformed aloag with &e rnetasedimentary rocks, Karlstrom and others (1981) mcluded that it
was intruded before or dduring development of the early foldnappe structure. However, only the lower part of the Phmtom
Lake Metamorphic Suite is cut by dated intrusions, so part of the
succession may be Proterozoic.
Cretaceow a d Early Tertiary time. The rotnge extends northward for about 100 km from the Colorado-Wyoming brcder
southeast d Laramie and then trends &rthwest for another 90
km (mate 2). P r c c a n b rocks are aLso exwsed on Casuer
~ o ~ & t a iasmall
n,
uplifted block about 15 &northwest of ihe
northwestern end of the main Laramie Mountains, and in erosional windows through Phanerozoic rocks in various places on
the east side of the range.
The Precambrian rocks in the southern half of the range are
chiefly phtonic rocks of Middle Proterozoic age,.including the
Sherman Granite and anorthosite, syenite, and related rocks of
the Latamie Anorthdte Complex. The Sherman Wte and the
L m d e Anorthde Complex are collectively referred to as the
btrvsive suite d @Mile Creek by Snyder (1984). They are pw
of the Middle Proteromic anorogenic suite &cussed in Chapter
4 (this volume). These Proterozoic plutonic rocks po&date
movement along the Cheyenne belt. and ocoupy its projected
trace n d m t of the Medicine Bow Mountains.
The Arch- rocks, which are widely exposed in the northern half of the k a m i e Mountains and on Casper Mountain,
&dudegranitic mks, *titic
g n b , and mveral sequence
of s u p m M rocks, all of which are cut by ma& dikes of Late
Arcbesn w d Proterozoic q e (Love and Christiansen, 1985, and
r d r e n m &&Atherein). Much of the kchean terrane in the
central and northwetem part of the h a m i e Mountains consists
of &te of the h m i e batholith of %die (1969), dated at 2.5
to.2.6 Ga by the R M r whole-rock method'(Hills and Amstrong,
197% P&mm, 1982; and 2. E. Peteman, Pnitfp;'tlcombtion, 1983). The batholith eontaias medium- to coarse-grained
grade t b t is massive to gneissic, and generally rich in biotite,
The grad& passets both southward and northward into d@mtisic gnthe "met83norphic wmpkP of Condie (1969). The
migmatite tenme south bt&e batholith was W b e d by Snyder
(1984) a the @cow a d rnetamorpbic complex of Laramie
River. It canaists of interleaved gmite+ granite gneiss, augen
s&, and
ite, with deform& sills, dikes, and l a s s of
mafic and &m&c rocks. Many of the graniticrocks are a p ently apophpa ~ f or, at legst synchronaus with, the Laramie
batholith. Somesmallbadies of graniti~to graasodioritic racks are
signkm1y yomger and tnrnsed the older granites. One of these
yielded zifcon with s U-Pbage of 1.7 Ga (K .R. Ludwig, oral
canmmication, 1984). Most of the granitic rocks are componats of migmatim that represent the erugt into which the h&usive parts of the complex were emplaced. Along the east side of
the central Laramie Mountaim, some of the rocks have an RbSr
model age 5f 3.2 Ga, but most fit on a 2.6 Ga wholetock
isochron. The lapred migmatites may represent arkoses or rhyolitic extrusives(Smithson and Hodge, 1969) that were deposited
1
Wyomingprovince
I
I
I
145
in the Middle Archean, and remobilized and recrystallized in the
Supracrustal rocks are exposed in the central Laramie
Late Archean. Graff and others (1982) report that some of the Mountains, near the northwestern end of the uplift, and on
conglomerates in the supracrustal sequence of the central Laramie Casper Mountain (Fig. 12). Those in the central part of the range
Mountains contain clasts that can be identified as migmatitic have been described as the Elmers Rock greenstone belt (Gsaff
rocks from this terrane. The main batholith contains roof pen- and others, 1982) and the metamorphic suite of Bluegrass Creek
dants of hornblende gneiss and schist, and interlayered quartzite, (Snyder, 1984). They occupy synforms or occur in homoclinal
mica schist, and iron formation. These and lenses of similar rocks sequences between domes of granitic gneiss interpreted as remoin the migmatitic terrane may either represent parts of the early bilked basement. The sequence consists of layered mafic metabasement or parts of the later supracrustal sequences.
volcanic rocks, chiefly massive to layered amphl'bolite, and a
SCALE
-
-
COVERING DEPOSITS
I
w
C
P
ANORTHOSITE
a LAUCOQABBRO
SYENITE.
MONZOSYA-NITE,
6 MONZOGRANITE
SHERMAN
i
-
-
)94*
--COL
-e
WYO
0.
Z L
INDEX MAP
WITH AS
MUCH AS
BASE DIKES
<
W
'2
META
SEDIMENTARY
ROCKS
:SHALE.QRAYWACKE. CONGLOldERATE
6 DOLOMITE)
VOLCANIC
ROCKS
(CHIEFLY
J
$I L L
0
Figure 12. Generalized geologic map of the central part of the Laramie Mountains (after Snyder, 1984).
I
R S Houston mi Others
-146
variety of interlayered metasdhentary rocks, all of which are
~ n m r p W
to amplrrhlite grade. Those with appropriate
c~mpositionscommonly contain maisthg d h m i t e , kyapite,
and mddwite. The preservation of primary features such i
s
crossbeds and graded beds in sedimentif, pillows in hva, and
emulate hyming in mafic sills (LaqstafT, 1984) makes is possible to establish the stratigraphic sequence with some confidence
in some areas. Most of the mebvolcanics are tholeiitic and they
locally con& welldeveloped mygdular pillows near the base
of the -ion. Layered amphibolites with random splays of amphibole on fbliation surfhave compositions of tboleiitic andesite or calc-dkalic basalt and ?re co~ldmonin the upper paa.
Metasedimentary rocks intercalated with the volGanics include
graywacke, marble,quartzite, iron kimation, a variety ofmi@,'
and conglommtes containing abmdmt m b b h and bosf
granitic r o c k Metapelites, metagrqw~ches,and tremolitic or chondroditic dolomites in the upper part of the seation mag Eie .
unoonformably above the volcanic sequence as q a W in Fig- ,
ure 13,The rocks are all interpreted to hare been &pasired ia a
shallow marine environment.
& o h m supracrustal mks in the. northwestern La&
Mountah md in Cmper Mountain have been descni by
Jbbnson aad Ms [1976), Gable (1987), and GaMe and others
(1988). Thew rac&s hclude nzetqraywacke, meWmdstone,
mewmt-ab,
and m%Wt.They are not i~
weU exposed
as those the south, theh st&@&w&seauenee is unagt&, and
.;* *-wi
/$ *.'
L~8.*-4,!&:
,
WEST
EAST
CENTRAL LARAMI
RANGE
41" 55' N, 1 0 5 ~ 2 0w'
*
HARTVI LLE
UPLIFT
25401
42°18' N, 1 0 4 ~ 4 2 ' ~
E
n
EXPLANATION
Carbonate
Shale.
Sandstone
graywacke
with conglomerate
layers
Felsic
volcanic
Basalt.
local
pillows
Ironformation
Peridotite
Granite
with mafic
intrusives
Unconlormity
Figure 13. Correlation of Archean metamorphic rocks in southeastern Wyoming.
Ages
in Ma
\
i
I
:.I
!
147
Wyoming province
their age is undetermined. Detritus in the clastic rocks suggests
that they were derived from source areas that included granite
and other crystalline rocks. These supracrustal rocks apparently
predate an episode of regional dynamothermal metamorphism at
about 2.8 Ga, and they may correlate with rocks in the Granite
Mountains for which a minimum age of 3.2 Ga has been established (US. Geological Survey, 1979). They are invaded by large
mafic to ultramafic plutons that were emplaced during the waning phases of regional metamorphism.
The Precambrian rocks of the central Laramie Mountains
may include as many as four depositionalsequences separated by
three plutonic episodes. The first two plutonic events at 2.6 and
1.7 Ga were probably synchronous with folding and regional
dynamothermal metamorphic events that produced the widespread aluminum silicate triple-point assemblages. The last was
emplacement of the Middle Proterozoic anorogenic plutons at
about 1.4 Ga, which produced purely thermal metamorphism in
a kilometer-wide contact zone. Multiple episodes of deformation
have been recognized, but the details of the early tectonic history
are poorly understood.
The voluminous mafic and ultramafic dikes in the Laramie
Mountains range from early porphyritic metabasalt through nonporphyritic tholeiitic diabase or basaltic komatiite to variously
altered peridotite and serpentinite. The dikes trend generally
northeast and have Sr isotope ratios and REE patterns that
suggest at least two distinct mantle sources (Snyder and others,
1985). Emplacement of the dikes predated at least one period of
metamorphism and deformation, because most diabase dikes
contain metamorphic garnet and some folded amphibolite dikes
display axial plane cleavage. However, it is not yet clear whether
dike emplacement was episodic or continuous and whether it
occurred mainly in the Archean or mainly in the Proterozoic.
HARTVIUE UPLIFT
$cAssA
ANTICLINE
ANTELOPE H I L L S
I
G.L. Snyder
SCALE
The Precambrian rocks of the Hartville Uplift consist of
medium- to high-grade metasedimentary rocks cut by three ages
of felsic plutonic rocks and two ages of m&c dike rocks (Fig.
14).
The earliest study of these rocks was by Smith (1903). Many
subsequent studies are summarized by Snyder (1980), Peterman
(1982), and Snyder and Peterman (1982). The earlier work recognized the major lithologies and some broad age groupings of
rocks; the later work developed a detailed stratigraphy and established the sequence of geologic events (Fig. 15). The Precambrian
rocks of the Hartville Uplift contain small base metal and uranium deposits, and large reserves of hematitic iron ore and
crushed rock (Harrer, 1966; Hausel, 1982).
The Whalen Group of Smith (1903) consists of 70 to 80
percent clastic or biochemical sediments and 20 to 30 percent
rnafic extrusives or volcanic-derived sediments. The group is divided into four formations that together compose more than
3,700 m of section (Fig. 15). These rocks, which contain rare
EXPLANATION
10
IS KILOMETERS
Granite of
Rowhid8 Buttes
Contact
Formofion of
Muskrat Canyon
Figure 14. Generalized geologic map of the H a M e Uplift (after
Snyder, 1980).
Figure 15. Diagram of facies changes in the W e n Group in the Hartville Uplift.
granitic pebbles, were probably deposited on a basement of 3.0
Ga sialic rocks (not exposed) that was later reactivated to intrude
them (Peterman, 1982; Snyder and Peterman, 1982). The sequence contains well-preserved primary structures, indicating
stratigraphic tops; geologic mapping has demonstrated open to
isoclinal north-northeast-trending folds tightly refolded along
east-west axes.
The lowest unit of the Whalen Group, the formation of
Muskrat Canyon, is predominantly carbonate rock that includes
nearly pure tremolite dolomite in the southern part of the uplift
and siliceous or tremolite dolomite with interlayers of schist,
graywacke, or quartzite in the central part. Thickness ranges from
900 m to more than 1,500 m; the base of the formation is not
exposed. The next formation, the metabasalt of Mother Featherlegs, consists of less than 120 m to as much as 1,500 m of
amphibolite or chlorite schist, originally Mg-tholeiite. The lower
half of the formation tends to be massive and locally contains
well-preserved pillows, amygdules, and agglomerate structure.
The upper half is more layered. A distinctive amphibolite containing calesilicate pods occurs locally near the contact with the
formation of Muskrat Canyon. The schist of Silver Springs consists largely of various types of schist, originally largely shales, but
including subgraywackes, graywackes, and clean quartz sandstones. The unit averages 1,200 m thick but is as much as 2,100 m
thick in the northeastern and southeastern parts. It is absent
locally in the central part of the uplift, probably because it is cut
out by an unconformity. The uppermost unit of the Whalen
Group is the formation of Wildcat Hills, which averages 450 m
thick. It consists largely of siliceous stromatoliticdolomite grading into dolomite and chondrodite dolomite. In the northern part
of the uplift it contains interlayers of pelitic schist and layered
amphibolite. Four types of long-ranging shallow-water benthic
microbial stromatolites (Hadrophycus, Stratrjera, Cohmnaefmta, and Gruneriu) are recogmzed in the two metacarbonate
units (Hofmann and Snyder, 1985). The uppermost exposed part
of the formation of Wildcat Hills is cross-cut by the oldest Late
Archean granite.
Large lenses of hematite occur near the contact between the
schist of Silver Springs and the dolomite of Wildcat Hills in
several localities (Carter, 1963 and references cited therein).
These deposits are believed to have been formed by ground-water
oxidation and enrichment of originally ferruginous beds (Bayley
and James, 1973). The hematite pods are most common where
the overlying carbonates are thinnest and where schist is silicified,
suggesting that some of the hematite may have been concentrated
by postdepositional ground-water leaching.
After deposition of the Whalen Group, intrusive episodes
alternated with metamorphic episodes from Late Archean to
149
Wyoming province
Middle Proterozoic. Plutonic rocks now constitute more than 40
percent of the exposed Precambrian rocks. The granite of Rawhide Buttes was emplaced at 2,580 Ma, as determined by a Rb-Sr
whole-rock isochron. Rb-Sr model ages for granitic gneisses in
the outlying areas of Twin Hills, Cassa anticline, and Antelope
Hills suggest that these rocks may have an earlier history dating
back to 2.9 Ga. Swarms of Fe-tholeiite dikes (now granular
quartz amphibolite) intruded the granite of Rawhide Buttes before latest Archean or earliest Proterozoic metamorphism. A
small pluton of the granite of Flattop Butte was emplaced at
1,980 100 Ma. This granitic pluton is nearly devoid of amphibolite dikes, suggesting it post-dated the tholeiitic dike swarm.
Geochemical features of the granite of Flattop Butte, its initial
* 7 ~ r / * 6ratio
~ r of 0.767, Sm-Nd model ages of 2.8 to 3.1 Ga, and
strongly S-type mineralogy and chemistry indicate that this granite formed from a melt of Archean granite and pelite (Snyder and
Peterman, 1982).
Between 1,980 and 1,740 Ma, all rocks of the Harmille
Uplift were regionally metamorphosed and penetratively deformed. Diorite of Twin Hills was emplaced at 1,740 10 Ma
(U-Pb zircon age). At 1,720 40 Ga (Rb-Sr whole-rock isochron), the granite of Haystack Range was emplaced. This granite
contains inclusions of the diorite of Twin Hills. Later events
include emplacement of small pegmatite bodies, local faulting
and cataclasis, resetting of muscovite (RbSr) and hornblende
(K-Ar) ages at 1,610 to 1,670 Ma, emplacement of undeformed
post-tectonic Mg-tholeiite diabase dikes perhaps at around 1,400
Ma, and resetting of biotite K-Ar ages at 1,300 to 1,350 Ma
(Peterman, 1982; Snyder and Peterman, 1982).
*
*
*
and metamorphism associated with both the Sevier orogeny during the Cretaceous, and Basin and Range extension during the
Neogene. This area is part of a core complex analogous to the
Shuswap terrane of northeast Washington and southeast British
Columbia (Armstrong, 1968). Progress has been slow in unraveling the Precambrian history of these complexes because of the
intense Phanerozoic overprints. However, geochronologic work
demonstrates the existence of Archean basement (Armstrong and
Hills, 1967; Compton and others, 1977).
In the Albion and Raft River Ranges, Archean rocks crop
out in a series of domes. Although Armstrong (1968) refers to
those in the Albion Range as mantled gneiss domes, a more
detailed study of-the Raft River Range and Grouse Creek Mountains by Compton and others (1977) suggests that the Archean
rocks there are part of a large autochthon partly concealed by
three generations of thrust sheets containing variably deformed
Paleozoic and Mesozoic metasediments.
Archean rocks in the Albion Range were named the Green
Creek Complex by Armstrong (1968). This complex includes
quartzo-feldspathic gneiss, schist, amphibolite, and quartzite,
which yield a poor Rb-Sr whole-rock isochron of 2,550 30 Ma
(Armstrong and Hills, 1967). Correlative rocks in the Raft River
Range and Grouse Creek Mountains are divided into four mappable units: quartz schist, amphibolite, metamorphosed trondhjemite and pegmatite, and metamorphosed adamellite. According to Compton and othen (1977), the adamellite yields an Rb-Sr
age of 2,460 170 Ma.
*
*
WASATCH MOUNTAINS AND ANTELOPE ISLAND
B. Bryant
i
R. S. Houston and K. E. Karbtrom
The oldest rocks identified in the Black Hills are gneissic
granites exposed near Nemo and in the Bear Mountain dome
(Fig. 16) (Kleinkopf and Redden, 1975). In the Nemo area the
gneissic granite includes the Little Elk Granite, a gray, coarse
grained, gneissic granite. Similar granite in the center of Bear
Mountain dome is associated with coarse-grained biotite schist
and gneiss. Both the Little Elk Granite, and the gneissic granite of
Bear Mountain are dated at about 2,500 Ma (Zartman and Stem,
1967; Rattk and Zartman, 1970).
Granitic rocks with minimum U-Pbzircon ages of 2,600 Ma
occur as screens in intrusive rocks of Tertiary age in the southern
Bear Lodge Mountains of northeastern Wyoming, approximately
10 km northwest of the Black Hills. Staatz (1983) suggests that
the granite bodies may be at the margin of a Precambrian uplift
whose center was replaced by the Tertiary intrusive.
ALBION AND RAlT RIVER RANGES
R. S. Houston and K. E. Karbtrorn
Precambrian rocks in the Albion and Raft River Ranges and
Grouse Creek Mountains (Fig. 17) were involved in deformation
Precambrian crystalline rocks are exposed in the Wasatch
Mountains in central Utah and on Antelope Island in Great Salt
Lake about 30 km to the west (Plate 2). The entire complex may
be allochthonous, for some structural interpretations suggest that
the major eastwarddirected Late Cretaceous to Early Tertiary
thrusts of the IdaheWyoming thrust belt root in a sole thrust that
would underlie the exposed crystalline rocks (Royse and others,
1975). The crystalline rocks compose the Farmington Canyon
Complex (Eardley and Hatch, 1940), which consists of
sillimanite-bearing pelitic schist and interlayered quartzo-feldspathic schist and gneiss. Small lenses and discontinuous layers of
amphibolite are widespread throughout the sequence, and layers
of quartzite are locally common in the southern part of the range.
Protoliths were probably quartz sandstone, feldspathicsandstone,
shale, and tuffaceous sandstone. Pelitic rocks in the southern part
of the range contain muscovite and sillimanite; farther north,
similar rocks are migmatitic and contain sillimanite-microcline
assemblages (Bryant, 1988). A few occurrences of relict hypersthene suggest that the terrane may once have been at granulite
grade.
In the northern Wasatch Range, metasedimentary rocks are
intruded by biotite-hornblende momgranite that contains foliation parallel to that in the wall rocks. The monzogranite is grada-
R S Houston and
Harney Peak Granite
Schist and phyllite
Iron Formation, schist,
and quartzite
I
I
Ea
. ....
. .
Roberts Draw Formation
and equivalents
Quartzite
.
- ....
Estes Formatiin
and equivalents
-
Amphibdite
and metagabbro
"Nemo Group"
Granite gneiss
SCALE
0
20 KILOMETERS
Figure 16. Generalized geologic map of the Precambrian rocks of the Black Hills (after Kleinkopf and
Redden, 1975; De Witt and ofhers, 1986; Redden and French, 1989).
1
',
tional with migmatitic gneiss, and contains large inclusions of
wall rocks, suggesting that it was derived by nearly in situ melting
of the country rocks.
Similar migmatitic gneiss and schist of sillimanite grade are
exposed on Antelope Island. There they have been intruded,
perhaps synkinematidy, by granite that is now foliated and
converted to a pyroxenehornblende gneiss (Bryant and Graff,
1980). Garnet-bearing biotitemuscovite granite gneiss, sheared
and recrystalked under medium-grade conditions, has been recovered in core from a weU about 8 km west of Antelope Island.
The complicated metamorphic history of the rocks of the
Farmington Canyon Complex is reflected in the complexity of
their isotopic systematics (Hedge and others, 1983). Rb-Sr
wholerock data on the metamorphic rocks do not define an
isochron. On an isochron diagram, data points for most rocks fall
between lines that wouid suggest an upper age limit of 3,600 Ma
and a lower age limit of 2,600 Ma, with an initial 87sr/g%r ratio
less than 0.705. The Sm-Nd isotope system also fails to define a
unique age for these rocks, but does indicate that crustal material
as old as 2,800 to 3,600 Ma was included in the metamorphic
complex. The Rb-Sr and Sm-Nd data can be interpreted as indicating: (1) that the age of the rocks is 3,600 Ma and that the
scatter of points is due to disturbance of the isotopic systems
during later metamorphisms, (2) that the age of the rocks is 2,600
Ma, but that they contain various amounts of inherited sedimentary components, or (3) that the rocks are of some intermediate
age and that the scatter of points is due to some combination of
(1) and (2).
Wyoming province
-
0
SCALE
10
2 0 KILOMETERS
151
The &neb&monzogrnite has a whole-rmk Rb-Sr kmchron
age of 1,808 & 34 Ma with an initial 87~/8"Srratio of 0.769.
The very high initial value is consistent with gealogic evidence for
derivation of the rock by melting of Archean rocks. Zircon from
the monmganite gneiss gives a U-Pb concordia age of 1,780 &
20, amistent with the Rb-Sr age. U-Pbdata for zircons from the
layered metamorphicrocks yield a spectnun of a3Pb/%b ages
ranging from 1,770 to 2,271 Ma. These data can be interpreted as
i n d i m that some or 9 of the zircons are older (perhaps much
older) than 2,271 Ma, that they were partly reset d u a a metamorphic event at about 1,770 Ma, and partly reset again during
an event at about 70 Ma. K-Araga ~n biotite and hornblende
range from 1,700 to 224 Ma and psobbly reflect both the Early
Proterozoic metamorphic event and shearing and partial rt!crystaIhation of the rock under gnxnsehist-facie mnditions in the
later Precambrian or the late Mesozoic or both. '
Zirconsfrom thegranite gneiss on Antelope Island and from
the core from the well to the west have U-Pb concordia ages of
about 2,020 Ma, distinctly m r e n t fitom themoonmgrdte gneiss
of the W a t c h Range. In fact, very few rocks of that age have
been reported frpm anywhere in the w a r n United States.
Bryant (1988) interprets the chemical and hobpic data as
indicating that the sedimentary and psssibly volcanic rocks of the
Farmingtoli Canyon Complex were depdted on oceanic crust
adjmnt to a continental margin of unknown trend prior to 2,600
Ma and that they were initially metmnorphmed at about 2,600
Ma. AmphihWfaciesmetamorpk and migmatization of the
rocks now ~ x p e in
d the W-h
Adomtab took place at
about 1,800 Ma, but m - b
and emplacement of granitic
rock in the part of the tenme now exposed on Antelope Island
may have taken place 200 m.y. or more earlier. The 2,020 Ma
age of the granitic g n e h on Antelope Island and in the subsurface to the west may reflect the eE.ects of Protermic deformation
on the margin of the aaton (Farmer and DePaolo, 1983,1984;
Stacey and Zubm, 1978).
LlTlW, BBLT AND LiTI'LE ROCKY MOUNTAINS
EXPLANATION
R.S. H o d n and K. E. Karbtrom
Lofe Precombrion ond
Phonerozoic
metosedimentary rocks
f
-
Green Creek Complex
Eor/v P#ferozoic
In the Liae Mt Mountains of west-centralMoatana (Plate
2), Catware a d gulp (1964) describe gnekes and scW t b t
form the bwement beneath the Middle Proterozoic Belt SuperArcheon granite
group. The rocks include paiaga;eiss, migmtite, grade gneiss,
and chlorite and biotite schists, all of which are intruded by
Late Precambrian
schists ond gneisses
mWorite and basalt dikes. They f m d that zirc0115 from the
Archeon metosediments
migmtite and whist are older than 2,450 Ma.
The Little Rocky Mountah contain Precambrian rocks
within
and marginal to Tertiary syenite stocks (Peterman, 1981).
Elba Ouartzite
The Precambrian rocks congist of hterlayered biotite schist,
Early Proterozoic
Low ong/e foult,
qmm-feldspathic gneiss, and minor quartzite and ampb'bolite,
teeth on upper plote
Figure 17. Generalized geo1ogic map of the Albion and Raft River presumabIy metamorphosed sxbmtaq and volcanic rocks.
Ranges and Grouse Creek Mountains (after Armstrong, 1968; and Seven samples including biotite gneiss, graniic gneiss, a m p h i
lite, biotite-hornblende granodioritic gneiss & h e an imprecise
&mpton and others, 1977).
r-
R S. Houston and Others
whole-rock Rb-Sr isochron that indicates an age of about source of uranium in the Tertiary deposits of the Wyoming prov2,550 Ma. Calculations based on a typical Archean initial ince. The granite in the Owl Creek Mountains is fractured, and
8 7 ~ r / 8 6 ~ratio
r of .701 suggest that the time of major meta- pitchblende veins are developed in the fiacture system, but the
pitchblende may not be magmatic.
morphism may have been substantially earlier.
Most mineral deposits in Archean rocks of the Wyoming
province are in supracrustal SLICXZ&O~~S. Between 1.2 and 1.6
MINERAL DEPOSITS
million tons of iron ore have been produced from the Nemo Iron
Formation in the eastern Black Hills (De Witt and others, 1986),
which is part of the Archean supracrustal succession cut by the
Archean rocks affected by granulite-facies metamorphism at 2,500 Ma Little Elk Granite. Iron formation has been identified
some stage in their history may not be as promising for mineral in almost all Archean supracrustal successions, and mines were in
exploration as lower-gradeArchean rocks, but experience in sim- operation until 1984 in the Hartville Uplift and in the South Pass
ilar areas (Anhauesser, 1981) indicates that mineral deposits in area in the southeastern Wind River Range. With the exception
granulite tenanes are most likely to occur in the metasedi- of small-scale production of talc and chlorite in the supracrustal
mentary-metavolcanic units. Potential deposits may include rocks of southwestern Montana, there are no mines currently
chromite in ultramafic bodies, iron ores in iron formation, and active in Archean rocks in the Wyoming pro+ce. The supranickel sulfide and platinum-group metals in mafic bodies. If most crustal rocks have not been thoroughly explored since the turn of
of the terrane underwent granulite-faciesmetamorphism early in the century, however, and there is a good potential for stratiform
its history, the rocks may be impoverished in such elements as K, sulfide deposits in several of the supracrustal successions (Hausel,
Rb, Th, and U (Tarney, 1976) and any mobile mineral concen- 1982). Gold-tungsten mineralization of the Jardinecrevice
trations may have been destroyed. While it is inappropriate to Mountain area in the Beartooth Mountains, which is probably
recommend the core of the Wyoming province as a region for stratiform in origin, is one of the prospects that may have evenmineral exploration, Late Archean intrusives such is the Still- tual economic potential (Seager, 1944). Mineralized quartz veins
water Complex have substantial promise for economic minerali- and shear zones in the supracrustal rocks, such as those of the
zation. The Stillwater Complex contains large reserves of South Pass area (Bayley and others, 1973), may have current
high-iron chromite, and nickel, copper, and platinum sulphides. economic possibilities for small precious-metal mines, but no
The mineralization in the basal u l t r d c zone of the wmplex large deposits have been identified.
has been known for more than 30 years (Howland, 1955),but the
discovery of the Merensky Reef-type sulphide layers containing PROTEROZOIC ROCKS
significant platinum values in the upper part of the complex was
not announced until the 1970s (Page and others, 1976). The R. S. Houston and K.E. KarIstrom
high-iron chromite and sulphides at the bottom of the Stillwater
Complex constitute large low-grade reserves and are not of imEarly Proterozoic metasedimentary rocks are only exposed
mediate economic interest, but the platinum-rich sulphide layers along the southern margin of the Wyoming province (Plate 2).
in the upper part are being investigated as potential mineral They include the Snowy Pass Supergroup of the Medicine Bow
Mountains, the Snowy Pass Group of the Sierra Madre, the Red
deposits.
Late Archean felsic intrusions can be divided into an older Creek Quartzite af the Uinta Mountains, Proterozoic rocks of the
group of tonalites and granodiorites with probable ages of 2,600 Black Hills,and pomi'bly rocks exposed in isolated outcrops
to 2,700 Ma, and a younger group of he-grained, high- within the Wo-Wyoming thrust belt. Although of limited expotassium granites with probable ages of about 2,500 Ma. No tent, the Proterozoic rocks are well preserved, especially in the
major mineral deposits have been found associated with the older Medicine Bow Mmtains, and record the Early Proterozoic sedintrusions, although modest-grade molybdenite and other sul- imentary and tectonic history of the margin of the Wyoming
phide mineral pockets occur in disseminated mnes in an intrusion province between 2,400 and 1,900 Ma.
near Schiestler Peak in the southwestern Wind River Range
(Benedict, 1982). Late Archean to Early Proterozoic granites are SNOWY PASS SUPERGROUP OF THE
found in the Black Hills, Owl Creek Mountains, Granite Moun- MEDICINE BOW MOUNTAINS
tains, Wind River Range, Teton Range, Seminoe Mountains,
Freezeout Hills, northern Laramie Mountains, northern Medicine R. S. Houston and K.E. KarIstrom
Bow Mountains, and northern Sierra Madre (Plate 2). No e m
nomic mineral deposits have been identified in these granites but
The Snowy Pass Supergroup in the Medicine Bow Mounthey are believed to have major economic significance. They are tains includes the Deep Lake and Libby Creek Groups, a metauranium-, thorium-, and potassium-rich and may have been sedimentary succession more than 10,000 m thick. The Deep
sources of uranium &om Archean to the present (Houston, 1979; Lake Group unconformably overlies Archean rocks, including
Stuckless, 1979). They are interpreted as a direct or indirect the Phantom Lake Metamorphic Suite and granite. Figure 10
Wyomingprovince
shows the stratigraphic units that comprise the Deep Lake and
Libby Creek Groups. The following description of these rocks is
summarized from Karlstrom and others (1983) and references
cited therein.
The Magnolia Formation is the basal unit of the Deep Lake
Group. It includes two members: a basal conglomerate member
containing muscovitic quartz-pebble conglomerate and paraconglomerate, and a quartzite member containing more mature subarkosic and arkosic quartzite and quart-granule conglomerate.
The conglomerate member crops out discontinuously in the cores
of anticlines and the limbs of synclines throughout the 35-kmlong outcrop area of the Deep Lake Group. Typically, the unit
occurs as lenticular zones dominated by polymictic paraconglomerate with interbeds of quartz-pebble conglomerate and
coarse-grained arkosic quartzite. At the northern limit of the
outcrop, the conglomerate member is radioactive and is dominated by muscovitic quartz-pebble conglomerate and coarsegrained muscovitic arkose. The quartz-pebble conglomerate is
pyritic and occurs in lenses ranging from the thickness of a single
pebble to compound zones 2 m thick.
The conglomerate member grades laterally and upward into
the quartzite member, which is troughcrossbedded, coarse
grained to granular, micaceous subarkose to arkose. Both the
conglomerate and quartzite members of the Magnolia Formation
are fluvial. The paraconglomeratesare interpreted as deposited in
alluvial fan systems, and the quartz-pebble conglomerate and
trough-crossbedded quartzite as deposited in braided rivers. The
radioactive minerals are believed to be fossil placer accumulations in river channels. The Magnolia Formation is overlain conformably by the Lindsey Quartzite, a trough-crossbedded.fluvial
quartz arenite and subarkose with thin phyllite.
The Campbell Lake Formation is a discontinuous paraconglomeratephyllite succession. The paraconglomerate contains
poorly sorted, subangular clasts of granite, quartzite, and phyllite
in a poorly sorted, micaceous arkose to subarkose matrix. The
origin of the Campbell Lake is uncertain. It may represent a
debris flow of some type or it may be a glacial deposit.
The Cascade Quartzite is an extensive unit as much as 850
m thick that overlies older beds unconformably. The Cascade is
characterized by clean, white, massive, pebbly quartz arenite and
subarkose. Quartz and black chert pebbles are in distinct layers as
much as 10 cm thick and locally occur in conglomerate beds 3 to
6 m thick. The Cascade Quartzite is probably of fluvial and
shallow marine origin.
The Vagner Formation is a three-fold unit consisting of
diamictite at the base (Fig. 18A), a middle unit of marble (Fig.
18B), and an upper unit of interbedded quartzite and phyllite.
The formation is interpreted as glacial.
Rocks of the Deep Lake and Libby Creek Groups are separated by a fault, but the amount of offset is uncertain.
The Rock Knoll Formation is the basal unit of the lower
Libby Creek Group. It is predominantly medium-grained
plagiocbrich arkose, but also contains phyllite and conglomerate. The conglomerate contains clasts of quartz, quartzite, and
i
153
granite. The Rock Knoll is interpreted as representing a glacial
retreat between episodes of glaciomarine sedimentation represented by the underlying Vagner Formation and overlying
Headquarters Formation.
The Headquarters Formation consists of a lower succession
of lenticular paraconglomerate with dropstones (Fig. 18C),
quartzite, and schist and an upper succession of laminated
(varved?) phyllite. It is interpreted as glaciomarine.
The Headquarters Formation is overlain conformably by the
Heart Formation, a quartzite unit with local beds of phyllite. The
Heart Formation is interpreted as prodelta and delta-front sediments associated with a prograding tide-dominated delta.
The upper part of the lower Libby Creek Group is a
quartzitedominated succession of three formations. The Medicine Peak Quartzite is medium to very coarse grained and contains pebbly zones and beds of quartz-pebble conglomerate; it is
approximately 1,600 m thick. The Lookout Schist is composed of
about 370 m of interlayered phyllite, subarkosic, arkosic, and
argillaceous quartzite. The Sugarloaf Quartzite (Fig. 18D) is
largely medium-grained quartz arenite about 600 m thick. The
lithology and primary structure of these three formations suggest
that they are deltaic, with the Medicine Peak and Sugarloaf repre
senting parts of a delta plain and the Lookout Schist representing
the delta front and prodelta.
The upper Libby Creek Group is in contact with the lower
Libby Creek Group along a major fault that has removed segments of the Sugarloaf Quartzite. This fault may have been a
thrust with large displacement that has been subsequently rotated
to near vertical. The upper Libby Creek Group must have been
deposited farther offshore than the lower Libby Creek Group.
The upper Libby Creek Group is divided into three formations:
the Nash Fork Formation, the Towner Greenstone, and the
French Slate. The Nash Fork Formation consists chiefly of
tan metadolomite with thick lenses of black phyllite. The metadolomite contains stromatolitic bioherms (Fig. 18E and F) in a
variety of shapes and sizes (Knight, 1968). The formation may
have been deposited on a shallow marine platform or carbonate
bank. The Towner Greenstone is predominately massive to
schistose amphiilite, locally with possible pillow structures and
lenses of coarse-grained sandstone and hegrained quartzite. The
interbedded sandstone lenses and possible pillows suggest a subaqueous origin for the Towner, but an intrusive origin for parts of
t eliminated. The French Slate is laminated
the unit c a ~ o be
black ferruginous and graphitic slate, and phyllite containing layers of hematite-cemented quartzite. Map thickness of the French
Slate is about 600 m, but an unknown thickness is removed by a
major fault (Houston and others, 1968). The French Slate was
probably deposited in a deep marine basin.
The ages of various units in the Snowy Pass Supergroup are
only loosely constrained. The Magnolia Formation, the basal unit
of the Deep Lake Group, lies on pink granite in the northeastern
Medicine Bow Mountains. The granite is undated, but may correlate with pink orthogneiss of the Sierra Madre dated by Premo
(1983) by the U-Pb zircon method as 2,665 to 2,683 Ma, or with
R S. Houston and Others
Figure 18. Primary structures in Early Proterozoic rocks in the central Medicine Bow Mountains.
A. Deformed marble in the Vagner Formation near Twin Lakes. B. Diamictite in the Vagner
Formation near Twin Lakes. C. Granite dropstone in Headquarters Formation near Twin Lakes.Clast
is approximately 2.5 cm long. D. Rhomboid ripple marks in Sugarloaf Quartzite. E and F. Algal
structures in Nash Fork Formation.
Baggott Rocks Granite of the western Medicine Bow Mountains
dated by him as 2,429 4 Ma. If so, the Snowy Pass Supergroup
is younger than about 2,430 Ma. The Gaps Intrusion, a small
pluton of leucocratic quartz diorite that cuts the Sugarloaf
Quartzite, has been dated at 2,075 75 Ma by C. E. Hedge
(personal communication, 1982) using the Rb-Sr wholarock
method. Thus, the lower Libby Creek Group is pre-2,000 Ma.
*
*
*
*
The Lookout Schist was metamorphosed at 1,710 60 Ma, and
the French Slate was metamorphosed at 1,620 425 Ma, as
determined by Rb-Sr whole-rock isochron (Hills and others,
1968). A probable age for the Deep Lake and lower Libby Creek
Groups is thus between 2,429 and 2,075 Ma, and a probable age
of the upper Libby Creek Group is between 2,075 and 1,700 Ma.
A continuing problem in evaluating relative ages of these groups
Wyomingprovince
is uncertainty of the magnitude of movement on inferred thrust
faults (Fig. 9). At present, we do not believe that any of the
thrusts inverts the stratigraphic order.
SNOWY PASS GROUP OF THE SIERRA MADRE
R. S. Houston and K. E. Karhtrom
i
+
W l y Proterozoic metasedimentary rocks of the Sierra
Madre (Fig. 9) are more deformed and metamorphosed than
those of the Medicine Bow Mountains, but rocks equivalent to
both the Deep Lake and Libby Creek Groups are present (Fig.
10). The Magnolia Formation is mostly quartzite, with local beds
of radioactive quartz-pebble conglomerate at the base. There are
no basal conglomerates or obvious regional unconformity as in
the Medicine Bow Mountains, and it is difficult to distinguish the
Magnolia Formation from underlying Bridger Peak Quartzite of
the Phantom Lake Metamorphic Suite. However, generally
coarser grain size and the presence of trough-crossbedding are
characteristic of the Magnolia.
The Singer Peak Formation is chiefly phyllite with minor
quartzite and conglomerate, and the Cascade Quartzite consists of
quartzite with layers of quartz-pebble and blackchert-pebble
conglomerate, as in the Medicine Bow Mountains. A U-Pb zircon
age of 2,092 rt 4 Ma on a pegmatitic phase of a metagabbro that
cuts Cascade Quartzite in the Sierra Madre (Premo, 1983) correlates well with that of the Gaps Intrusion of the Medicine Bow
Mountains and establishes a minimum age for deposition. It also
suggests that gabbroic sills and dikes throughout the Snowy Pass
Supergroup are related to the same Proterozoic tectonic event.
The Bottle Creek Formation is a heterogeneous succession
consisting of quartzite, phyllite, angular-clast paraconglomerate
(diamictites), and metacarbonate. The Bottle Creek may correlate
with the Vagner Formation or with parts of the Rock Knoll,
Headquarters, or Heart Formations of the Medicine Bow Mounbins (Fig.10).
The Copperton and Slaughterhouse Formations, preserved
in incomplete sections in fault slim in the Sierra Madre (Fig. 9),
are also possible correlatives of parts of the Libby Creek Group of
the Medicine Bow Mountains. The Copperton Formation consists of a lower, coarse-grained, highly sheared, kyanitsbearing
quartzite correlated with the Medicine Peak Quartzite; a middle
laminated unit consisting of alternating quartzite and phyllite
correlated with the Lookout Schist; and an upper sheared quartzite correlated with the Sugarloaf Quartzite. Its thickness is
The Slaughterhouse Formation is composed of metadolomite, quartzite, phyllite, and metachert. It is correlated with the
Nash Fork Formation (Fig. lo), although none of the stromatolites characteristic of the Nash Fork have been recognized in the
Slaughterhouse. The Slaugherhouse is fault bounded, and the
stromatoliticsection may have been removed.
- .
.- * . .
. a , ~' r-: ~
RED CREEK QUARTZITE
The Red Creek Quartzite is exposed in a series of fault
blocks in the northeastern Uinta Mountains (Fig.
. - 19).
. The formation consists of alternating massive white quartzite and phyllite
units (Hansen, 1965; Graff and others, 1980). The phyllite units
are composed of alternating beds of phyllite and quartzite, and
locally contain beds of marble. The exact age and depositional
environment of the Red Creek Quartzite is uncertain. It is in fault
contact with quartz-feldspar gneiss and related rocks which Graff
and others (1980) called the Owiyukuts Complex and which they
dated at 2,700 Ma. However, recent studies by G. A. Swayze
(written communication to Reed, 1989) have shown that the
Owiyukuts Complex is composed entirely of potassium-metasomatized Red Creek Quartzite. He quotes C. E. Hedge (oral
communication, 1987) as stating that the 2,700 Ma Rb-Sr date
on the complex is meaning1ess, and suggests that if the correlation
of the Red Creek Quartzite with the lower Libby Creek Group of
the Medicine Bow Mountains is correct (Graff and others, 1980),
a more realistic date for metasomatism and metamorphism is
around 1,800 Ma. Swayze has also found that the amphiilites
that Hansen (1965) interpreted as post-metamorphic were actually emplaced before folding and metamorphism. The 1,550 Ma
K-Ar age on hornblende from the amphibolites (Graffand others,
1980) is interpreted as a cooling age following the peak of regional metamorphism.
BLACK HILLS
R. S. Houston and K.E. Karktrom
The Black Hills uplift of southwestern South Dakota (Plate
2; Fig. 16) is a broad Laramide anticline with Precambrian metasedimentary, metavolcanic, and plutonic rocks exposed in the
core. The metasedimentary rocks of the Black Hills have been
studied in detail only in limited areas, and the relationship between areas is not fully understood. A modem compilation of the
Precambrian geology of the entire uplift by Kleinkopf and Redden (1975), a review of the general geology by Bayley and James
(1973), and a summary of the uranium potential of the Nemo
area by Redden (1980) serve as the basis for this discussion.
The central Black Hills consist of a sequence of metamorphosed graywacke, shale, basalt, chert, and iron formation shown
as eugeoclinal rocks in Figure 16. Two horizons have been dated
as 1.89 to 1.97 Ga (Chapter 2). These rocks are intruded by the
1,740 Ma Harney Peak Granite in the southern Black Hills, and
by amphibolite and pegmatite dikes throughout the area.
The Archean rocks previously descn'bed appear in two
domes, each of which is surrounded by quartz-rich metasedimentary rocks that are quite different in character and are probably
older than the eugeoclinal rocks of the central Black Hills. Thus,
R S.Houston and Others
--
--
UTAH
SCALE
0
WYOMING
1 COLORAD?
.
5
10 K I L O M E T E R S
I
I
EXPLANATION
Figure 19. Generalized geologic map of the Precambrian rocks of the northeastern Uinta Mountains
(after Graff and others, 1980).
the Proterozoic rocks may occupy a north-northwest-trending
synclinorium. Redden (1980) believes that the older quartz-rich
succession adjacent to the Archean domes are platform-type
metasediments of Early Proterozoic age.
The Little Elk Granite of the Nemo District, dated as 2,500
Ma by Zartman and Stem (1967), is interpreted by Redden
(1980) to be in fault contact with overlying metasedimentary
units. However, he suggests that most of the metasediments are
not only younger than the Archean granites, but deposited unconformably on them.
The oldest rocks in the metasedimentary sequence in the
Nemo District are shown as the "Nemo Group" on Figure 16.
The basal part of the sequence consists of an unnamed schist unit
and the Nemo Iron Formation, a chert-hematite-magnetiteiron
formation interbedded with phyllite. The overlying Boxelder
Formation contains chloritic paraconglomerate of the Greenwood Tongue, radioactive conglomerate of the Tomahawk
Tongue, and crossbedded quartzites. The uppermost unit in the
"Nemo Group" fs the Benchmark Iron Formation, a thin iron
formation prese~edlocally beneath an unconformity at the base
of the overlying Estes Formation, a sequence of coarse conglomerate and quartzite. The overlying Roberts Draw Formation is a
dolomitegraphiticphyllite unit, which marks the end of deposition of quartz-rich platform deposits in the Nemo area. The amphibolite and metagabbro and quartzite units of Figure 16 are
considered part of the Vanderlehr Formation by Redden and
French (1989) and are believed by them to be correlative with
the Estes and Roberts Draw Formations.
I
I
157
Wyoming province
Redden (1980) believes the Greenwood Tongue of the Boxelder Formation is an alluvial fan sequence derived from the west.
He interpreted the Tomahawk Tongue as fluvial deposits derived
from the northeast. The entire Boxelder Formation is a finingupward sequence that represents a generally transgressive depositional condition that eventually gave rise to shallow marine
depositional of the Benchmark Iron Formation.
The lenticular distribution of various lithofacies, the variable
source areas for the conglomerates, and the fining-upward character of the sequence suggest deposition in a small but progressively widening fault-bounded basin. These features are very
similar to those seen in the basal Deep Lake Group of the Medicine Bow Mountains, where deposition probably took place during the early stages of r i f t i i along the southern margin of the
Wyoming province (Karlstrom and Houston, 1984). A rift setting for the Boxelder Formation is also possible. The Blue Draw
Metagabbro, a gabbroic sill that intrudes the Boxelder Fornation,
has been dated as 2,090 10 Ma by R. E. Zartman (Redden,
1980). This is about the same age as tholeiites that intrude the
Deep Lake Group and that are interpreted as related to rifting
(Karlstrom and Houston, 1984).
Redden (1980) suggests that the Estes Formation was deposited in deepwater fans shed from active basin-bounding
growth faults across older units, which were already folded and
faulted. The Estes Formation apparently represents a continued
trend toward deeper-water sedimentation in a fault-bounded
basin, a trend that eventually gave rise to the deepwater shales
and graywackes of the eugeoclinal sequences in the central Black
Hills. This structural and sedirnentologic setting is typical of a
fault-bounded rift basin, and it is possible that the Black Hills
succession was deposited in a northwest-striking rift basin (Houston, 1986). If so, the rifting began prior to 2,090 Ma and ceased
before intrusion of the Harney Peak Granite (1,740 Ma).
There are no operating mines elsewhere in Early Proterozoic
rocks in the Wyoming province. At the turn of the century, a
number of gold prospects were identified in the Medicine Bow
Mountains, and a major copper deposit was discovered in the
Sierra Madre. The gold prospects were in quartz veins near or at
contacts between quartzite of the Magnolia Formation and large
gabbroic sills. The main mining district was the Gold Hill district
(Fig. 9) in the central Medicine Bow Mountains (Houston and
others, 1968). The copper deposits of the Sierra Madre were
found in rocks of the Snowy Pass Group and were said by
Spencer (1904) to be in veins and fractures in quartzite. He
suggested that the copper was derived from the numerous gabbroic sills that cut the Proterozoic rocks, but Houston and others
(1975) could not demonstrate a correlation between the copper
mineralization and gabbroic intrusions. There wasonly one highgrade copper deposit in the Sierra Madre, the Ferris-Haggarty
Mine (Fig. 9), and since this mine closed there has been no
further production from the Encampment mining district. The
Ferris-Haggarty ore body may have been a stratiform deposit.
In the late 1970sand early 1980s, interest in developingnew
uranium reserves for the United States was intense, and a number
of Precambrian successions were prospected for worldclass uranium deposits. Substantial uranium resources were discovered in
quartz-pebble conglomerate in the Magnolia Formation in the
Medicine Bow Mountains, but the grade was too low to be of
economic interest (Karlstrom and others, 1981). Geochemical
anomalies and the presence of uranium in shear zones and veins
suggest the northern Medicine Bow Mountains may host other
types of Proterozoic uranium deposits and may be a good area for
continued exploration should the price of uranium rise significantly (Houston and others, 1984).
MINERAL DEPOSITS
G. L. Snyder and R. S. Howton
R.S. Houston and J. C. R e 4 Jr.
Mafic and ultramafic intrusive rocks are reported from
nearly all parts of the Wyoming province. Ages span most of the
time between 700 and 3,000 Ma, but there seem to be gaps near
950 200 and 2,350 100 Ma, hiatuses apparently also present
in Canada (Gates and Hurley, 1973). Some early K-Ar age determinations are much younger than modern Rb-Sr, Sm-Nd, and
U-Pb determinations on the same rocks. Most matic intrusive are
dikes, and most ultramafic intrusives are elongate lenses. Dike
orientations are dominantly northeast in the southeast half of the
province, where most dikes may be Archean, and northwest in
the northwest part of the province, where most dikes are probably Proterozoic (Plate 2). This difference in orientation may reflect ditlerent stress regimes. However, Wooden (1975) notes that
"the dike swarms of Montana differ from those of the Canadian
Shield in that they are more limited in occurrence and generally
not separable into compositional or age groupings on the basis of
strike." Several workers have noted that zones of weakness occupied by early dikes have been reoccupied, sometimes repeatedly,
*
I
I
Proterozoic miogeoclinal rocks of the Black Hills are host to
some of the most important mineral deposits in the Wyoming
province (De Witt and others, 1986). These include the banded
iron formations in the Boxelder Creek and Estes Formations in
the Nemo area (estimated to contain reserves of 250 to 330
million tons), uraniferous quartz-pebble conglomerates in the
Boxelder Creek Formation near Nemo, and, most importantly,
the syngenetic stratiform precious metal deposits of the Lead
District in the northern Black Hills. The latter are auriferous
zones within carbonate- and oxide-facies iron formations in the
Homestake, Montana Mine, and Rockford Formations. They are
apparently controlled by hot springs, biologic actions, or sedimentation that contributed gold, arsenic, and minor base metals
to the strata. More than 37 million ounces (1,150 metric tons) of
gold have been produced from the Homestake Mine in the Lead
district.
MAFIC AND ULTRAMAFIC INTRUSIVE ROCKS
*
*
-.-.
..
R S. Roixmn d CMws
158
by later dikes (Prinz,1964;Armbmtmacher, I=, Wooden and
oztaets, 1979).
Mapping of d c btrdves, espechlly dikes, has generally
~ ~
not been sup@men&d by ~ o a 0 or 1gcmchmxioal
d y to the extent necesasry to interprt3f
wamdeby Prim(1964)hrtheBm
he
recopbed two Nods of Anhean and one perid ofhotercmie
dike emplacmmt. T
Y
m
I dike can be d&hg&hed by orhagtion aad &emistry. If s i m h &es of dike sets are m& eke
where in the Wyoming province, it may be possible to interpret
their
sigdkmce.
Detailed &cwsins of the 8eolqjy and chemisfq of d c
@us
rocks of the Wyoming province are in Snyder and others
(1989, 19'90).
formed when rocks to tibe m
tEr were tlmtn&wr~ud ova er
m
m
a
@
c
~
~
t
later in the defbrmtkd
1987).
The Cheyenne belt has been in^^ as a suture dong
whi& klmd am W developed to thi:south
atiwbd w the
rifted nwgh of an A r k mntinmt probably &$ween l,&OQ
and 1,700 Ma @%Usand Houston, 1979;aiarlsttom aad Houston,
1984). Hi& and Hkwton (1979) reported thaS 1,73Ma graites
fncheSim&drewdr@~Mh~elgain&mzoaeand
that massive 1,645 Ma ~ P k igd8eis mimat the dmw zone.
CHEYENNE BELT
R. S; Houston and K.1% KwLsbont
The Cheyenne belt is i xone of nny1oniW rocks in southern Wyoming that separates AreBe~nrmh of the Wyoming
provlaGe from Protermic rocks eqmd to the south (RW2).
In the Makine Bow Mountah, it is a zone as muoh as 10 k133.
wide made up of northatt-sbikhg, subvertical, ~]ikyZio&e
zones tbn, tbwsfs ha ttre Snowy P w
that separatepmtmively d e f o d blocks of vmbw fithob@ o m more stmdw breland thwt belt, W:h may have cwas Ear north as the Cbaite
(Fi.9). Ikfhgmetasxhem rocks of Late k h a n erecl be mut%srnWym& pro*
and M l y Protemic age lie on &&em basement north of the &un&ins. They suggested that burial af Archa bemmt by
mad subsequent upW and swim at a b t 1$300
belt; to the south are eugeoclhd r t l e t a r y and metavol- ~~t
canic rodccr of Early Proterozoic
with no h o r n ArchIkk a@ lowered K-Axmineral ages in the southern mar& of
bmanent. En the Sierra Madre, the Cheyenne belt is a single the Wpo&g province (Peteman and IBWeth, 197%).
arcme l~ylonitem e stdchg generally eastdwest, but s u d ria
other mqxctti to the Medicine Bow QBQm- The p ~ o j d o nof
the Cbpwne'helt in t
k J a d e Ivkmtdns is matpied by
post-tatazk mh*, but 'anoudier of: Raxdwkm racks in
the Whw Hilb f*. 12) contains a zpne of m%beast-s-g
myloni& &st Chaff rwnd others (1982) bave intwp~&xias the
eastemmost &xt@ndunof the Cheyenne I&, Bower, gnyder
(1984) found that tltb mylonib zone is enthly w&h the Arc h ~suggshg
,
that it is not the continuation ofthe Cheyenne
belt.
ate m y known, and exi9thg matmw map
The Cheyme belt is a funWenta1g&~il~gic
diiwxmtintaiq,
t
h Mdm (Fig. 20)are bwd on a very fkw
and its i n ~ ~ idiuences
o n our view ~ f f ; h gJ3arly Froteromk re&m@or I h s of vdsrkde quality. Not dl of these lines are
tectmic evolution of the southern W y o province
~
(KmWrom
and Howbn, 1984). It is also a
in deep mtal
&own nust lx &&
&hampand
structure, as shown by s W c and gravity studies (Albendinger ma&&h cmtd veloc@ stm&m are highly specmldve. Nevand others, l9.Q Johamn and others, 1984). These st.wti~e9srtg- eqth~ka,&ws data s-t
M
r
r
b & p b d 33 to 41 1Em in the
g a t that crust south of the belt h thicker mdfor less d e m
~
~
pMcker cmO
v (50 Ism)
~ to the,north, esfst, and
that to the north. It is also a metamorphic discmlimity along -5
ansl sbinner crust (25 km) to the wesa 11-$ 6@f~
Lamaide
which upper amphibolhEacies rockg a d tnignstites to the am4 p m & m deformation, relrrtiyrdy abmpt
ia
south arejw&posed with gmmscWW rmbwdime to tbe cxmtd tgicseem to coiaei* with rhe bcxmbdes of tile
north (Duebdorfer md Houstan, 1986). titruCtur4 studies d pmoinm. In particularPthe c h g e in cmsM thicknessin southern
mylonites &ow several
of rnovanent. S t q I y p1w&g Wyoming ruugbly miaddm with the Chepam beit d with a
hatiam and quartz c-axis &btic% b v e been interpreted as indi- mjoc gravity gdii&Irt. This Zed J o b mil &ers (1984)to
catkg mvergmce acrm the belz These feetmay haw suggest Phrts variation in mtdt h i b aad m p i t i o n m y
~~
Wyomingprovince
SCALE,
0
4 0 I(IL0METERS
Figure 20. Present crustal thickness in the central United States (modified from Allenby and Schnetzler, 1983). Contours in kilometers; contour interval 5 km.Dots indicate control points where thicknnless has been
determined by seismic refraction.
be in part a relict of Proterozoic deformation across the belt.
Thicker crust in Colorado was interpreted as due to thickening
during accretion by northward thrusting of Proterozoic arc sequences, and variations in density as due to juxtaposition of different crustal blocks. The decrease in crustal thickness westward
into Utah and Idaho presumably reflects the rather abrupt transition into the zone where continental crust was thinned during
Tertiary extension of the Basin and Range province (Braile and
others, 1974). Crustal thickness in the northern Wyoming province is very poorly known. In east-central Montana and the
Dakotas crustal thicknesses as much as 55 km have been reported
(Warren and Healy, 1973) while estimates in southwestern Montana and Idaho range between about 33 km (Sheriff and Stickney, 1984) and 45 km (Braile and others, 1974).
The Bouguer gravity field over the western United States
generally reflects Mesozoic to Recent tectonic features (Eaton
and others, 1978). The only feature of the gravity field that may
be related to Precambrian tectonic features is the pronounced
southward 50 to 100 mgal decrease in southern Wyoming, which
coincides with the Cheyenne belt (Fig. 21). Johnson and others
(1984) suggested that this anomaly is related to Proterozoic de-
159
formation for the following reasons: (1) it is intermediate in scale
between the regional negative anomaly centered over the continental divide (Hildenbrand and others, 1982), and local
perturbations related to mountain uplifts, and is at a high angle to
the regional gradient; (2) mean elevations along the profiles are
relatively constant, so that the anomaly cannot be explained as
due to Cenozoic crustal thickening; (3) the gradient is apparent
on north-south gravity profiles entirely within the Precambrian
uplifts, so that effects of local Laramide uplifts and sedimentary
basins are minimal, and (4) the gradient coincides with the trace
of the Cheyenne belt. Gravity models (Johnson and others, 1984)
suggest the crust in Colorado is less dense and/or thicker than the
crust in the Wyoming province, and that this change takes place
along the Cheyenne belt. Thus, the thicker crust in Colorado,
which is also shown by refraction data, may partly reflect Proterozoic tectonics, as well as superimposed Laramide crustal
structure.
Aeromagnetic data for the Wyoming province have recently
been compiled by Rush and others (1983). The central part of the
province (Plate 2) is dominated by arcuate anomaly trends. These
anomalies, although discontinuous, have high intensities and
sharp gradients and outline a circular area some 450 km in diameter that roughly coincides with the area in which K-Ar
radiometric ages are older than 1,800 Ma. The area is bordered
by a zone characterized by lower-intensity magnetic anomalies
and weaker magnetic gradients. The change in magnetic charao
ter from core to margin is more conspicuous than are the margins
of the Wyoming province itself. The borders of the Wyoming
province are particularly obscure on the northeast, northwest,
and west. On the northeast there is a gradual transition into
northwest-trending anomalies characteristic of the concealed
Trans-Hudson orogen. Magnetic data are sparse along the western border of the province; highly magnetic Tertiary volcanics
obscure basement trends along the northwestern part of the
border. The southern boundary of the Wyoming province is
marked by a strong magnetic anomaly coincident with the
Cheyenne belt (Indelicato and Karp, 1982) and by several parallel anomalies that are both concentric to the core and parallel to
the Cheyenne belt. Parallel trends south of the Cheyenne belt
reflect northeast-striking foliation, lithologic contacts, and shear
zones within the accreted Proterozoic terrane.
Magnetic trends are also of some use in outlining more local
basement trends. Many of the Precambrian-cored uplifts of the
Wyoming province are associated with positive magnetic anomalies bordered by steep gradients that correspond to Laramide
range-bounding faults. Examples of such anomalies are in the Big
Horn Mountains, Wind River Mountains, Owl Creek Uplift, and
Uinta Mountains (Plate 2). Some other magnetic features of the
uplifts can be directly related to magnetic lithologies, for example,
exposed iron formation in the Black Hills, anorthosite in the
Laramie Range, layered mafic complexes of the Medicine Bow
Mountains and Sierra Madre, and the Atlantic City greenstone
belt of the southern Wind River Mountains are all expressed as
prominent positive anomalies. A third group of anomalies indi-
R S. Houston and Orhers
MJ
1
I
-1.b
-130
I
-$
I
I
0
I
I
40
I
,
M
DISTANCE FROM CHEYENNE
~~
I
I
IN)
I
Figure 21. Gmity may of the Wyomiag
a d
awes the *ems
belt.
~ . B i o ~ ~ ~ i t y ~ g a y &m250Laoa(3tnIBdap ~ ~ ~ ~ ~
brand athers, 1982). B. Gravity pd&s and ge&ty ma%& wmm the Cbepnt~.beitdong.@ A,
B, and C (&omJ o h n and d e n , 1984).
cate the strike of subsurface basement features: for example, the
positive anomaly under the Green River Basin, which may indicate a buried dome of basement rock south of the Wind River
Mountains; a northeast-striking anomaly in the Powder River
Basin, which is parallel to several lineaments postulated by Slack
(1981) to be ancestral basement highs associated with Precambrian shear zones; and a high in the southern Laramide Basin,
which may indicate an anorthosite or a layered mafic complex
beneath Mesozoic and Cenozoic sediments (Plate 2).
While palmmagnetic data are potentially important in determining the position of the Wyoming province relative to other
Precambrian cratonic elements of North America during the Archean and Proterozoic, much of the data are currently in disarray
(Roy, 1983). Large motions of North America relative to the
magnetic pole occurred in the Precambrian, but uncertainties in
dating of paleomagnetic samples make it ambiguous as to
whether the Superior, Slave, and Wyoming cratons have followed similar polar wander paths.
Pole positions for Precambrian rocks of the Wyoming province have been reported in Bergh (1970), Spall (1971), and
Larson and others (1973). The available data do not yet permit
construction of a meaingful polar wander path for the province.
The many new isotopic ages on Precambrian rocks of the Wyoming province provide excellent opportunities for new palmmagnetic studies, which may ultimately make it possible to
construct such a path.
I
110
I
,
I00
l
BELT MMI
n
~
~
TECTONIC MODELS
R. S. Houston, K. E. KarIstrom, E. A. Erslev, and
C. D. &ost
Knowledge of the tectonic history of the Wyoming province
is fragmentary, and evolutionary models are highly speculative,
especially for the older rocks. Areas such as the eastern Beartooth
Mountains where detailed geochemical studies have been made
are incompletely understood structurally, and areas such as the
Tobacco Root Mountains that have been carefully mapped have
received little geochemical attention. The effects and cause of
Proterozoic structural and thermal disturbances are poorly
known, although widespread resetting of K-Ar ages indicates
their importance.
The northern Wyoming province is dominated by domains
of gneiss formed under relatively high P-T metamorphic conditions and by an unusual diversity of Archean metasedimentary
suites. High-grade tonalitic gneiss complexes of Middle Archean
age have been documented in both ends of the Madison Range,
Beartooth Mountains, Bighorn Range, Wind River Range, Owl
Creek Mountains and Granite Mountains. Granulitefacies rocks
from the Tobacco Root Mountains, Ruby Range, Highland
Mountains, Black Tail Range, and Teton Range may also be
Middle Archean in age. The close correlation of U-Pb zircon ages
from these rocks in southwest Montana and the western Beartooth Mountains (P. Mueller, 1987, personal communication)are
~
Wyoming province
consistent with regional correlation into the Pre-Cherry Creek
Metamorphic Complex. The relationship of these rocks with
those to the east and southeast remains uncertain.
The fragmentary tectono-stratigraphic evidence for the Late
Archean evolution of northwestern Wyoming province supports
mobilistic tectonics similar to today's plate tectonics. Mueller and
others (1983,1984) have suggested that the batholithic rocks of
the eastern Beartooth Mountains represent the underpinning of an
Andean arc. Citing the diversity of lithostratigraphic and geochemical signatures, they proposed that the Beartooth Mountains
and southwest Montana represent a collection of Archean allochthonous terranes juxtaposed by eastward subduction of intervening
oceanic crust under the Beartooth arc. While this model explains
the general distribution of the lithologies, it is not compatible with
the lithostratigraphiccorrelations presented earlier in this chapter.
If the Cherry Creek Metamorphic Suite and the older, PreCherry Creek Metamorphic Complex are correlative from the
Beartooth Range to the Ruby Range, then this area must have
been a contiguous entity during Late Archean sedimentation.
Intrusions associated with the Beartooth arc cross-cut these rocks
in the western Beartooth Range, suggesting immediate proximity
of the arc. In this scenario, the Cherry Creek Metamorphic Suite
represents a stable shelf assemblage sloping eastward into a
deeper turbidite trough next to the Beartooth arc. The abundance
of metavolcanic strata throughout the Cherry Creek sequence,
whose stratigraphy is remarkably similar to the back-arc Coronation geocline of the Wopmay orogen (P. F. Hoffman, personal
communication to Erslev, 1987), is consistent with deposition in
a back-arc setting. In the western Beartooth Mountains, thrust
sheets of amphiboliteand quartzite overlain by andalusite-bearing
metaturbidites may represent the floor of the back-arc basin foreshortened on to the craton margin during the subsequent compression of the Beartooth orogeny. The moderate pressure
conditions of metamorphism during this event suggest a continental collision of southwest Montana with the central
Wyoming province.
Late Archean supracrustal rocks in central and southern
Wyoming may represent parts of back-arc basins, island arcs, or
microcontinents, and may even include fragments of continental
margin deposits. Possible back-arc basin sequences include those
in the Wind River Canyon in the central Owl Creek uplift, in the
Seminoe Mountains, and the greenstone belts in the southern and
central Laramie Range. The Phantom Lake Metamorphic Suite
of southern Wyoming, the Whalen Group of the Hartville uplift,
and the succession in the southern Wind River Mountains contain relatively clean quartzites and stromatolitic limestones that
may have been deposited on continental margins. Widespread
invasion by Late Archean granite, uncertainties as to the age and
geochemistry of the supracrustal sequence, limited exposure, and
failure of mappers to distinguish basement from cover in many
areas makes it difficult to develop a detailed accretionary model,
but all these sequences were added to the craton before about
2,700 Ma.
161
Late orogenic and post-orogenic thermal events affected
parts of the Wyoming province between 2,600 and 2,400 Ma.
Several sets of mafic dikes were emplaced during this interval,
and there was a thermal disturbance of the margins of the province between 1,950 and 1,700 Ma. Archean rocks along the
northwestern margin of the Wyoming province were thrust southeastward along ductile shear zones, resetting K-Ar ages and exposing deeper structural levels (Erslev and Sutter, 1990).
However, the continental crust in most of the Wyoming province
seems to have been thick and thermally stable by 2,700 to
2,600 Ma.
Early Proterozoic rocks have not been identified in the
northern part of the Wyoming province. They are present but
poorly understood along the western boundaries. Along the
southern and eastern boundaries, Proterozoic sequences are preserved, and enough information is available to discuss .the evolution of this margin.
Early Proterozoic metasedimentary rocks lie unconformably
on or are in fault contact with Late Archean basement and supracrustal rocks in the Sierra Madre, Medicine Bow Mountains,
Black Hills, and northern Uinta Mountains (Plate 2). Prior to
deposition of these strata the Wyoming province had been assembled and invaded by Late Archean granites. The stabilized
craton was eroded and weathered, and an extensive braided
drainage system developed. The basal Early Proterozoic quartzpebble conglomerates of southern Wyoming and the Black Hills
were deposited in this environment. We do not know how extensive these rivers and streams were because the deposits are preserved at the Wyoming province margins and have not been
recognized elsewhere. The Early Proterozoicfluvial deposits contain zircons as young as 2,400 Ma and are cut by 2,100 Ma
intrusions (Premo, 1983; Redden, 1980). Stratigraphic studies
both in the Black Hills (Redden, 1980) and southern Wyoming
(Karlstrom and others, 1983) suggest that these early fluvial
materials were deposited in response to rifting, which may have
continued after 2,100 Ma.
The Early Proterozoic rift margin was probably roughly
parallel to the present trace of the Cheyenne belt. The change in
trend from northeast-southwest in the Medicine Bows to eastwest in the Sierra Madre may reflect two legs of an early triple
junction. The size of the block rifted from the craton is unknown;
it may have been a large continent or only a small fragment. In
either case, the southern block exerted important controls on
early rift sedimentation. Paleocurrent and sedimentary data suggest that both the fluvial Deep Lake Group and deltaic lower
Libby Creek Group in the Medicine Bow Mountains were deposited in a northeast-trending basin bounded on the south by continental crust (Karlstrom and others, 1983). By the time of
deposition of the upper Libby Creek Group, the southern block
was gone and a carbonate shelf, open to the south, had developed. Extensive tholeiitic sills and dikes emplaced at 2,100 Ma in
both the Sierra Madre and Medicine Bow Mountains were probably related to rifting along the continental margin.
-
k,T
162
R S. H o w n and Others
The pattern of magnetic anomalies (Plate 2) and structural
trends-s
that the Early Prderozaic rocks in the Hack I-iiils
were deposited in a north-northwest-trending rift basin. The
Wymiag province may have been separated from the Superior
province by rifting in the Early Proterozoic. Basement slrmples
from drill holes east of the Black Hills show only Proterozoic agss
in a north-northwest-trending band some 200 km wide, indcating the presence of a wide Proterozoic mobile belt betwm the
Archean cratons. Thus, it is quite possible that Proterozoic metasedimentary s u m i o n s of the southern Wyoming province
developed indepmdently from those of the Superior province in
similar depo&tional settings on the margins of two widely separated Archerrn cratons. However, because of the remarkable
dmihity in sedimentary d o n s , &e great thkhem of these
miogeoclinal s u d o m , and apparent conbmporaneity of riftrelated tholeiitic intrwiom in the Wyoming and Superior povinces, we prefer a model that involves rifting of a Late Archem
supercontinent prior to closure of the rift-formed ocean basin to
form the intervening Trans-Hudson orcgen (Chapter 2).
Although we suggest that the Wyoming province was
welded to the Superior provke prior to Early R o t e m i c rifting, we do not suggest that the two provinces httd fhe same
history prior to their assembly in the Late Archean. Middle Archean racks are widely distributed in the northern and central
parts of the Wyoming province, whereas the most ancient rocks
of the Lake Superior area are in the south. The Wyoming province probably evolved in part by addition of daoconthents
along the southeastern margin, but the southern Superior province does not seem to have evolved in this manner. Om view is
that the Late Archean was the time of formation of a large
continental mass from assembly of smaller independent fragments. The Wyoming province does not appear to be a simple
western exWnsion of fhe Superior province interrupted by a
north-nortbwest~trendingmobile belt.
The interpretaton that the Wyoming and Superior provinces
were separated by Early Protermic lrifting of a Late Archean
supermntinent comes from two source& the similarity of miogeoclinal sedimentary successions of Early PMtaozoic age, which
are preserved in isolated remnants along the south a$r&of both
cratons (Houston, 1986); and the widespread episode of tousion
of basalt and alkaline rocks at about 2,100 Ma inbath provincm.
The intrusions have been inqreted as rift-related in the Superior
province (Anderson and Burke, 1983), the Black IIills (Wden,
1980), and southeastern Wyoming (aarlstrom and others, 1983).
Differences in present geology between the southern and eastern
margins of the Wyoming province are the result of iniW rift
differences and differences in timing and c o n f i i o n of plate
convergence following rifting. The southern margin is the result
of colliion of an Qpen continentaI margin shelf with Protermic
island araq Black Hills deformation resuited from closure of an
aulocogen at a high angle to the margin. Deformation of the
southm Superior province apparently involved collision of
several Arch- microcontinental blocks with the main mass of
the Superior province.
If this interpretatin is correct, the depositional sequences
along the southern r n of the
~ Sqmrbr sad Wyoming provinces should be cornlatable, and the depositi~nalhistories of the
Black
and southern Wyoming might be expected to mirror,
each ather.
Similarities in Ethology and stratigraphy suggest that the
Deep Lake and lower Libby Creek Grsupg of southern Wyoming
are m e a t i v e with the Huronian Supergroup of the Superior
province [KwMrom and others, 1%1; Young, 1970). The lower
Black Hills stramsion in the Nemo area (Redden, 1980) is about
the same age as the Huronian Supergroup and contains a fluvial
swcmsion that gmda upward into a marine s d o n . In both
the Black Hi& aand southern Wyoming, the younger mekwdimentary and m e t a v o l ~ csuacession resernbies the Marquette
Range Supergroup of the Lake Superim area. The upper Libby
Creek Group of southern WyomJag is k t correlated with the
lower part of the Marque& Range Supergroup; tke Nash Fork
Formation resembles the Kona Dolomite of the Marquette
Rage, and the French Slate resembles the Wewe Slate of the
Marquette Range Supergroup. The upper Libby Creek Group is
bracketed between 2.1 and 1.7 Ga; the ~q~
Range Super& ~ pd the Lake Superior area is bracketed between 2.4 and 1.9
Qa (Van Schus, 1976). It is thus possible that tbe Wyoming
rocks may be a Werent age than the Marquette Range SuperPtho1ogic resemblance between the Kona
group, but the sDolomite d h Mmquette Range and the dolomites of the Wash
Fork Formation, and the fact that similar units are not present
higher in the section in the Marquetb Range or elsewhere in the
Lake Superior area, support the correlation. The rocks above the
"Nemo Group" i~
the uppr Blwk Hilk succession are
l i t h o l o g i similar
~
to the Menominee and bwaga Groups-the
upper part of the Marque%& Range Supergroup. Both sequences
contain iron f d o n , volcanic rocks, slate, and graywacke
(Bayley and James, 1973). Two horizons wit& the Black Hi@
d o n have bieen dated as 1.89 and 1.97 Ga (Sims and othem,
this volume), and the Hemlock Formation of Baraga Group in
M a r q u e b a g e has been dated as 1.9 Ga (Van Schmus, 1976).
Thus it apthat reasonable IiU~ostratigmpbkcorrelations can be extended from m t h m Wyoming northeastward
along the ma- mgin of the Superior province, sugpting
the possibility of a 2,Wknz-long continental margin, the record
of which is preserved only in isolated outcrops of the miogeocline. These carrelations are discussed En more detail by Karlstrom and others f 1981).
ProWmoio collbional events have okured this early rift
margin, and diftkmnt areas have had quite merent tectonic bhories. Convergence of rift b a s h began in the Superior and Black
Hills areas.-bd arcs and previously rifted con;Einentdm e n &
were apparently dkling' with the Superior provhx at 1,900 to
1,800 Ma, at about the same time as ocean b a s h of unknown
width were closing in the Black Hills and northern areas of the
Trans-Hudson orogen. Collisio~lsof Proterozoic i s h d arcs with
the southern margin of the Wyoming province across the
Cheyenne belt took place at about 1,750 Ma, perhaps as a con-
,
i
j
1
Wyomingprovince
tinuation of a major Early Proterozoic assembly of continental
crustal blocks in North America, somewhat analogous to assembly of Phanerozoic supercontinentsbetween 450 and 250 Ma. In
this view, both the Late Archean and Early Proterozoic tectonic
history of North America can be thought of as dispersal and
re-assembly of supercontinents.
It is obvious that a critical need exists for additional geochronological, geological, and geochemical data if we are to
correctly interpret the geologic history of the Wyoming province.
Additional geophysical and geologic data from the subsurface are
needed to fill the extensive gaps between exposures of Precambrian rocks. It would be a serious mistake to assume that geologic
mapping of exposed Precambrian rocks in the province is adequate or complete. Many areas will benefit by adding geochronological and geochemical information without additional geologic
mapping, but large areas will require integrated geochemical,
petrologic, and geochronologic studies combined with careful
new geologic mapping to satisfactorily unravel the complex record of Precambrian events.
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MANUSCRIPT
ACCEFTED BY THE SOCIETYOCTOBER 7,1987
NOTES ADDED IN PROOF
Most of the material for this chapter was prepared in 1984 and revised in
1987. Inasmuch as descriptive geology was emphasized, major revisions are unnecessary. However, since 1984 geologic, geochemical, and geochronological
studies in the Wyoming province and adjacent areas have added significantly to
our understanding of the geologic history. The following is a brief summary of
some key contniutions since 1987.
The western, northern, and eastern margios of the Wyoming province have
been better defined The western margin can be extended to the Ruby Mountains
of Nevada (Lush and others, 1987) and the northern margin may coincide with a
northeast-trending set of geologic features, the Great Falls tectonic zone (O'Neil
and L o p , 1985; O'Neil and others, 1988), which has been traced from the Idaho
batholith through central Montana, and into Saskatchewan (Plate 1). New evidence, which is reviewed in our discussion of the Black Hills, supports the concept
that the Black Hills are part of the Wyoming province and that the eastern margin
of the Wyoming province is buried beneath Phanerozoic cover just east of this
uplift.
Recent geochronological and geochemical studies of the gneisses of the
Wyoming province demohate that the province contains some of the oldest
crust in North America and that this old crust extends farther south than we
anticipated in 1984. For example, there are gneisses in the central Wind River
Range of Wyoming that record U-Pb zircon evidence of at least two granulite
facies metamorphic events at ca. 3.2 and 2 7 Ga (Koesterer and others, 1987;
Aleinikoff and others, 1989).These gneisses also contain zircons, some of which
may be xenocrystic, with ages of ca. 3.8,3.65,3.35, and 2.85 Ga (Aleinikoff and
others, 1989). We do not know how far south the Early and Middle Archean
gneisses may extend, nor can we be certain that they represent fragments of
microcontinents welded together or remnants of a continuous basement that once
extended to the southern margin of the continent. Support for a Middle Archean
basement in the south comes from model ages of ca. 3.2 Ga determined on older
gneisses of the Granite Mountains (Fischer and Stacey, 1986) and the Hartville
Uplift of Wyoming (Snyder and Peteman, 1982).
.
169
Wyomingprovince
These geochronological studies coupled with the Early and Middle Premnbrim ages of gneiss and metasediment in southwest Montana and the Beartooth
Range indicate that the Wyoming province is significantly older than the Canadian Superior province and that it is not a western extension of the Superior
province.
We have considered the Black Hills of South Dakota part of the Wyoming
province on geologic evidence. Additional evidence for this suppition is from
geochemical and geochronological studies of granites of the Black Hills. Gosselin
and others (1988) note that the ages and geochemical evidence for long-lived
crustal sources for Black Hills Archean granites indicate a similarity to granites of
the Wyoming rather than the Superior province.
In one of the most recent tentative correlations of rock units of the Black
Hi- the quartzite-undivided and amphibolite and metagabbro units shown on
Figure 16 as parts of the Vanderlehr Formation are correlated with the Estes and
Roberts Draw Formations. Iron formation, schist, and quartzite along with the
eugeoclinal schists and phyllites (Fig. 16) are parts of the younger eugeoclinal
succession (Redden and French, 1989, p. B52-B53, Table B4).
The post-1984 literature on the Wyoming province includes a number of
papers that give more detailed information on Archean supercrustal rocks and
several summary papers that present plate tectonic models for genesis of the
Archean successions. In Wyoming, s u p e m t a l successions in the South Pass area
of the Wind River Range have been reviewed by Hull (1988) and by Hausel
(1991), the Casper Mountain supercrustal rocks by Gable and others (1988). the
northern Laramie Mountains by Snyder (1986), and the northern Medicine Bow
Mountains and Sierra Madre by Houston and others (1992). An interesting
collection of summary papers on Archean geology of southwest Montana and
northwest Wyoming is in Lewis and Berg (1988). These Montana articles present
geologic and geochemical evidence that plate tectonic processes may have operated during the assembly of this area. In the South Pass area of the southern Wind
River Range, Harper (1986) has d e s c n i a part of the s u p e m t a l succession as
dismembered ophiolites, and Hull (1988) has evaluated various plate tectonic
models that might apply to the rock units. Houston (1992) has suggested a plate
tectonic model for the origin of Late Archean supercrustal rocks of southern
Wyoming, but this model relies on questionable age relationships.
The Proterozoic of the Wyoming province has had its share of attention by
geologists, geochemists, and geochronologists. The Black H i were deformed
during formation of the Proterozoic Trans-Hudson orogen (Gosselin and others,
1988), and the effects of this orogeny may extend to the Hartville Uplift and
possibly as far west as the eastern Laramie Mountains of Wyoming (Houston,
1992). The Cheyenne belt of southern Wyoming has been interpreted as part of
the northern boundary of the Central Plains omen of Sims and Peterman (1986).
which is now known to have formed during a separate and later event than either
the Trans-Hudson orogen or the Penokean orogen of the Great Lakes area.
Houston (1992) has reviewed the Proterozoic geology of the Wyoming province
and suggested that certain critical geochronological problems need to be solved
before the relationship between the Cheyenne belt and the Trans-Hudson and
Penokean orogens can be evaluated fully.
More detailed geochronologic (Premo and Van Schmus, 1989) and structural studies (Duebendorfer, 1986) of the Cheyenne belt tend to support the
earlier hypothesis that this is an area where Proterozoic island arcs were attached
to a Proterozoic passive margin. In the Medicine Bow Mountains, structural
studies by Duebendorfer (1986), Duebendorfer and Houston (1986, 1987), and
by Houston (1992) have shown that the Cheyenne belt consists of three blocks
separated by zones of intense mylonitization. Ball and Farmer (1991) present Nd
isotope evidence that suggests that the northern block is Archean gneiss, the
middle block consists of metamorphosed intermixed sedimentary dehitus from
both Archean and Proterozoic sources, and the southern block is chiefly metam o r p h d Proterozoic volcanics and sediments. Study of granitoids that intrude
the Cheyenne belt and of metamorphic rocks and granite south of the Cheyenne
belt indicates that the Cheyenne belt proper is underlain by Archean crust, but
that no Archean crust is present to the south of the Cheyenne belt (Ball and
Farmer, 1991). These isotope studies of Ball and Farmer (1991) give us the first
evidence of Archean crust beneath the Cheyenne belt and verify the puzzling lack
of Archean south of the belt. If the Cheyenne belt simply represents an area where
Proterozoic island arcs collided with and overrode a Proterozoic passive margin,
as suggested by Karlstrom and Houston (1984), Archean crust beneath the Proterozoic passive margin must have extended some distance south. Its absence may
indicate a more complicated history for the Cheyenne belt. As suggested by
Houston (1992), the Cheyenne belt proper may preserve evidence of an early
collision and the island arc terrain may be exotic.
The Cheyenne belt extends southwest into the eastern Sierra Madre of
Wyoming, but in the west-central Sierra Madre the Cheyenne belt is severely
disrupted by a later cataclastic fault system (Duebendorfer and Houston, 1990).
This cataclastic fault system includes two major northwest-sbiking dextral strike
slip faults in the east-central Sierra Madre that appear to merge with east-west
thrust faults. The cataclastic faults are interpreted as part of a north-vergent,
thrust-tear system that cut across the Cheyenne belt as rocks of the upper plate
were transported north (Duebendorfer and Houston, 1990).
Bryant (1988) has suggested that the southern margin of the Wyoming
province in northern Utah is approximately midway between Salt Lake City and
Provo. He also suggests (Bryant, 1988, p. 48) that rocks of the Little Willow
Formation that are near the margin may be derived from Early Proterozoic
sediments deposited on the margin. However, there are no outcrops of the
Cheyenne belt west of the Sierra Madre (Plate 2). The location of the southern
Archean-Proterozoic boundary has been defined by neodymium isotope mapping
(Bennett and DePaolo, 1987). It extends west of the last outcrop through southern
Wyoming, northern Utah, and northern Nevada to the 8 7 ~ r / 8 6 ~.706
r line in
north-central Nevada (Plate l), where Archean and Early Proterozoic rocks are
absent (Bennett and DePaolo, 1987).
One of the most interesting results of the neodymium isotope studies is the
identification of Nd-model ages of 2.0 to 2.4 Ga south of the Wyoming province
boundary (Bennett and DePaolo, 1987). The absence of significant outcrops of
rocks exhibiting qstakation ages in the 2.0 to 2.4 Ga span within or south of the
Wyoming province led Bennett and DePaolo (1987) to rule out a new crustforming event. Instead they suggested that these Nd-model ages resulted from
mixing of Archean crustal material with newly formed Proterozoic arc material
near the Wyoming province boundary. The 2.0 to 2.4 Ga Nd-model ages are
present in a fairly broad area south of the Wyoming province in central and
western Wyoming, northern Utah, and northern Nevada (Bennett and DePaolo,
1987, Fig. 3). However, in the Sierra Madre and Medicine Bow Mountains, this
event has been identified in only one small area of the Cheyenne belt proper (Ball
and Farmer,l991); no crust of this age has been identified south of the Cheyenne
belt. Is this phenomenon due to the exotic nature of crust south of the Cheyenne
belt or, as suggested by Ball and Farmer (1991), is it because 2.0 to 2.4 Ga crust in
the vicinity of the C h e y e ~ ebelt was thrust over the craton and eroded?
Finally, we note again that our review of the Wyoming province emphasizes
geology. Fortunately, a recent review of the Archean geology of Wyoming (Frost
and Frost, 1992) emphasizes the geochemistry of the rocks of the province. We
hope that these two reviews will give the reader a good start in understanding this
interesting area.
-
Aleinikoff, J. N., Williams, I. S., Compston, W., Stuckless, J. S., and Worl, R. G.,
1989, Evidence for an Early Archean component in the Middle to Late
Archean g n e k of the Wind River Range, west-central Wyoming; Conventional and ion microprobe U-Pb data. Contributions to Mineralogy and
Petrology, v. 101, p. 198-206.
Ball, T. T., and Farmer, G. L., 1991, Identification of 2.0 and 2.4 Ga Nd model
age crustal material in the Cheyenne Belt, southeastern Wyoming; Implications for Proterozoic accretionary tectonics at the southern margin of the
Wyoming craton: Geology, v. 19, p. 360-363.
Benneq V. C., and DePaolo, D. J., 1987, Proterozoic crustal history of the
western United States as determined by Neodymium isotopic mapping: Geological Society of America Bulletin, v. 99, p. 674-685.
Bryant, B., 1988. Geology of the Fannington Canyon Complex, Wasatch Mountains, Utah: U.S. Geological Survey Professional Paper 1476,54 p.
Duebendorfer, E. M., 1986, Structure, metamorphism, and kinematic history of
the Cheyenne Belt, Medicine Bow Mountains, southeastern Wyoming
[PhD. thesis]: Laramie, University of Wyoming, 323 p.
170
R. S.Houston and Others
Duebendorfer, E. M., ead Houston, R S, 1990, Structural analysis of a d u d e
brittle Precambrian shear zone in the Siena Madre, Wyoyomiag; Western
extension of the Cheyenne Belt PrecamW Reseanoh, v. 48, p. 21-39.
Frost, 6. D., and Frost,B. R, 1992, The Archean history of the Wyoming
Province, fn Snoke, A. W., and Steidbmnn, J. R, eds., The geology of
WyoGeological Survey of Wyoming, Memoir 5 (in press).
Gosselin, D. C., Papike, J. J, ban,R. E., Petaman, Z. E., and Laul, J. C.,
1988, Archean rocks ofthe Black Hills, South Dakota; Reworked basement
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Harper, G. D., 1986, D i s m e m b d Archean ophiolite in southeast Wind River
Mountains, Wyoming-remains of Archan ocean
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Hausel, W. D., 1991, Economic gmlogy ofthe South Passgranitegreenstone belt,
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Hamton, R S., 1992, Proterozoic geology of Wyoming, in Snoke, A. W., and
Steidtmann, J. R, eds., The geology of Wyoming Geological Survey of
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Houston, R. S., Karlstrom, K. E,W,P. J., and Flurkey, A. J., 1992, New
stratigraphic subdivisions and M t i o n of subdivisions of Late Archean
and M y Prokromic metwdmentary and metavolcanic rocks ofthe Sierra
Madre and Medicine Bow Mountains, southern Wyoming. U.S. Geological
Survey Professioaal Paper 1520,56 p.
Hull, J. M., 1988, S.bucttlre and tectonic evolution of Archean supemubb,
southern Wind River Range, Wyoming [PhD. thesis) Rochester, N.Y.,
University of Rochester, 280 p.
Koesterer, M. B,and 5 others, 1987, Development of Archean crust in the
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Geology Special Publieation 96,195 p.
Lush, A. P., Snoke, A. W., and Wright, J. E., 1987, AUochthonous Archean
basement in the northern East Humboldt Range, Nevada: Geological Society
of America Abst;nlds with Prcpm, v. 19, p. 752.
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~
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and
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Premo$W. R., end Van Sch8lus,W. R., 1989, Zircon geachranology of Precambrian rocks in southmtem Wyoming and northern Colorado: Geological
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S i P. k,and Pctermao, Z. E., 1986, Early Proterozoic Qntral Plains orogen;
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the Hightowex SW 7.5 minute qwka@es (Bart A) and parts ofthe Fletcher
Park and Johnson Mountai~~
7.5 minute q u d n q k (Part B), Albany and
H~tteCounties, WyoU.S. Geological Survey Open File Report 86201, scale 124.000.
Printed in U.S.A.