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Geological Society of America Bulletin
Sedimentology, detrital zircon geochronology, and stable isotope
geochemistry of the lower Eocene strata in the Wind River Basin, central
Wyoming
Majie Fan, Peter G. DeCelles, George E. Gehrels, David L. Dettman, Jay Quade and S. Lynn Peyton
Geological Society of America Bulletin 2011;123;979-996
doi: 10.1130/B30235.1
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Notes
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Sedimentology, detrital zircon geochronology, and
stable isotope geochemistry of the lower Eocene strata
in the Wind River Basin, central Wyoming
Majie Fan†, Peter G. DeCelles, George E. Gehrels, David L. Dettman, Jay Quade, and S. Lynn Peyton
Department of Geosciences, University of Arizona, Tucson, Arizona 85721, USA
ABSTRACT
Shallow subduction of the Farallon plate
beneath the western United States has been
commonly accepted as the tectonic cause for
the Laramide deformation during Late Cretaceous through Eocene time. However, it remains unclear how shallow subduction would
produce the individual Laramide structures.
Critical information about the timing of individual Laramide uplifts, their paleoelevations at the time of uplift, and the temporal
relationships among Laramide uplifts have
yet to be documented at regional scale to address the question and evaluate competing
tectonic models. The Wind River Basin in
central Wyoming is filled with sedimentary
strata that record changes of paleogeography
and paleoelevation during Laramide deformation. We conducted a multidisciplinary
study of the sedimentology, detrital zircon
geochronology, and stable isotopic geochemistry of the lower Eocene Indian Meadows
and Wind River formations in the northwestern corner of the Wind River Basin in order
to improve understanding of the timing and
process of basin evolution, source terrane unroofing, and changes in paleoelevation and
paleoclimate.
Depositional environments changed from
alluvial fans during deposition of the Indian Meadows Formation to low-sinuosity
braided river systems during deposition of
the Wind River Formation. Paleocurrent
directions changed from southwestward to
mainly eastward through time. Conglomerate and sandstone compositions suggest that
the Washakie and/or western Owl Creek
ranges to the north of the basin experienced
rapid unroofing ca. 55.5–54.5 Ma, producing
a trend of predominantly Mesozoic clasts
giving way to Precambrian basement clasts
†
E-mail: mfan1@uwyo.edu
upsection. Rapid source terrane unroofing
is also suggested by the upsection changes
in detrital zircon U-Pb ages. Detrital zircons
in the upper Wind River Formation show
age distributions similar to those of modern
sands derived from the Wind River Range,
with up to ~20% of zircons derived from the
Archean basement rocks in the Wind River
Range, indicating that the range was largely
exhumed by ca. 53–51 Ma. The rise of these
ranges by 51 Ma formed a confined valley in
the northwestern part of the basin, and promoted development of a meandering fluvial
system in the center of the basin. The modern
paleodrainage configuration was essentially
established by early Eocene time.
Carbon isotope data from paleosols and
modern soil carbonate show that the soil CO2
respiration rate during the early Eocene was
higher than at present, from which a more
humid Eocene paleoclimate is inferred. Atmosphere pCO2 estimated from paleosol
carbon isotope values decreased from 2050 ±
450 ppmV to 900 ± 450 ppmV in the early
Eocene, consistent with results from previous studies. Oxygen isotope data from paleosol and fluvial cement carbonates show that
the paleoelevation of the Wind River Basin
was comparable to that of the modern Great
Plains (~500 m above sea level), and that local
relief between the Washakie and Wind River
ranges and the basin floor was 2.3 ± 0.8 km.
Up to 1 km of post-Laramide regional net uplift is required to form the present landscape
in central Wyoming.
INTRODUCTION
The eastern portion of the Cordilleran orogenic belt in the western interior USA consists of
the Laramide structural province, a region of Precambrian basement-involved uplifts, and intervening sedimentary basins that developed during
Late Cretaceous–Eocene time (Dickinson and
Snyder, 1978; Bird, 1988; DeCelles, 2004). The
deformation partitioned what had been a continental-scale foreland basin that developed east of
the Cordilleran thrust belt, ~1000–1500 km inland from the subduction zone. Modern regional
elevation is ~1.5 km with the range summits at
>4 km. Similar basement-involved uplifts in the
foreland of a major Cordilleran-style orogenic
system are present in the Sierras Pampeanas in
South America, where flat-slab subduction is intimately associated with the region of intraforeland basement deformation (Jordan et al., 1983;
Wagner et al., 2005). Although shallow subduction of the Farallon plate underneath North
America is the commonly accepted tectonic
cause for the Laramide deformation (e.g., Dickinson and Snyder, 1978; Bird, 1988; Saleeby,
2003; Sigloch et al., 2008; Liu et al., 2010), it
remains unclear how shallow subduction would
produce the individual Laramide structures and
the extent to which Laramide deformation may
be viewed as an eastward propagation of the
greater Cordilleran strain front (Erslev, 1993;
DeCelles, 2004). Critical information about
the timing of individual Laramide uplifts, their
paleoelevations at the time of uplift, and the temporal relationships among Laramide uplifts have
yet to be documented at regional scale to address
such questions.
Many previous structural, sedimentological,
thermochronological, and paleoelevation studies have shown that deformation and uplift in the
Laramide region is highly variable in time and
space (e.g., Love, 1939; Keefer, 1957; Soister,
1968; Seeland, 1978; Gries, 1983; Winterfeld
and Conard, 1983; Flemings, 1987; Dickinson
et al., 1988; DeCelles et al., 1991; Gregory and
Chase 1992; Cerveny and Steidtmann, 1993;
Crews and Ethridge, 1993; Omar et al., 1994;
Strecker, 1996; Hoy and Ridgway, 1997; Wolfe
et al., 1998; Dettman and Lohmann, 2000; Crowley et al., 2002; Fricke, 2003; Peyton and Reiners, 2007; Fan and Dettman, 2009). Tectonic
models explaining deformation during shallow
GSA Bulletin; May/June 2011; v. 123; no. 5/6; p. 979–996; doi: 10.1130/B30235.1; 12 figures; 2 tables; Data Repository item 2011013.
For permission to copy, contact editing@geosociety.org
© 2011 Geological Society of America
979
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Fan et al.
slab subduction include basal shear traction
(Bird, 1988), thrust and back thrust connected
to a master detachment in lower crust, which is
possibly linked to the Sevier thrust belt (Erslev,
1993), lateral injection by mid-crustal flow from
the overthickened Sevier orogenic hinterland
(McQuarrie and Chase, 2001), lithospheric buckling in response to horizontal endload (Tikoff
and Maxson, 2001), and isostatic rebound by
removing the ecologized Shatsky conjugate plateau on the Farallon plate (Liu et al., 2010). PostLaramide modification of the lithosphere in the
Laramide region may have played a vital role in
shaping the modern landscape. Mechanisms of
modification include: (1) thermal uplift due to
asthenospheric upwelling by removing the subducted slab or thickened mantle lithosphere beneath the western USA (Dickinson and Snyder,
1978; Humphreys, 1995; Sonder and Jones,
1999); (2) subcontinental-scale dynamic subsidence by induced asthenospheric counterflow
above the subducted slab (McMillan et al., 2002;
Heller et al., 2003; McMillan et al., 2006); (3) regional uplift caused by isostatic rebound of lithosphere due to climate-driven erosion (Pelletier,
2009); and/or (4) thermal upwelling associated
with the initiation of the Rio Grande Rift (Heller
et al., 2003; McMillan et al., 2006). Quantitative
data on paleoelevation, source-terrane unroofing
and exhumation, climate, and paleogeography
within a precise chronological context are required to evaluate these competing hypotheses.
In this paper we report results of a multidisciplinary study of the sedimentology, sandstone
petrography, detrital geochronology, and stable
isotope geochemistry of the early Eocene basin
fill in the Wind River Basin, central Wyoming.
From these data we reconstruct the paleogeography, paleoclimate, source-terrane exhumation,
tectonic setting, and basin evolution.
REGIONAL GEOLOGY
Stratigraphy and Age Control
The Wind River Basin in central Wyoming is
one of the many Laramide intermontane basins
formed to the east of the Sevier thrust belt. The
basin is surrounded by reverse fault-bounded,
basement-involved uplifts, including the Wind
River Range on the southwest, the Washakie
Range, Owl Creek Mountains, and Bighorn
Mountains on the north, the Casper arch on the
east, and the Granite Mountains on the south
(Fig. 1). Strata of Late Cretaceous–Eocene
age are exposed along the basin margin, and
have been studied extensively in the past century (Keefer, 1965; Soister, 1968; Courdin and
Hubert, 1969; Seeland, 1978; Phillips, 1983;
Winterfeld and Conard, 1983; Love and Chris-
980
Figure 1. (A) General map of the United States showing location of study area relative to the
major tectonic features and source terranes of detrital zircons (modified from Dickinson and
Gehrels, 2009). Black area stands for the Wind River Basin. (B) General map of Wyoming
showing the age of Precambrian basement surrounding the Wind River Basin (Frost et al.,
2000; Kirkwood, 2000). Gray lines represent the Wind River and its tributaries. Stars represent locations of modern river sand for detrital zircon U-Pb age analysis: (a) East Fork;
(b) Little Popo Agie River. (C) Geological map of northwestern Wind River Basin. Squares
represent locations of measured sections (modified from Love and Christiansen, 1985).
tiansen, 1985; Flemings, 1987). Among them,
only Courdin and Hubert (1969) and Flemings
(1987) presented modal petrographic analysis
of Late Cretaceous and Paleocene strata in the
western and southern basin. Flat-lying, lower
Eocene, Wind River Formation occupies the
center of the basin, with the greatest thickness along the east-northeastern basin margin
(Keefer, 1965). In the northwestern corner of
the basin, the slightly older Indian Meadows
Formation was tilted and displaced on the top
of the Wind River Formation by several northwest-striking thrust faults (Love, 1939; Keefer,
1957; Soister, 1968; Seeland, 1978; Winterfeld
and Conard, 1983; Flemings, 1987). A moderately angular unconformity between the Indian
Meadows Formation and underlying Mesozoic
rocks has been observed along the southwestern flank of the Washakie Range (Winterfeld
and Conard, 1983; this study). Our study focuses on the northwestern corner of the Wind
River Basin (Dubois–East Fork area, Fremont
County), where downcutting by the Wind River
exposed ~700 m of the Indian Meadows Formation and Wind River Formation along the north
bank, which is the thickest outcrop of the lower
Eocene strata in the basin.
The ages of the Wind River and Indian Meadows formations are constrained by North America land mammal ages (GSA Data Repository
Fig. DR11) (Love, 1939; Keefer, 1957; Winterfeld and Conard, 1983), whose boundaries in the
early Eocene have been revised based on correlation and calibration with radioisotope ages
and magnetostratigraphy in the Green River
Basin (Clyde et al., 1997; Robinson et al., 2004;
Smith et al., 2008). The presence of Canitius,
Hyopsodus, and Diacodexis places the Indian
Meadows Formation in the early Wasatchian
1
GSA Data Repository item 2011013, Figure
DR1: Chronostratigraphy of the early Eocene Wind
River Basin; Figure DR2: Photos of zircon grains;
Figure DR3: U-Pb concordia diagrams for detrital
zircon samples; Table DR1: Modal petrographic
point-counting parameters; Table DR2: Modal
petrographic data; Table DR3: Zircon U-Pb data;
Table DR4: Stable isotope data, is available at http://
www.geosociety.org/pubs/ft2011.htm or by request
to editing@geosociety.org.
(Wa1–Wa3, 55.5–54.5 Ma), and the presence of
Heptodon and Lambdotherium places the Wind
River Formation in the late Wasatchian (Wa6–
Wa7, ca. 53–51 Ma) (Clyde et al., 1997; Robinson
et al., 2004; Smith et al., 2008). The youngest
U-Pb ages of detrital zircons collected from
fluvial sandstones in this study cluster at 62.3 ±
1.3 Ma, which is older than the mammalian age
and thus provides no additional constraint on
depositional age. The contact between the Indian Meadows Formation and the Wind River
Formation is complex but appears conformable
in the East Fork area (Winterfeld and Conard,
1983; this study).
Tectonic Setting
The Wyoming craton includes granite
gneisses and minor amounts of metavolcanicmetasedimentary rocks older than 2.6–2.7 Ga
(Frost and Frost, 1993; Frost et al., 2000). Paleoproterozoic metavolcanic and metasedimentary rocks intruded by numerous plutons were
accreted onto the Wyoming craton along the
Cheyenne belt at ca. 1.86 Ga (Frost and Frost,
1993). Rifting of the western margin of Laurentia during Neoproterozoic–early Paleozoic time
accommodated the siliciclastic sedimentary
and metasedimentary rocks deposited under
predominantly marine environments (Levy and
Christie-Blick, 1991). During Paleozoic–early
Mesozoic time, the location of modern western
USA was in the Cordilleran miogeocline, and a
thick succession of carbonate rocks and subordinate siliciclastic rocks was deposited in a passive margin setting (Devlin and Bond, 1988).
From Late Jurassic on, the backarc crustal shortening and thickening caused by the subduction
of the Farallon plate produced the Cordillera
thrust belt, and a thick succession of marine
and marginal marine sediments was deposited
in Cordillera foreland when the Western Interior
Seaway inundation was terminated by the Late
Cretaceous time (e.g., DeCelles, 2004). During
ca. 80–40 Ma, Laramide deformation tectonically partitioned the foreland basin in Wyoming, Montana, and Colorado, and produced the
basement-involved uplifts and intervening
basins such as the Wind River Basin.
Geological Society of America Bulletin, May/June 2011
USA
(1.0
–
ille
mid-continent
(1.3–1.5 Ga)
Se
vie
r th
Yavapai-Mazatzal
(1.6–1.8 Ga)
Amarillo-Wichita
(ca. 0.525 Ga)
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USA
Mexico
he
ec )
p
am Ga
-C .58
n
ta 0
ca .54–
u
Y (0
East Mexico
magmatic arc
(0.23–0.29 Ga)
N
1000 km
120°W
20°N
Be
art
B
River
oo
th
M
Suwanee
(0.54–0.68 Ga)
80°W
20°N
C
C
s
Mtn )
orn
a
Bigh –2.9 G
(2.8
Absaroka
volcanic field
(51–46 Ma)
Washakie Range
MT
Bighorn
tns
40°N
Gre
nv
elt
tb
WY
r us
40°N
Superior (>2.5 Ga)
Canada
1 .3
Cordilleran
accretion
and
batholiths
(<0.25 Ga)
100°W
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Wyoming-HearneRae (>2.5 Ga)
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(0 pp
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Trans-Hudson
(1.8–1.9 Ga)
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Basin analysis of the early Eocene Wind River Basin
Washakie
Range
Owl Creek Mtns
(2.8–2.95 Ga)
a
W
in
d
Ri
ID
UT
Sevier thrust front
ve
r
b
Wind River Basin
Granite Mtns
(3.0–3.3 Ga)
Wind River
Range
(2.5–2.8 Ga)
2DB
Dubois
1DB
43.5°N
4DB
d
in
W
WY
3DB
er
iv
Uinta Mountains
R
CO
ge
Quarternary
Post-early Eocene deposites
Wind River Formation
India Meadows Formation
Mesozoic rocks
Paleozoic rocks
Precambrian basement
an
R
100 km
Geological Society of America Bulletin, May/June 2011
10 km
109°W
981
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Fan et al.
SEDIMENTOLOGY
Depositional environments are interpreted
on sedimentological analysis basis of four
measured sections, totaling ~750 m in thickness (Fig. 2). Paleocurrent directions are based
on 211 measurements of imbricated clasts
(10 per station) and trough cross-stratification
(method I of DeCelles et al., 1983). The Eocene
sedimentary rocks in the basin are described by
Love (1939), Keefer (1957), Seeland (1978),
Soister (1968), Winterfeld and Conard (1983),
and Flemings (1987), but lithofacies have not
been described in detail. Lithofacies that were
documented in the field are typical of fluvial deposits and are well understood in terms of depositional processes (summaries in Miall, 1978,
1996). Therefore, we provide a summary table
of the most common lithofacies and physical
processes in this study, and we focus on interpreting lithofacies assemblages and the corresponding depositional systems (Table 1).
Indian Meadows Formation:
Stream-Dominated Alluvial Fan Association
Description
The Indian Meadows Formation consists of
interlayered gray and yellow conglomerate,
sandstone, and brick-red siltstone with paleosols, totaling 315 m in thickness (Fig. 2B).
The Formation rests on an erosional unconformity on the upper Cretaceous Cody shale or
the Wind River Formation in our studied area
(Fig. 3A). The lower 100 m of the Formation
(11–110 m of section 3DB) is divided into
two 40- to 50-m-thick intervals of intercalated
pebble-cobble conglomerate (Gch, Gct, Gcm,
and Gcmi), medium-grained sandstone (Sh,
Sm, and St), siltstone (Fm), and paleosol (P),
sandwiching a 15 m interval of cobble-boulder
982
A
f
c
Sand
Basement volcanic
and metamorphic rock
Mesozoic
sandstone
Gravel
Pebble
Cobble
Boulder
vf m vc
Clay
Silt
Four schools of thought exist on the tectonic
history of the Wind River Basin during Laramide
deformation. Keefer (1965) suggested that the
basin experienced three stages of subsidence
in response to three uplift events in surrounding mountains, with the climax of deformation
happening to the Casper Arch, southern Bighorn
Mountains, and Owl Creek Mountains during
the early Eocene. In contrast, Flemings (1987)
concluded that the maximum rate of uplift and
deformation was during Maastrichtian time and
was caused by the thrusting of the Wind River
Range and Granite Mountains. Moreover, of
regional scale, Gries (1983) suggested that the
NW-SE–oriented uplifts were generally earlier than the E-W–oriented uplifts, and Brown
(1988) and Erslev (1993) argued progressive
SW to NE arch development.
Paleosol nodule
Chert
Mesozoic–Paleozoic
carbonate
Quartzite
Covered
Planar cross-stratification
Siltstone contains cobble
Stratigraphic correlation
Trough cross-stratification
Erosional contact
Reverse fault contact
Detrital zircon sample
Sandstone petrology sample
Normal grading
Clast imbrication and lamination
N
E
n
Paleocurrent direction and number (n) of measurements (imbricated
W clasts unless otherwise noted; tr—trough cross-stratification)
S
Figure 2 (on this and following page). (A) Stratigraphic symbols used
in the stratigraphic sections in (B) and (C). (B) Measured sections
of the Indian Meadows Formation in the northwestern Wind River
Basin. (C) Measured sections of the Wind River Formation in the
northwestern Wind River Basin. (B) and (C) include lithofacies,
paleocurrent directions, clast composition, and stratigraphic correlations (dotted lines). Section locations are marked in Figure 1C.
Lithofacies descriptions are summarized in Table 1. Abbreviations:
c—coarse; f—fine; m—medium; vc—very coarse; vf—very fine.
conglomerate (Gcm and Gmm) (Fig. 3B). Units
of lithofacies Gch, Gct, and Gcm (~5 m thick)
are lenticular, overlie basal scour surfaces, and
exhibit crude, upward-fining trends. In general, conglomerate grades rapidly upward into
sandstone. Beds of lithofacies Sm, Sh, and St
are composed of two or more, ~2- to 4-m-thick
sandstone packages separated from each other
by scour surfaces. Plane-parallel laminations of
a few millimeters thick are present only in the
upper ~0.5 m of many packages. The mudstone
and siltstone beds (Fm) usually contain paleosol carbonate nodules.
The middle 100 m of the Indian Meadows
Formation (7–110 m of 4DB and top of 2DB)
is dominated by pebble-cobble conglomerate
composed of lithofacies Gcm, Gcmi, Gct, and
Gcp. Two- to 4-m-thick beds are stacked to form
lenticular bodies up to 30 m thick, which extend
several kilometers laterally (Fig. 3C). These
kilometer-scale lithosomes have erosional bases
and generally fine upward. Siltstone lithofacies
Fm forms beds 1–2 m thick between these conglomeratic bodies; although these units are massive, they lack distinctive paleosol features.
The upper 110 m of the Indian Meadows
Formation (110–215 m of 4DB) is composed of
intercalated lenticular pebble-cobble conglomerate (Gct, Gcm, Gmm, Gcmi, and Gcp), mediumto coarse-grained sandstone (Sm, Sh, and St)
(Fig. 3D), siltstone (Fm), and paleosol (P). The
sandstones contain floating granules and pebbles
and generally cap the fining-upward conglomer-
ates. Siltstone of lithofacies Fm (~2 m) is sandy
and contains floating granules. Paleosol carbonate nodules are developed in most siltstone intervals, and organic-rich layers up to 1 m thick
are preserved above some of the carbonate accumulation horizons. In some instances, several
paleosols are stacked atop one another.
Interpretation
The assemblage of lithofacies in the Indian
Meadows Formation is characteristic of alluvial fan depositional systems (Nemec and Steel,
1984; DeCelles et al., 1991; Stanistreet and
McCarthy, 1993; Miall, 1996). The lower interval was deposited in the proximal alluvial fan by
gravel-bed braided channels and debris flows.
The thin bodies of conglomerate containing
sedimentary structures were deposited as longitudinal bar deposits in small flashy, highly unsteady braided streams (Hein and Walker, 1977;
Miall, 1978). The sandstones and siltstones were
deposited on the tops or flanks of bars during
waning flow or in slackwater ponds, such as
abandoned channels (Rust, 1972; Miall, 1978).
The unstratified, poorly sorted, matrix-supported
conglomerates represent viscous debris flows
triggered by catastrophic events (Nemec and
Steel, 1984; Schultz, 1984; DeCelles et al., 1991;
Stanistreet and McCarthy, 1993). The poorly
sorted, clast-supported conglomerates represent
high-concentration floods that were not in equilibrium with typical regime bedforms (Pierson,
1981; DeCelles et al., 1991).
Geological Society of America Bulletin, May/June 2011
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Basin analysis of the early Eocene Wind River Basin
2DB
4DB
210
B
300
Gcm
200
180
160
Gcmi, Gcm
Gcp
280
Sm,Fm
Sm
P
Fm,Sm
P
Gch,Gct
Gcp
260
Gcm,P
240
140
Gcm,Gct
Gcp
120
Gmm
Gch
Fm
Sm,P
Gcm,Fm,P
Gmm
22tr
100
C
10
Gcm,Fm
Gmh,Gct
P
Gmm,Fm
Gcm,Gch
P
Gmm,P
Gcm
Gmm
Gcm
P
P
Sm
P
220
1DB
10
120
Gch,Gchi
Gct,Gcp
200
100
Sm
Sm,Fm
Gct, Gch
P
Gmm,Sm
Gcm,Gct
Gcp
Gmm,Sm
180
80
80
Sm,P
Gch,Gct
Gcmi,Gcp
60
P
Gct, Gcmi
P
Gch,Gct
Gcm,Gcp
P
140
Gcmi,Sm
10
3DB
Gch,Gct
Gcmi,Gcp
120
Thickness (m)
105
Gcm,Gcmi
20
Gct,Gcm,Sh
Sm,Fm
10
Gcm,Sm
P
Gcm,Fm
Gcm
Gch,Gcm
St,Sh
Gch,Fm
80
100
0
Sh,Fm
Gmm,Fm,P
Gcm,Gchi,Gct
Gct
Sm,Fr,Fm,P
Gct,Gch
Sm
Gcm,Sm,Fm,P
80
60
0
c s sand g p c b
10 20
0
Gmm
Sm,Fr,P
P
Sm,Sh
Gcm,Fm
Gch,Gct
P
Fm
Gcmi,Sh
Sm,Fr
P
Sm,Sh,Fr
Gchi
40
20
0
10
St
Fm,P
Gch,sh
Sm,P
Gch,Gcm
P
St,Sr
Sm
Gct,St
9tr
P
St,Sm
Gcmi,Gct 10
P
10
Gct,Gcp
Gch
Fm,P
c s sand g p c b
10
Gcm,Gct,Fm,P
Sm,P
10
Gch,St
Fm
K
Shale
40
20tr
Gcm
40
20
P
Gct, Gmmi
c s sand g p c b
Fm
60
60
Gcm,Fm,P
Gch,Fh,Fm,P
Gmm,Sm,Fm,P 10
Gcp,Gcmi
Sh
100
Gct
Gch,Gct,Gcp
Fm
Gct,Gch,Gcp
40
160
Gcm,Sm
Fm,P
Gch,Gct,Sh,St
P
12tr
Gch,St
Sm,Sr,Sh
Gch,St
18tr
P
Sh,Fr,P
Fm, Sm
Sh,St
Gcp,Gch
P
10
10
St
Gcm,Gct
Gcmi,Sh,Sm
Gcmi
10
c s sand g p c b
Figure 2 (continued).
Geological Society of America Bulletin, May/June 2011
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Fan et al.
TABLE 1. LITHOFACIES AND INTERPRETATIONS USED IN THIS STUDY*
that the sediments were shed from the Washakie
Range and/or western Owl Creek Mountains by
transverse rivers (Figs. 1B and 1C).
Interpretation
The assemblage of lithofacies in the Wind
River Formation is characteristic of low-sinuosity, gravel-bed braided river depositional
systems (Hein and Walker, 1977; Miall, 1978,
1996). The conglomerates represent braided
channel deposits, and the sandstones were deposited on the tops or flanks of bars or in abandoned channels. The stacked lenticular thin beds
of horizontally stratified and cross-stratified
sandstone are interpreted as crevasse-splay deposits (Miall, 1978, 1996). The lack of floating
pebbles in siltstones suggests the lack of sandy
mudflows (Pierson, 1981). Thick paleosol carbonate horizons suggest that the floodplain of
the braided rivers was vegetated and the paleosols are well developed when the mean annual
precipitation was high, and the climate strongly
seasonal (Retallack, 2005).
The paleocurrent direction varies between
eastward and southwestward, but the most
prominent direction was eastward (Fig. 2C).
This suggests that the sediments were mainly
derived from the Wind River Range to southwest, with local input from the Washakie Range
and/or western Owl Creek Mountains to north
(Figs. 1B and 1C).
Overall, the depositional environment in the
northwestern Wind River Basin evolved from alluvial fan into low-sinuosity fluvial, with paleocurrent directions changing from southwestward
to mostly eastward during 55.5–51 Ma. This
evolution suggests that surface slope between the
upland sediment source terranes and the basin decreased through time (Stanistreet and McCarthy,
1993; Miall, 1996).
Wind River Formation:
Braided River Association
SANDSTONE PETROGRAPHY
AND PROVENANCE
Description
The Wind River Formation contains intercalated, lenticular, gray pebble-cobble conglomerates (Gcm, Gcmi, Gcp, Gct, and Gch) with
basal scour surfaces (Fig. 3E), sandstones (Sh,
Sr, St, and Sm), purple-red siltstones (Fr and
Fm), and stacked calcisols (P), totaling ~260 m in
thickness (Figs. 2C and 3F). The lithofacies are
in many ways identical to those of the upper Indian Meadows Formation discussed already, and
therefore only the differences will be described.
Floating pebbles do not occur in the sandstones
and siltstones. Many thin lenticular sandstone
beds (0.5–2 m thick) of lithofacies Sh and Sr are
encased in the siltstone intervals. Locally, two to
three sandstone beds are stacked on top of each
other, and display an upward-coarsening textural
trend. Calcisols are developed in most siltstone
intervals and are often stacked as three to four
paleosols of ~1 m thick (Fig. 3F).
Methods and Description
Lithofacies
code
Description
Interpretation
Massive, matrix-supported pebble to boulder
Deposition by cohesive mud-matrix debris
Gmm
conglomerate, poorly sorted, disorganized,
flows
unstratified
Pebble to boulder conglomerate, poorly sorted, clast- Deposition from sheet floods and clast-rich
Gcm
supported, unstratified
debris flows
Pebble to boulder conglomerate, moderately sorted, Deposition by traction currents in unsteady
Gcmi
clast-supported, stratified, imbricated
fluvial flows
Deposition by large gravelly ripples under
Pebble to cobble conglomerate, well sorted, clastGct
traction flows in relatively deep, stable
supported, trough cross-stratified
fluvial channels
Pebble to cobble conglomerate, well sorted, clastDeposition from shallow traction currents in
Gchi
supported, horizontally stratified, imbricated (long
longitudinal bars and gravel sheets
axis transverse to paleoflow)
Pebble to cobble conglomerate, well sorted, clastDeposition from shallow traction currents in
Gch
supported, horizontally stratified
longitudinal bars and gravel sheets
Deposition by large straight-crested gravelly
Pebble to cobble conglomerate, well sorted, clastGcp
ripples under traction flows in shallow
supported, planar cross-stratified
fluvial channels
Massive very coarse- to medium-grained sandstone,
Sm
Deposition from small gravity flow
sometimes contains gravel
Migration of small 2D and 3D ripples under
Fine- to medium-grained sandstone with small,
Sr
weak unidirectional flows in shallow
asymmetric ripples
channels
Migration of large 3D ripples (dunes) under
Medium- to coarse-grained sandstone with trough
moderately powerful unidirectional flows in
St
cross-stratification
large channels
Upper plane bed conditions, or flash floods
Fine- to medium-grained sandstone, horizontally
Sh
under unidirectional flows, either strong or
stratified
very shallow
Deposition from suspension or weak traction
Fr
Siltstone, horizontally laminated, or small ripples
current in overbank area
Deposition from suspension in ponds and
Fm
Massive siltstone, sometimes bioturbated
floodplain
P
Massive siltstone with carbonate nodules
Calcic paleosol
*Modified from Miall (1978) and DeCelles et al. (1991).
The thick conglomerate beds in the middle
interval of the Indian Meadows Formation are
typical of shallow gravel-bed braided river channels, with cross-stratification interpreted as bar
accretion sets, and the upward-fining trends reflecting a progressive decrease in flow strength
(Hein and Walker, 1977; Miall, 1996). The absence of debris flow facies and extensive lateral
distribution of this interval indicates the depositional environment was an extensive gravelly
braidplain (Rust and Gibling, 1990). The upper
interval was deposited in a stream-dominated alluvial fan environment, with the sandstone and
siltstone containing floating pebbles deposited in
sandy mudflows that spilled out of channels onto
the preexisting ground surface as sheet-flood deposits (e.g., Pierson, 1981; DeCelles et al., 1991).
The paleosols have distinct massive pedogenic
nodular carbonate intervals, and are classified as
calcisol (Mack et al., 1993). Well-developed calcisols are typical of floodplain deposits and represent relatively low sedimentation rates (Kraus,
1999). Overall, the depositional environment
changed from debris flows–dominated alluvial
fans into stream-dominated alluvial fans.
Paleocurrent data show consistently southwestward flow directions (Fig. 2B), suggesting
984
Standard petrographic thin sections were
made from 16 medium-grained sandstone samples that were collected while measuring stratigraphic sections. The thin sections were stained
for potassium and calcium feldspar and pointcounted (450 counts per slide) using a modified Gazzi-Dickinson method, in which crystals
larger than silt-sized within lithic fragments are
counted as monocrystalline grains (Ingersoll
et al., 1984). The point-counting parameters are
listed in Table DR1 (see footnote 1), and modal
data are given in Table DR2 (see footnote 1). The
modal sandstone analyses are augmented by 11
clast-counts (at least 100 clasts per station).
Monocrystalline quartz is the primary constituent of all samples. Feldspar mainly consists
of potassium feldspar and plagioclase, with a
K/P ratio that varies between 0.6 and 4.7. Lithic
Geological Society of America Bulletin, May/June 2011
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Basin analysis of the early Eocene Wind River Basin
A
Interpretation
B
SW
Indian Meadows Formation
20 m
30 mm
Wind River Formation
C
D
2 mm
E
F
30 m
1m
Figure 3. Photographs of northwestern Wind River Basin outcrops. (A) Angular unconformity
between the Wind River Formation (below, horizontal) and the Indian Meadows Formation (moderately dipping to northeast). (B) Disorganized conglomerate of lithofacies Gcm in
the lower Indian Meadows Formation. (C) Imbricated conglomerate of lithofacies Gcmi in the
middle Indian Meadows Formation. (D) Coarse sandstone containing floating pebbles (Sm) in
the upper Indian Meadows Formation. (E) Cross-stratified conglomerate of lithofacies Gct
in the Wind River Formation. (F) Overview of the Wind River Formation.
grain populations are dominated by micritic
carbonate. Volcanic grains constitute <5% of
total lithic grains. Other lithic grains include
phyllite, siltstone, and quartzite derived from
sedimentary and metasedimentary strata. Biotite, muscovite, pyroxene, olivine, glauconite,
and chlorite are the most common accessory
minerals. The Indian Meadows and Wind River
formations sandstones have average modal
compositions of Qm:F:Lt = 57:10:33, Qt:F:L =
77:11:13; and Qm:F:Lt = 41:26:33, Qt:F:L =
56:26:18, respectively. The Indian Meadows
Formation contains greater amounts of quartz,
and lesser amounts of feldspar than the Wind
River Formation (Fig. 4). All of the Eocene
sandstones that we studied are dominated by
calcite cements (Fig. 5).
Clasts of dolostone, limestone, and sandstone dominate the conglomerates in the Indian
Meadows Formation. Carbonate clasts with
brachiopod fossils are abundant in the lower
interval of the Indian Meadows Formation. The
sandstone clasts are quartzose and have distinct
red color. Granite gneiss clasts first appear in the
middle interval of the Indian Meadows Formation, and the content increases upsection, reaching ~20% at the top of the Indian Meadows
Formation. Granite (~45%) and carbonate
(~40%) clasts dominate the conglomerates in
the Wind River Formation.
The sandstone modal petrographic data show
that the lower Eocene sandstones in the northwestern Wind River Basin were derived from a
recycled orogenic provenance (Fig. 4). Potential source terranes in Wyoming include: Cretaceous sandstone, siltstone, and chert-pebble
conglomerate; Jurassic and Triassic eolian
sandstone and marine limestone; Jurassic–
upper Permian glauconitic siltstone, sandstone,
and limestone; Permian–upper Cambrian limestone, dolostone, and sandstone; middle-upper
Cambrian micaceous shale; Cambrian coarsegrained quartzite; and Archean granite gneiss
(Love and Christiansen, 1985).
Lithic grains of carbonate, sandstone, mudstone, and phyllite were probably derived
from the Phanerozoic sedimentary cover strata
exposed along the northern and southwestern flanks of the Wind River Basin. The large
amount of well-rounded monocrystalline quartz
grains in the Indian Meadows Formation suggests the sediment was recycled from underlying quartzose sedimentary units, mostly from
Jurassic and Triassic eolianite. The occurrence
of glauconite suggests that nearby Mesozoic sedimentary rocks were major sediment
sources. The abundant feldspar grains in the
Wind River Formation were derived from Precambrian basement rocks.
The red sandstone clasts in the conglomerates
are of recycled Jurassic and Triassic eolian and
fluvial sandstones, and the limestone and dolostone clasts were derived from the Phanerozoic
marine carbonate units. The carbonate clasts
with brachiopod fossils are of Mississippian
and Pennsylvanian age (Love and Christiansen,
1985). The granite clasts were derived from the
Precambrian basement rocks, which are now exposed in the Laramide ranges.
Overall, the Eocene sandstones and conglomerates of the northwestern Wind River Basin
exhibit a simple unroofing sequence, with initial source terranes dominated by Mesozoic
and upper Paleozoic rocks giving way to lower
Paleozoic and Precambrian basement rocks
upsection.
DETRITAL ZIRCON U-Pb
GEOCHRONOLOGY
Samples and Methods
Six samples of medium- to coarse-grained
sandstone from the Indian Meadows Formation
(3DB69, 4DB4.5, 4DB191, and 2DB273) and
Wind River Formation (2DB2 and 2DB257),
one sample of a granite cobble (2DB2cobble),
and two samples of modern river sands (East
Geological Society of America Bulletin, May/June 2011
985
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Fan et al.
Qm
zoic age constitute the largest population on
age-probability diagrams. Sample 2DB2cobble
has a mean U-Pb zircon crystallization age of
2.61 Ga (Fig. 7). A single zircon grain in sample
LP produced an age of 47.7 ± 2.1 Ma, but the
other age groups are similar to those in the Eocene samples. The nine detrital zircon grains
with concordant ages in sample EF cluster at
2.7–3.2 Ga. The interpretation of detrital zircon
ages in the western USA is relatively well understood (Armstrong and Ward, 1993; Gehrels
et al., 1995; Roback and Walker, 1995; Stewart
et al., 2001; Torres et al., 1999; Gehrels, 2000;
Gehrels et al., 2000; DeGraaff-Surpless et al.,
2002; Dickinson and Gehrels, 2003; Link et al.,
2005; Dickinson and Gehrels, 2009). Here, we
summarize the sources of specific age groups in
Table 2. The major source terranes are presented
in Figure 1A.
Qt
CB
CB
RO
RO
MA
MA
F
Lt
Qp
Lv
F
L
mean
Wind River Formation n = 8
mean Qm
Indian Meadows Formation n = 6
Ls
P
K
Figure 4. Ternary diagrams showing modal-framework grain
compositions of sandstones from lower Eocene strata of the northwestern Wind River Basin. Provenance fields after Dickinson and
Suczek (1979). RO—recycled orogen; CB—continental block;
MA—magmatic arc. Framework components are explained in GSA
Data Repository Table DR1 (see footnote 1), and data are listed in
Table DR2 (see footnote 1).
Fork [EF], Little Popo Agie River [LP]) were
collected and processed by standard methods
for separating zircons (Gehrels et al., 2000).
The Little Popo Agie River and East Fork rise in
Precambrian crystalline basement of the Wind
River Range and Washakie Range, respectively
(Fig. 1B). The former cuts an ~2-km-thick Paleozoic and Mesozoic sedimentary section on the
northeastern flank of the Wind River Range,
whereas no Paleozoic and Mesozoic sedimentary strata are exposed along the East Fork.
Zircon grains in sample EF are euhedral, and
~200 μm long, whereas those in sample LP are
similar to those from Eocene sandstones, subhedral to anhedral, and mostly 50–100 μm in
diameter (Fig. DR2 [see footnote 1]).
About 100 individual zircon grains were
analyzed for the early Eocene sandstone
samples and sample LP. All 19 zircons extracted from sample EF and 23 zircons from
sample 2DB2cobble were analyzed. U-Pb geochronology of zircons was conducted by laserablation–multi-collector–inductively coupled
plasma–mass spectrometry (LA-MC-ICP-MS)
at the University of Arizona LaserChron Center. Details of analytical procedures and data
processes are described in Gehrels et al. (2008).
The ages presented are 206Pb/238U ages for grains
<1000 Ma and 206Pb/207Pb ages for grains with
206
Pb/238U ages >1000 Ma (age-probability plots
in Fig. 6; concordia diagrams in Fig. DR3 [see
footnote 1]). Analyses that yielded isotopic data
of acceptable discordance, in-run fractionation,
986
Interpretation
and precision are shown in Table DR3 (see
footnote 1). Nine of the 19 zircon grains from
sample EF show concordant ages. There is no
systematic correlation between shape, color, angularity, and U-Pb age for zircons in the Eocene
sandstones and sample LP.
Results
The youngest detrital zircons recovered from
the Eocene basin fill are in the 60 to 80 Ma
range, which is older than the depositional age
derived from North American land mammal
ages. Ages of pre–80 Ma zircons generally fall
within the following intervals: 80–220 Ma, 230–
290 Ma, 330–760 Ma, 1.0–1.3 Ga, 1.3–1.5 Ga,
1.6–1.8 Ga, and 2.5–3.0 Ga. Zircons of Protero-
A
Plagioclase
The euhedral morphology of zircon grains
suggests sample EF was derived from weathered granite in the Washakie Range and/or the
Owl Creek Mountains. The zircon U-Pb ages
are consistent with the crystallization age of the
granodiorite gneiss in west Owl Creek Mountains
(ca. 2.8 Ga) (Kirkwood, 2000). The youngest zircon in sample LP is ca. 48 Ma; this grain was
most likely derived from the Absaroka volcanic
field (Armstrong and Ward, 1993). Approximately 20% of the detrital zircons analyzed in
sample LP are in the 2.5 to 2.7 Ga age range,
which is consistent with zircon U-Pb ages of
the granitic plutons in the Wind River Range
(2.55–2.67 Ga) (Frost et al., 2000). The anhedral
morphology and small grain size suggest most
of the zircon grains in sample LP were recycled,
presumably from Mesozoic sandstones. This is
supported by the fact that Mesozoic sandstones
in western USA have similar age spectra, including: (1) zircons from the Cordilleran magmatic
arc in California and Nevada (80–220 Ma) and
B
Calcite
Glauconite
Weathered K-feldspar
250 μm
Plagioclase
250 μm
Figure 5. Photographs of petrographic thin sections of sandstone from the Wind River Formation (in cross-polarized light).
Geological Society of America Bulletin, May/June 2011
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Basin analysis of the early Eocene Wind River Basin
2.61 Ga
East Fork
n=9
2400
2600
2800
3000
3200
3400
Little Popo Agie River
n = 113
Wind River Formation
2DB257
n = 108
Normalized age probability
2DB2
n = 88
Indian Meadows Formation
4DB191
n = 97
2DB273
n = 89
4DB4.5
n = 94
3DB69
n = 114
0
500
1000
1500
2000
2500
3000
3500
Age (Ma)
Figure 6. U-Pb age-probability diagrams of detrital zircons for samples in this study. The curves represent a sum of the probability distributions of all analyses from a sample, normalized so that the areas beneath the probability curves are equal for all samples. Thick dashed line
represents 2.61 Ga, which is marked as the reference of Archean zircons from the Wind River Range (see Fig. 7). Sources of different zircon
age peaks in this study are summarized in Table 2.
Geological Society of America Bulletin, May/June 2011
987
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Fan et al.
0.56
0.54
2700
Figure 7. U/Pb concordia diagrams for zircons from a granite cobble in the base of section
2DB. The age fits the crystallization age of the granite in the
Wind River Range. MSWD—
mean square of weighted deviates.
206Pb/ 238U
0.52
2600
0.50
2500
0.48
0.46
Mean age = 2606 ± 24 Ma
MSWD = 4.4
n = 23, 68.3% confidence
0.44
0.42
10
11
12
13
207Pb/ 235U
Permian–Triassic magmatic arc sources in eastern Mexico (230–280 Ma); (2) recycled zircons
from the Devonian–Mississippian Antler orogenic belt in Nevada and Idaho (330–450 Ma);
(3) zircons derived from the Appalachian orogenic belt (360–760 Ma), and Grenville terrane
(1.0–1.3 Ga) in the eastern USA; and (4) recycled zircons from the Yavapai-Mazatzal orogen
(1.6–1.8 Ga) (Roback and Walker, 1995; Dickinson and Gehrels, 2003, 2009).
Detrital zircon geochronology of the modern
river sands sheds light on the Eocene detrital
zircon age spectra in the following aspects: (1) zircon grains recycled from Mesozoic sandstones
with mostly Grenville and Yavapai-Mazatzal
ages, dominate the zircon population in fluvial
basinal sediments; (2) basement exposures
along the modern basin margin contribute only
~20% of the total zircon grains to basinal sediment; and (3) zircons from the Washakie Range
(2.7–3.1 Ga) are older than zircons from the
Wind River Range (ca. 2.6 Ga).
The evolution of the detrital zircon age spectra in the Indian Meadows Formation shows a
rapid source terrane unroofing in 55.4–54.5 Ma
(Fig. 6). The detrital age spectrum of the lower
part of the formation (3DB69 and 4DB4.5) is
similar to that of the Jurassic eolianite in Utah
and Wyoming, with abundant Grenville (ca. 1.0–
1.3 Ga) and lesser amounts of Yavapai-Mazatzal
(1.6–1.8 Ga) age grains (Dickinson and Gehrels,
2003, 2009). The proportion of grains derived
from the Cordilleran magmatic arc increases in
the middle formation (sample 2DB273). This
change corresponds to the inferred increase of
fluvial discharge in the alluvial fan association.
The age spectrum of the upper formation (sample 4DB191) is dominated by grains of YavapaiMazatzal age, reflecting a recycled source in
Paleozoic strata (Gehrels et al., 1995; Stewart
et al., 2001) when the Washakie and/or western
Owl Creek ranges were unroofed.
The detrital zircon age spectra in the Wind
River Formation show sediment shed from the
Wind River Range dominated the basinal sediment in 53–51 Ma. The ca. 2.61 Ga zircon age
from the granite cobble from the lower Wind
River Formation matches the age of basement
granite in the Wind River Range (Frost et al.,
2000). Abundant granite clasts and Archean
zircons (ca. 2.6 Ga) in the lower Wind River Formation (sample 2DB2) suggest that Wind River
Range basement rocks were a major source by
early Eocene time. The detrital zircon age spectrum changes into the shape of modern sand
sample LP at the top of the Wind River Formation (sample 2DB257), with predominant
Yavapai-Mazatzal ages, a large component of
Mesozoic ages indicating derivation from the
Cordillera magmatic arc, and ~20% Archean
ages with sources in the Wind River Range
basement rocks.
In summary, detrital zircon geochronology
of lower Eocene sandstones in the Wind River
Basin shows: (1) the source terrane of the Indian
Meadows Formation (ca. 55.5–54.5 Ma) was
most likely the Washakie Range and/or Owl
Creek Mountains, consistent with southwestward
paleoflow directions. Sediment derived from
these sources records an unroofing sequence
from upper Mesozoic sandstone to lower Paleozoic strata; (2) grains derived from the Wind
River Range began to enter the depositional system by ca. 53–51 Ma; and (3) the basement of
the Wind River Range was eroded and exposed
to its present extent by 51 Ma.
STABLE ISOTOPE GEOCHEMISTRY
Methods
TABLE 2. MAJOR U-Pb AGE POPULATIONS OF DETRITAL ZIRCONS IN MODERN
RIVER SAND AND EARLY EOCENE SEDIMENT IN THE WIND RIVER BASIN
Minimum Maximum
age
age
Population
(Ma)
(Ma)
Source regions
Challis and Absaroka
South-central Idaho, northwest Wyoming, southeast Montana
42
60
Volcanic Supergroup
(Armstrong and Ward, 1993)
Cordilleran magmatic arc
Central and southern California, Nevada, south-central Idaho
(mainly Sierra Nevada
80
220
(Armstrong and Ward, 1993)
and Idaho batholiths)
Permian–Triassic
230
290
Eastern Mexico (Torres et al., 1999)
magmatic arc
Southern and eastern United States; recycled from
Appalachian orogenic belt
330
760
Mesozoic to Paleozoic miogeocline (Dickinson and
Gehrels, 2003, 2009)
Recycled from Neoproterozoic to Paleozoic miogeocline
Grenville orogeny
950
1300
(Dickinson and Gehrels, 2003, 2009)
Recycled from Cordilleran miogeocline (Dickinson and
Midcontinent magmatism
1304
1532
Gehrels, 2003, 2009)
Recycled from Cordilleran miogeocline (Dickinson and
Yavapai-Mazatzal province
1600
1800
Gehrels, 2003, 2009)
Mainly from exposed Archean craton; also recycled through
Cordilleran miogeocline (Frost et al., 2000; Kirkwood,
Archean Wyoming craton
2500
3200
2000; Dickinson and Gehrels, 2003; Frost and Fanning,
2006; Dickinson and Gehrels, 2009)
988
The basic approach of this study and the analytical methods for carbonate isotope analysis
follow that of DeCelles et al. (2007). Modern
soil carbonate formed under cobble clasts from
at least 50 cm depth in active soil horizons was
collected. Paleosol carbonate nodules were collected from the middle of paleosol beds at least
50 cm below the paleosurface (Fig. 8). We also
analyzed the δ13C values of eight dominant
plant species in the study area collected during
the growing season. Organic matter δ13C was
measured using a Costech EA combustion unit
connected to a Finnigan Delta Plus XL mass
spectrometer via a Conflo III split interface. The
precision is better than ±0.1% (1σ). All the isotope values of carbonate and organic matter are
reported relative to PeeDee belemnite (PDB),
and that of water relative to standard mean
ocean water (VSMOW).
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Basin analysis of the early Eocene Wind River Basin
A
B
Organic-rich A horizon
2 mm
Evaluation of Diagenesis
Paleosol carbonate
200 mm
Figure 8. Field photographs and images of thin sections of the carbonate nodules in the early
Eocene Wind River Basin: (A) paleosol in the Wind River Formation; (B) micritic paleosol
carbonate nodule with microsparite.
0
Evaporation
–2
Marine
–4
18
δ O (PDB)
–6
–8
–10
Paleosols
Secondary vein
Marine limestone and coral fossil
Modern soil
Fluvial cement
Corrected paleosols
–12
–14
–16
–18
–15
to –0.3‰. The δ13C values of secondary veins
and fluvial cements vary between –9.3‰ and
–5.8‰ and –7.0‰ and –3.2‰, respectively.
The δ13C values of eight species of modern
plants are –24.5‰ ± 1.0‰, consistent with the
values obtained from Holocene soil organic
matter by Amundson et al. (1996).
Low soil respiration rate
–10
–5
0
5
10
13
δ C (PDB)
Figure 9. Results of the stable isotope analyses in this study. See
text for an explanation of the correction of paleosol isotope values.
PDB—PeeDee belemnite.
Results
The δ18O values of all marine and nonmarine
carbonate in the Wind River Basin range from
–17.6% to –1.5‰ (Table DR4 [see footnote
1]; Fig. 9). Recycled marine limestone clasts
and coral fossils have the highest δ18O values,
–4.9‰ to –1.5‰, whereas early Eocene fluvial
cements have the lowest values, from –17.6‰
to –11.5‰. The δ18O values of paleosols and
isotope values of fluvial cements do not show
systematic change through the section (Fig. 10).
Early Eocene paleosol micrites have δ18O values between –9.8‰ and –8.5‰, display little
variability, and are similar to the δ18O values of
early Eocene paleosol nodules in the Bighorn
Basin (Koch et al., 1995). The δ18O values of
secondary veins in the paleosol nodules are
generally lower than those of paleosol nodules,
ranging from –14.1‰ to –9.4‰. The δ18O values of modern soil carbonates vary between
–10.3‰ and –6.3‰.
The δ13C values of all marine and nonmarine
carbonate in the Wind River Basin range from
–9.5‰ to +4.4‰ (Table DR4 [see footnote
1]; Fig. 9). Recycled marine limestone clasts
and coral fossils have the highest δ13C values,
of –4.9‰ to –1.5‰. Early Eocene paleosols
have δ13C values between –9.5‰ and –5.6‰,
with the highest values from the middle and
lower Indian Meadows Formation (Fig. 10).
These paleosol δ13C values are similar to the
values of early Eocene paleosols in the Bighorn Basin (Koch et al., 1995). The δ13C values of modern soil carbonates are generally
higher than paleosols, and range from –4.5‰
Several lines of evidence argue that early
Eocene paleosol carbonate nodules and fluvial
carbonate cements in the Wind River Basin are
not diagenetically reset, and therefore that the
isotopic results are primary and thus can be used
to reconstruct paleoelevation and paleoclimate.
First, carbonate petrography study shows the
paleosol carbonates are dominantly micritic
(Fig. 8). Recrystallization or cement infilling
is not observed in thin sections. Calcite cement
constitutes up to ~50% of the sandstone, consistent with the high initial porosity of weakly
compacted sands (Fig. 5). Such high percentages characterize early diagenetic calcite, since
early cementation occurs prior to significant
compaction (Quade and Roe, 1999, and references therein; Sanyal et al., 2005). Second,
the δ18O values of a few sparry calcite veins in
paleosol nodules are lower than the values of
micrite in the paleosol nodules (Fig. 9). This
suggests that the calcite veins probably formed
by fluids at temperatures higher than at the Eocene soil surface. If paleosol nodule micrite
underwent diagenetic resetting during the formation of secondary sparry veins, the isotopic
values of micrite would be similar to the values
of secondary veins. Third, the δ18O and δ13C
values of recycled marine limestone clasts and
fossil corals in conglomerates immediately adjacent to paleosols and sandstones are similar
to those from unaltered Mesozoic–Paleozoic
marine carbonate (Veizer et al., 1999) but significantly higher than the values from paleosol
and fluvial carbonate cement. The δ18O and δ13C
values of soil nodules, secondary veins, fluvial
cement, and marine limestone are distinct from
each other, and form clusters in the δ18O versus
δ13C diagram (Fig. 9). Diagenesis would tend to
homogenize the isotopic values of marine limestones, micrites, and secondary spars. The large
variation of fluvial cement δ18O values may be
due to contamination during sampling by marine limestone detritus, which have relatively
positive values (Figs. 9 and 10).
Oxygen Isotopes and Paleoaltimetry
The δ18O value of ancient local meteoric
water can be calculated from paleosol δ18O
values and constrained temperature. Pedogenic
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989
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Fan et al.
Wind River Formation
(53–51 Ma)
600
m
500
400
?
Indian Meadows Formation
(55.5–54.5 Ma)
300
200
100
Feldspar grains
Granite clasts
Archean zircons
0
sand
0
10
20
30
40
Percentage (%)
Paleosol
Fluvial cement
Paleosol
Fluvial cement
50 –20
–15
–10
–5 –10
–8
–6
–4
–2
δ13C (PDB)
δ18O (PDB)
Figure 10. Simplified stratigraphy, carbon and oxygen isotope values of paleosol and fluvial cement, percentage
of granite clasts, Archean zircons, and feldspar through section. Vertically ruled areas are depositional lacunae.
PDB—PeeDee belemnite.
calcite is thought to form in isotopic equilibrium with soil water during the warm season
(Quade et al., 1989; Cerling and Quade, 1993;
Quade et al., 2007; Breeker et al., 2009). The
temperature of the warm season is usually between the mean annual temperature and maximum summer temperature. Soil water derives
from rainfall that may undergo evaporative
enrichment in δ18O, especially in arid settings.
We assume the summer temperature in Wyoming in early Eocene was ~30 °C based on an
early middle Eocene temperature reconstruction in France, which has a similar mean annual temperature as Wyoming (Andreasson
and Schmitz, 1996). The δ18O values of early
Eocene precipitation can then be estimated by
using soil nodule formation temperatures of
18.6–30 °C and standard calcite fractionation
factors (Kim and O’Neil, 1997). The calculated
δ18O values of early Eocene precipitation in the
Bighorn and Wind River basins are –6.8‰ ±
1.9‰ (VSMOW), which are consistent with the
values derived from the δ18O values of mammal
tooth enamel in the early Eocene Bighorn Basin
and Powder River Basin (Fricke, 2003) and the
values derived from freshwater bivalve fossils
in the lower Eocene in the Powder River Basin
(Fan and Dettman, 2009). For comparison with
modern water, 1‰ is added to the δ18O value
of early Eocene precipitation to account for the
990
lower seawater δ18O value in the early Eocene
(Zachos et al., 1994).
The δ18O value of groundwater can be calculated from fluvial cement δ18O values and constrained temperature. Early diagenetic cements
in fluvial sandstone formed in equilibrium with
shallow groundwater (Quade and Roe, 1999;
Mack et al., 2000). The deposition temperature
of the early cements should be close to mean
annual temperature since it formed near the surface (18.6 ± 2.4 °C). The calculated δ18O value
of groundwater is –16.6‰ ± 0.5‰ (VSMOW),
using the most negative δ18O value of fluvial
cement (–17.6‰, PDB). This value (–16.6‰)
is comparable with the unweighted mean δ18O
value of modern precipitation in the south flank
of the Wind River Range (–15.2‰, Pinedale,
Wyoming, elevation of 2.4 km) after adding 1‰
for the lower early Eocene seawater δ18O value
(Zachos et al., 1994) (http://www.uaa.alaska.edu/
enri/usnip). The large difference between the
δ18O values of groundwater and local precipitation in the early Eocene Wind River Basin
suggests the groundwater was not sourced in local basin precipitation, but rather from the surrounding highlands.
At present, rainfall in the high elevations of
the Rocky Mountains is dominated by winter precipitation from the Pacific Ocean and
the Arctic Ocean. The low elevation regions
in the Laramide province likely received abundant summer precipitation from the Gulf of Mexico and small amounts of winter precipitation
from the Arctic Ocean (Bryson and Hare, 1974;
Dutton et al., 2005). Today, the elevation along a
west-east transect through the Rocky Mountains,
Laramide province, and Great Plains decreases
from 3–4 km to 0.5 km, which strongly influences the δ18O value of modern precipitation and
the local isotopic lapse rate (Fig. 11). By early
Eocene time, crustal shortening in the Sevier had
caused significant crustal thickening, and formed
an elevated hinterland plateau to the west of the
Laramide province (DeCelles, 2004). Therefore
the paleotopography of Wyoming in early Eocene time was probably similar to present, with
the main vapor source from the ancestral Gulf of
Mexico. This justifies the comparison of the δ18O
values of the early Eocene precipitation with the
modern isotopic patterns.
Warmer global temperature in early Eocene
time could affect the relationship between the
δ18O value of precipitation and local temperature. Therefore, we compare the δ18O values of
early Eocene precipitation in the Wind River
Basin with modern summer precipitation isotopic gradient in the region (Fig. 11). The δ18O
values of early Eocene precipitation (–5.8‰ ±
1.9‰ after correction) in the Wind River and
Bighorn basins are similar to the values of the
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Basin analysis of the early Eocene Wind River Basin
Eocene paleosol, Wind River Basin (55–51 Ma)
Eocene paleosol, Bighorn Basin (54–53 Ma)
Eocene fluvial cement, Wind River Basin (55–51 Ma)
Modern summer mean precipitation
Modern unweighted mean annual precipitation
4
–18
–16
3
–14
–12
2
–10
–8
1
δ18O (SMOW)
Elevation (km)
–20
–6
0
Rocky Mountains
Laramide Province
–4
Great Plains
–2
120
115
110
105
100
95
90
Longitude (°°W)
Figure 11. Oxygen isotope data of early Eocene paleosol carbonate in the Wind River Basin
compared to the oxygen isotopic gradient of modern precipitation along an elevation transect of Rocky Mountains–Laramide province–the Great Plains. The light-gray curve is the
modern elevation profile along 43°N. Isotope data of Eocene paleosol carbonate in the Bighorn Basin are from Koch et al. (1995). Latitudes of sampling site for modern precipitation
δ18O values range between 40°N and 46°N. Modern precipitation isotope data are compiled from the data set of the United States Network for Isotopes in Precipitation (USNIP)
(www.uaa.alaska.edu/enri/usnip); Simpkins (1995); Harvey and Welker (2000); Harvey
(2001); Harvey (2005); Gammons et al. (2006). SMOW—standard mean ocean water.
modern summer precipitation at the same latitude of the Great Plains (~–5.5‰, Mead, Nebraska, elevation of 352 m) (Harvey, 2001), but
significantly higher than the values of modern
summer precipitation in the Laramide province
(–12.6‰, Butte, Montana, elevation of 1653 m)
(Gammons et al., 2006), suggesting the paleoelevation of the early Eocene Wind River Basin
and Bighorn Basin was low (~500 masl). The
low basinal elevation of the Wind River Basin
was stable in early Eocene because there is no
systematic variation of the δ18O value of paleosols through the stratigraphic section (Fig. 10).
The difference (9.8‰ ± 2.0‰) between the δ18O
values of early Eocene river water and basinal
rainfall suggests that the elevation difference between the surrounding Laramide ranges and basin floor was 3.4 ± 0.7 km, based on the isotopic
lapse rate of –2.9‰/km derived from modern
precipitation in the USA (Dutton et al., 2005).
Such an elevation difference is higher than the
modern difference of 1.5–2.5 km. The high
estimated elevations of the Laramide ranges
(~4 km) from our study are consistent with the
elevation estimates derived in previous paleobotanical studies for the early Eocene (Gregory
and Chase, 1992; Wolfe et al., 1998).
Carbon Isotopes, Paleoclimate, and pCO2
The δ13C value of soil carbonate is determined
by the relative proportion of C3 to C4 plants in
moist climate, and by the extent of local plant
cover in arid climate where soil respiration rates
are low (Quade et al., 1989; Cerling and Quade,
1993; Breeker et al., 2009). Modern climate in
the Wind River Basin is semiarid, with the mean
annual precipitation of ~200 mm. Our analyses show that the most common grasses and
the dominant plant, sagebrush, in the sparsely
vegetated basin are C3, with the δ13C values of
–24.5‰ ± 1.0‰. Soil carbonate formed in equilibrium with CO2 respired from pure C3 plants
will have a δ13C value of –12‰ to –10‰ (Cerling and Quade, 1993). The relatively high δ13C
values of modern soil carbonate in the Wind
River Basin (–2.4‰ ± 2.1‰) clearly suggests
mixing of atmospheric CO2 with plant-derived
CO2 owing to low soil respiration rates in the
local semiarid climate (Fig. 9).
The δ13C values of early Eocene paleosols
(–7.0‰ ± 2.0‰) in the Wind River Basin are
significantly lower than the values of modern
soil carbonate (Fig. 9), which suggests a wetter
climate during the Eocene. This is supported by
several other lines of evidence: (1) calcisols are
well developed in the floodplain of braided rivers in early Eocene Wind River Basin. The thick
(>50 cm) reddish argillic horizons that typify
Wind River Formation paleosols point to a relatively moist setting; and (2) abundant leaf fossils
of mixed deciduous and evergreen plants in the
Wind River Formation show high mean annual
temperature up to 21 °C and paleoprecipitation
up to 2000 mm/yr (Hickey and Hodges, 1975;
Wilf et al., 1998).
However, the δ13C values of early Eocene
paleosols are higher than the values of soil carbonate formed in wet settings today, in the presence of pure C3 biomass. This can be explained
by the presence of C4 plants or higher pCO2
during early Eocene time. There is no robust
evidence for the existence of C4 plants in early
Eocene (Cerling and Quade, 1993). This justifies the use of δ13C values of paleosols to estimate ancient atmospheric pCO2 levels based on
a diffusion-based model developed by Cerling
(1992), under the assumption of elevated soil
respiration rates expected in moist soil-forming
conditions. δ13C values of soil carbonate are determined by the δ13C value of soil CO2, which is
a mixture of soil-respired CO2 and atmospheric
CO2. Therefore, the δ13C value of soil CO2 is a
function of the concentration (pCO2) and δ13C
value of CO2 in the atmosphere, and the concentration and δ13C value of soil-respired CO2.
The δ13C value of ancient soil CO2 can be calculated from the measured δ13C value of paleosols,
using the temperature-dependent fractionation
equation (Romanek et al., 1992). The concentration of soil-respired CO2 is a function of soil
respiration rate and soil depth. Carbonate nodules were collected at least ~50 cm below the
paleosurface (the top of the paleosol profile). In
our reconstruction we assume: (1) soil respiration rates ≥4 mmol/m2/hr (Cerling, 1992; Ekart
et al., 1999); (2) the δ13C value of CO2 in the
atmosphere is assumed to be near pre-industrial
values of –6.5‰, and the values of plant-derived
CO2 to be –24.5‰; and (3) an exponential CO2
production function (Quade et al., 2007).
Using the average value of the three most
positive δ13C values of paleosols in the lower
Indian Meadows Formation, atmospheric pCO2
at ca. 55 Ma is estimated to have been 2050 ±
450 ppmV (based on 1σ of the δ13C value of the
paleosol carbonate). Using the average value of
the more negative δ13C values in the upper Indian
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991
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Fan et al.
B
Ra
ng
e
C
rR
an
ge
an
ge
W
as
Ri
ve
rR
ak
ie
ive
e
ier
Sev
Paleozoic carbonate rock
W
as
h
Wi
nd
R
an
ge
rR
ive
ki
?
ha
ki
Ra
e
Ra
ng
e
ng
e
?
1 km
Mesozoic siliciclastic rock
r
e
evi
Precambrian basement
ge
an
rR
ve
Ri
ha
?
S
Wind River
Basin
W
as
Wi
nd
R
Cenozoic sedimentary rock
W
in
d
A
Owl Creek
Mountains
d
in
W
ier
Sev
NW
Sevier
Washakie
Range
?
St
De
re
a
al mlu do
via m
l f in
an at
s e
br
isall flow
uv –
ial do
fa min
ns a
te
Br
a
id
d
ed
riv
er
d
s
Figure 12. Paleogeographic sketch maps of the northwestern Wind River Basin in early Eocene: (A) 55.2–55.0 Ma; (B) 55.0–54.5 Ma;
and (C) 51.0–53.0 Ma.
Meadows and Wind River formations, atmospheric pCO2 from 54 to 51 Ma is estimated to
have been 900 ± 450 ppmV. The calculated high
pCO2 values and declining early Eocene trend,
based on paleosols in central Wyoming, are
consistent with previous estimates of Ekart et al.
(1999) using the same method, and by Pearson
and Palmer (2000) using boron isotope ratios of
planktonic foraminifer shells.
REGIONAL PALEOGEOGRAPHY
Our reconstruction of the paleogeography
in the northwestern Wind River Basin during
the early Eocene is divided into three temporal
stages constrained by the relative stratigraphic
position of particular lithofacies associations,
the paleoelevations of the surrounding ranges,
the change in sedimentary provenance recorded by detrital zircon age spectra, and absolute ages based on fossil assemblages in the
basin (Fig. 12).
The lower Indian Meadows Formation was
deposited unconformably on upper Cretaceous
Cody shale in the East Fork area. The depositional hiatus formed at least partly by folding
in front of the Washakie Range (Winterfeld and
Conard, 1983). Deposition evolved from small
braided rivers into debris flows on proximal
alluvial fans along the range front. Southwardflowing debris flows delivered mainly Mesozoic and Paleozoic detritus. The gravitational
force associated with steep slope is important
to the initiation of debris flows (Stanistreet and
McCarthy, 1993; Coussot and Meunier, 1996).
Therefore, we infer that the Washakie Range to
992
the north of the region was high during the early
Eocene (Fig. 12A), which is also supported by
stable isotope paleoaltimetry (~4 km).
Deposition in the upper Indian Meadows
Formation was dominated by shallow braided
rivers on alluvial fans (Fig. 12B). Southward
flow delivered late Paleozoic clasts from the
Washakie Range into the basin, and the Precambrian basement core in the Washakie Range
was gradually exposed. The change from debris
flows to braided rivers suggests the slope between the mountain range and basin floor decreased (Stanistreet and McCarthy, 1993). Such
a change of slope may have been caused by the
rapid erosion of the mountain range and accumulation of sediment in wet climate.
Braided rivers flowing east dominate the
sedimentary environment of the Wind River
Formation (Fig. 12C). We argue that the change
of paleoflow direction was caused by the development of a northeastward paleoslope along the
Wind River Range during or after the deposition
of the Indian Meadows Formation, and before
deposition of the Wind River Formation. Paleoslopes along both the Wind River and Washakie
ranges formed a confined valley during the
early Eocene, and low-sinuosity rivers were
developed in the narrow basin when discharge
and sediment yield were high. The dominant
northeastward paleocurrent direction and granite cobbles derived from the Wind River Range
suggest that the range became the major source
terrane of sediment in the northwest part of the
basin from ca. 53 to 51 Ma. The abundance of
basement granite clasts, feldspar, and Archean
zircons suggests that the basement core of the
Wind River Range had been exposed by ca. 53 Ma
(Fig. 10). Low-sinuosity rivers that developed
upstream could have merged into a large eastflowing meandering river in the basin (Soister,
1968; Seeland, 1978). Therefore, a paleodrainage pattern similar to the present one was developed by late early Eocene time in the northwest
Wind River Basin. The local relief based on
oxygen isotope ratios was 2.3 ± 0.8 km, which is
also comparable to modern relief in the region,
but with lower basin floor.
IMPLICATIONS FOR TECTONICS
Rapid Late Paleocene–Early Eocene Uplift
of Basement-Involved Ranges
Both the Wind River and Washakie ranges
had an elevation contrast of 2.3 ± 0.8 km relative to the basin by late Paleocene–early Eocene time, and the basin floor stood at ~0.5 km
(asl). Rapid late Paleocene–early Eocene tectonic uplift and growth of surface topography
throughout western Wyoming are supported by
evidence for synorogenic sedimentation, reflection seismic records, and thermochronology.
Alluvial fan conglomerates of late Paleocene–
early Eocene age, indicating intense unroofing of Laramide source terranes, also occurred
along the east flank of the Beartooth Range,
east of the Bighorn Range, north of the Uinta
Mountains, and on the south flank of the Wind
River Range (Gries, 1983; DeCelles et al., 1991;
Crews and Ethridge, 1993; Hoy and Ridgway,
1997). Seismic and borehole records from the
south flank of the Wind River Range show that
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Basin analysis of the early Eocene Wind River Basin
Precambrian rocks are in thrust contact above
lower Eocene sedimentary rocks (Gries, 1983,
and references therein). Thermochronological
studies of the Beartooth Range and Wind River
Range also show that of the mountains exhumed
4–8 km in late Paleocene–early Eocene time
(Omar et al., 1994; Peyton and Reiners, 2007).
This exhumation amount is significantly higher
than the thickness of Phanerozoic sedimentary
cover on the Precambrian cores.
Uplift of the Wind River Range may have
started in Late Cretaceous as a paleosol developed at the base of the Paleocene deposits on
the east flank of the Wind River Range (Gries,
1983). Uplift during Late Cretaceous would
be consistent with data from numerous other
Laramide uplifts and basins. For example, the
Hoback Basin to the west of the Washakie and
Wind River ranges contains thick uppermost
Cretaceous and Paleocene sediments with
abundant conglomerates (Love, 1973); and the
upper Cretaceous–lower Paleocene Beaverhead Conglomerate and Maastrichtian Sphinx
Conglomerate in southwestern Montana indicate uplifting of the Blacktail-Snowcrest and
Madison-Gravelly arches, respectively (Haley,
1986; DeCelles et al., 1987). It is not clear
how such evidence of Late Cretaceous tectonism of the Laramide ranges was expressed
in terms of paleoelevations, but the abundance
of coarse-grained alluvial fan deposits containing proximal sediment-gravity flows, as well as
provenance studies, indicate substantial topographic relief (Graham et al., 1986; DeCelles
et al., 1991). Large-scale tectonic exhumation
and unroofing of the Laramide ranges in western
Wyoming during late Paleocene–early Eocene
time could have been enhanced by the relatively
humid paleoclimate. Studies in the Himalaya
and Washington Cascades have shown that high
precipitation is coupled with rapid erosion and
sediment denudation rate (Reiners et al., 2003;
Thiede et al., 2005).
Post-Early Eocene Regional Uplift
Currently the Laramide basins in Wyoming
have an average elevation of ~1.5 km, roughly
1 km higher than their elevations during the
early Eocene, based on oxygen isotope ratios of
basinal precipitation recorded in paleosol nodules. Therefore, the Laramide basin floors in
Wyoming must have reached their present elevation after early Eocene time.
Three mechanisms are proposed to explain
the surface uplift of Laramide basins after early
Eocene: (1) removal of Farallon slab or thickened mantle lithosphere in western USA in middle Cenozoic time (Dickinson and Snyder, 1978;
Humphreys, 1995; Sonder and Jones, 1999; Liu
et al., 2010), (2) late Cenozoic isostatic rebound
of lithosphere due to climate-controlled or enhanced sediment erosion (Pelletier, 2009), and
(3) late Cenozoic mantle thermal uplift caused
by the thermal upwelling associated with the
initiation of the Rio Grande Rift (Heller et al.,
2003; McMillan et al., 2006).
Southward migration of post-Laramide magmatism in the northern Basin and Range is argued as the evidence for southward removal of
Farallon slab or mantle lithosphere (Humphreys,
1995; Sonder and Jones, 1999). The resulting
asthenospheric upwelling could have caused
regional uplift. Post-Laramide magmatism occurred in northwestern Wyoming and Idaho at
ca. 50 Ma (Armstrong and Ward, 1993), which
predicts the timing of regional uplift in the
Laramide province in Wyoming. However,
there is no robust evidence suggesting high basinal elevation before early Eocene (Wolfe et al.,
1998; Sjostrom et al., 2006; this study). Therefore it is unlikely that the removal of Farallon
slab caused the post-early Eocene regional elevation gain in the Laramide foreland.
Thermo-geochronological and sedimentological data suggest an episode of Miocene exhumation in the Laramide ranges, which could
have caused regional surface uplift (Cerveny
and Steidtmann, 1993; Omar et al., 1994;
Strecker, 1996; Crowley et al., 2002; Peyton
and Reiners, 2007). A recent reconstruction of post-Laramide basin fill also shows
that more than 1 km of incision occurred in
the Laramide basins and the western Great
Plains since late Miocene time (McMillan
et al., 2006). Therefore, the Miocene uplift
due to climate-enhanced erosion and/or mantle
dynamic uplift most likely formed the present
landscape in central Wyoming (Heller et al.,
2003; McMillan et al., 2006). However, more
paleoelevation data covering Cenozoic time
are required to evaluate the amount of regional
uplift during the late Cenozoic. Our research
does not preclude dynamic uplift between
early Eocene and late Miocene time.
CONCLUSIONS
This multidisciplinary study of the lower
Eocene strata in the northwestern corner of the
Wind River Basin improved our understanding of the timing and process of basin evolution and source terrane unroofing, changes in
paleoelevation of basin floor and surrounding Laramide uplifts, and paleoclimate during
Laramide deformation.
(1) Petrographic, detrital zircon U-Pb geochronology, and paleocurrent analyses show the
Washakie Range and western Owl Creek Mountains to the north of the basin experienced rapid
unroofing at 55.5–54.5 Ma, and the Wind River
Range to the southwest became the dominant
sediment source terrane by ca. 53–51 Ma. Rapid
tectonic uplift of the two mountains in the late
Paleocene–early Eocene formed a confined valley, and a paleodrainage similar to the present
with rivers flowing from both ranges was formed
in the northwestern Wind River Basin by 51 Ma.
The sedimentary environment evolved from alluvial fan into low-sinuosity, gravel-bed braided
rivers, and paleocurrents changed from southwestward into mostly eastward.
(2) Detrital zircon grains equivalent to the
depositional age are absent in early Eocene
sediment. Recycled detrital zircon grains derived from the Grenville and Yavapai-Mazatzal
orogens are present in Mesozoic and Paleozoic
sedimentary rocks exposed in the Laramide
ranges, and dominate the zircon populations of
both modern river sand and early Eocene sandstones in the Wind River Basin. Presently exposed crystalline Archean basement in the Wind
River Range contributes ~20% of the total zircon grains to basinal sediments. Zircons of different ages from Precambrian basement rocks
exposed in Laramide ranges are a potentially
useful tracer of sedimentary provenance.
(3) A more humid climate in the early Eocene
Wind River Basin is inferred from the high soil
CO2 respiration rate compared to today. Calcisols were extensively developed in the floodplain deposits. Atmosphere pCO2 estimated
from paleosol δ13C values decrease from 2050 ±
450 ppmV to 900 ± 450 ppmV during the early
Eocene, consistent with previous reconstructions of pCO2 for this period.
(4) The elevation of the Wind River Basin
was comparable with the modern Great Plains,
on the order of ~500 m (asl), and the local relief
between the Washakie Range and western Owl
Creek Mountains, Wind River Range, and Wind
River Basin was 2.3 ± 0.8 km during the early
Eocene. Post-Laramide regional uplift of ~1 km
is required to form the present landscape in
west-central Wyoming, which was most likely
caused by regional Miocene–Pliocene uplift and
exhumation.
ACKNOWLEDGMENTS
This research was supported by the Robert J.
Weimer Fund from (SEPM) Society for Sedimentary Geology, a Donald L. Smith award from Rocky
Mountain Section of the Society for Sedimentary
Geology, a ChevronTexaco scholarship, ExxonMobil
grant, and National Science Foundation grant EAR0732436 (to the Arizona LaserChron Center). We
thank Tamara Goldin and Dan Ross for assistance in
the field, Joel Saylor and Facundo Fuentes for help
with petrographic work, and Antoine Vernon for picking zircons. Reviews and editorial comments by Eric
Erslev, Kenneth Ridgway, Robert H. Rainbird, and
Brendan Murphy improved both text and figures.
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993
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Fan et al.
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