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Geomorphic controls on lacustrine isotopic compositions: Evidence from the Laney Member, Green River Formation, Wyoming

Amalia C. Doebbert1,†, Alan R. Carroll1, Andreas Mulch2,3, Lauren M. Chetel1, and C. Page Chamberlain2 1Department of Geology and Geophysics, University of Wisconsin–Madison, Madison, Wisconsin 53706, USA 2Department of Geological and Environmental Sciences, Stanford University, Stanford, California 94305-2115, USA 3Institut für Geologie, Universität Hannover, 30167 Hannover, Germany

ABSTRACT INTRODUCTION interpretation of downstream isotopic records consequently requires understanding the evolu- Oxygen isotope values from lacustrine car- Oxygen isotope records from lacustrine car- tion of upstream drainage networks. bonate in the Laney Member of the Green bonate, paleosol carbonate, and phosphatic The Green River Formation of Wyoming River Formation (Wyoming) exhibit a sud- continental faunal remains are commonly used provides a unique opportunity to test the signifi - den, basinwide, ~6‰ upsection decrease in to recognize local and regional climatic varia- cance of δ18O variation produced by changes in δ18O at ca. 49 Ma. 40Ar/39Ar geochronology tion, and the Paleogene basins of western North regional drainage patterns. Green River Forma- constrains the duration of the isotopic shift to America have been the setting for several such tion strata are among the most extensively stud- ≤~200,000 years. This change coincides with studies. δ18O records from these basins have ied in the world, providing a well-constrained a sudden change in lake type, from balanced- been used to support estimates of Paleocene stratigraphic and chronological framework (e.g., fi lled in the lower LaClede Bed to overfi lled in seasonality and Paleocene–Eocene warming Bradley, 1929, 1964; Eugster and Hardie, 1975; the upper LaClede Bed, as well as an increase (Dettman and Lohmann, 1993; Fricke et al., Surdam and Wolfbauer, 1975; Surdam and Stan- in the proportion of calcitic (>80% calcite 1998; Koch et al., 2003), to infer Eocene re- ley, 1979, 1980; Smoot, 1983; Roehler, 1973, out of total carbonate by X-ray diffraction gional atmo spheric circulation patterns (Fricke, 1992, 1993; Carroll and Bohacs, 1999; Smith [XRD]) samples from 32% to 73%. The δ18O 2003), and to estimate paleoaltitude (Norris et al., 2003, 2008). Furthermore, several authors shift is correlatable through several locations et al., 1996, 2000; Dettman and Lohmann, have suggested that changes in regional drainage across the Greater Green River Basin, and 2000). These studies demonstrate the potential infl uenced the depositional history of the Green also coincides with a previously observed shift for δ18O to document orogenic uplift histories River Formation. Dynamic interbasin drainage to less radiogenic 87Sr/86Sr. Minimum δ18O and associated climatic effects, a possibility relationships are cited as a control on the evolu- values observed are the same as values pre- which has generated particular interest in δ18O tion of sediment dispersal in the Laramide basin viously reported in aragonitic bivalves from records from intermontane settings worldwide system (Seeland, 1985; Dickinson et al., 1988), the same unit, indicating that low δ18O in this (e.g., Chamberlain et al., 1999; Garzione et al., while in the Greater Green River Basin previ- record is not diagenetic. A simultaneous shift 2000; Kleinert and Strecker, 2001; Poage and ous work has documented drainage diversion to evaporative conditions in the Uinta Basin Chamberlain, 2001; Rowley et al., 2001; Blis- by uplift (Pietras et al., 2003), sediment infi ll- to the south indicates that the δ18O shift and niuk and Stern, 2005; Kent-Corson et al., 2006). ing (Surdam and Stanley, 1980), and the expan- lake-type change are not driven by regional Most δ18O-based paleoclimate and paleo- sion of drainage networks during lake highstand climatic cooling and/or humidity increase. elevation studies rely on a simplifying assumption (Rhodes et al., 2002). This paper presents δ18O We propose that all of these observations re- that the upstream drainage network supplying data across a lithologically defi ned lake-type sulted from the capture by Lake Gosiute of a water to the system remains unchanged during change within the Laney Member of the Green river that drained higher elevations in central the period of record. However, upstream changes River Formation, and considers the hypothesis or north-central Idaho. Mass-balance model- in the long-distance regional drainage patterns of that a large negative shift observed in these data ing of Eocene Lake Gosiute indicates that rivers, lakes, and groundwater have the potential provides geochemical evidence for a drainage capture of a river with an annual average to affect these downstream δ18O records. Particu- capture event. discharge of ~20 billion m3/a (slightly larger larly in mountainous settings, precipitation δ18O than the modern Snake River) and δ18O of values can be highly variable within a small geo- GEOLOGIC SETTING –24‰ standard mean ocean water (SMOW) graphic area, and downstream basins may merge or lower would be capable of producing the waters that have traveled hundreds of km from Deposited by Lake Gosiute in the Greater observed change. A more likely alternative their sources. Several factors may produce rapid Green River Basin between ~53 and 48.5 Ma is a river with less negative δ18O and greater catchment reorganization causing abrupt down- (Smith et al., 2008), the Green River Formation discharge. For example, if river waters had stream sedimentary changes, including drainage records a wide range of depositional conditions a δ18O composition of ~−16‰, an estimated capture by knickpoint migration (e.g., Zaprowski including fl uvial-lacustrine, fl uctuating-profundal river discharge of ≥50 billion m3/a would pro- et al., 2001), landslide activity (Korup, 2005), lacustrine, and evaporative lacustrine (cf. Carroll duce the same effect. and drainage diversion by tectonic uplifts (e.g., and Bohacs, 2001). Lacustrine deposition Dickinson et al., 1988; Okay and Okay, 2002; reached its maximum areal extent (~40,000 km2) †E-mail: [email protected] Smith et al., 2008; Davis et al., 2009). Accurate during deposition of the LaClede Bed of the

GSA Bulletin; January/February 2010; v. 122; no. 1/2; p. 236–252; doi: 10.1130/B26522.1; 9 fi gures; 2 tables; Data Repository item 2009190.

236 For permission to copy, contact [email protected] © 2009 Geological Society of America Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Geomorphic controls on lacustrine isotopic compositions

Laney Member (Roehler, 1993). Airfall tuffs Green River Formation coincided with the peak Estimates of relief between the basin and inter bedded with lacustrine strata provide means and initial decline of the Early Eocene Climatic surrounding uplifts also differ widely. Conifer of correlation and excellent age control for Green Optimum (Zachos et al., 2001; Smith et al., pollen observed in the Green River Formation River sediments (Smith et al., 2003, 2008). 2003, 2006, 2008). Locally, estimates based has been inferred to represent elevations greater The Greater Green River Basin is a portion on leaf size and morphology (Wing and Green- than 3000–4000 ft (~900 m–1.2 km), which of the Sevier foreland (Fig. 1), which was seg- wood, 1993; Greenwood and Wing, 1995; Wilf, may indicate as much as 5000 ft (~1.5 km) of mented by basement-cored Laramide uplifts be- 2000) as well as fl oral and faunal assemblages relief for basin-bounding uplifts (MacGinitie , ginning in the Maastrichtian (Dickinson et al., (Leopold and MacGinitie, 1972; Wing and 1969). Norris et al. (1996, 2000) report δ18O 1988; DeCelles, 2004). Lacustrine deposition Greenwood, 1993; Markwick, 1994) indicate from micritic lacustrine carbonates in the in the basin coincided with the late stages of that temperatures rarely or never dropped below Wilkins Peak Member of the Green River For- Laramide uplift in southern Wyoming (Dickin- freezing at basin elevations. Leaf margin and mation of as low as –15.8‰ PeeDee belemnite son et al., 1988; Carroll et al., 2006). The adja- leaf area analysis of the Little Mountain fl ora (PDB) (14.6‰ SMOW) and propose that snow- cent Sevier plateau had already reached high in the upper Wilkins Peak Member (ca. 50 Ma; melt-derived water originating at an elevation of elevations by the late Cretaceous, although com- Smith et al., 2008) and lower Laney Member >3 km was entering the lake from surrounding pression continued through the Early Eocene as (ca. 49.3 Ma; Smith et al., 2008) suggest mean uplifts. However, low δ18O has been shown by indicated by motion on eastern Sevier thrusts annual temperatures of 19.6 ± 2.1 °C and mean Morrill and Koch (2002) to have an association as recent as ca. 50 Ma (DeCelles, 2004). Cor- annual precipitation of 75.8 + 33.2/−23.2 cm with diagenesis in bivalves from the Luman dilleran collapse, represented by the formation in the Greater Green River Basin (Wilf, 2000). Tongue and Laney Member of the Green River of extensional basins (Constenius, 1996) and The younger Green River fl ora collected near Formation, which they suggest could also ac- crustal-scale extensional faulting (Mulch et al., Bonanza , Utah (ca. 47.3 Ma; Smith et al., count for sporadically low δ18O values in lacus- 2007), initiated between ~50 and 48 Ma at the 2008) indicate a slightly cooler and wetter cli- trine sediments from the Wilkins Peak (Morrill latitudes of the Greater Green River Basin and mate, with mean annual temperatures between and Koch, 2002; Fricke, 2003). 49–47 Ma farther north. 14.3 ± 2 °C and 15.2 ± 2 °C (Wing and Green- Laney Member sediments that overlie the wood, 1993) and mean annual precipitation of Absaroka Volcanic Province youngest rocks displaced by the Wind River 84 + 36.2/−25.3 cm (Wilf et al., 1998). How- Paleobotanical assemblages suggest that Thrust on the northern basin margin suggest that ever, these conditions still represent a warm, paleo relief in the Absaroka Volcanic Province this Laramide fault was inactive by ~49–50 Ma seasonally dry, setting. was 10,000 ft (~3 km) or more (Fritz, 1980). (Steidtmann and Middleton, 1991). In the south- Today, the volcanic edifi ce maintains modern ern Greater Green River Basin, Dickinson et al. Regional Paleoelevation Estimates local thicknesses of as much as ~1.5 km fol- (1988) cites fl uvial conglomerates deposited lowing extensive erosion (Smedes and Prostka, along the margin of the Washakie Basin as late Greater Green River Basin 1972). δ18O evidence from riverine bivalves as 35–40 Ma as evidence that Laramide activity Eocene elevation of the Greater Green River supports 2.5–3 km of local relief in the Late continued through the Late Eocene. However, Basin is highly contentious. The difference be- Cretaceous–Paleogene Laramide Rockies around the expansion of low-energy depositional en- tween enthalpy estimates from the inland Little the Powder River Basin and Hell Creek Basin, vironments adjacent to the North Uinta Thrust Mountain fl ora and those of a coeval coastal as well as the presence of persistent snowfi elds suggests decreased subsidence during Laney assemblage has been cited as evidence that the in the mountains of Montana and northern Wyo- deposition, an indication that sedimentation elevation of the basin was greater than 2.1 km ming (Dettman and Lohmann, 2000). If these outpaced load-induced subsidence during the (Wolfe et al., 1998). However, this method is estimates are correct, some areas of high Paleo- mid-Eocene (Beck et al., 1988). likely to overestimate paleolatitude if “coastal” cene elevation in Montana and Wyoming may The Absaroka Volcanic Province, located assemblages represent embayments or valleys have persisted into the Eocene, producing high to the north of the Greater Green River Basin with warmer temperatures than are characteris- relief even while mean elevation was falling (Fig. 1), produced predominantly andesitic and tic of the coastline as a whole (Fricke and Wing, (e.g., Mulch and Chamberlain, 2007). An abun- basaltic volcanics beginning ca. 55 Ma (Harlan 2004). In an attempt to circumvent this problem, dance of conglomerate during the early Eocene et al., 1996; Hiza, 1999; Feeley et al., 2002). Fricke and Wing (2004) compared the Niland, in deposits of the Wind River, Bighorn, Gros Volcanism in the Absarokas built large strato- Sourdough, Latham, and Little Mountain fl ora Ventre, and Hoback basins (Love, 1939; Keefer, cones and generally progressed from NW to to a regression-based, predicted mean paleotem- 1957, 1965; Dorr et al., 1977; Krause, 1985) also SE, gradually approaching the Wind River perature for coastal sites between ~41 and ~45 attests to signifi cant relief in the area. Mountains and northern margin of the Greater degrees north latitude. This alternative approach Green River Basin (Smedes and Prostka, 1972; yielded an average temperature difference of Challis Volcanic Field and North-Central Sundell, 1993; Feeley and Cosca, 2003). Farther only 3.2 °C, corresponding to elevations of only Idaho west, the Challis Volcanic Field (Fig. 1) was ac- ~0.6–1.3 km. Additional studies also suggest The paleorelief of the Challis Volcanic Field tive between ~50 and ~45 Ma (McIntyre et al., that the basin was at relatively low elevation. is a subject of debate. Hydrogen isotope values 1982; Fisher et al., 1992), and also produced Bradley (1929) inferred from structural and ero- from southeastern British Columbia indicate large volumes of calc-alkaline volcanic material. sional history that the basin fl oor was less than precollapse elevations as great as ~4000 m in 1000 ft (~300 m) above sea level in the Eocene, parts of the North American Cordillera (Mulch Paleoclimate and similar altitudes are suggested by compari- et al., 2004, 2007). The Bitterroot metamorphic sons of the Eocene Green River, Little Moun- core complex and Pioneer core complex, asso- Oceanic δ18O records indicate that the early tain, Boysen, Tipperary, and Lost Cabin fl oras ciated with the Cordilleran collapse, bracket Eocene was the warmest part of the Cenozoic in western Wyoming to the elevations of similar the Challis Volcanic Field to the north and (Pearson et al., 2007), and deposition of the modern assemblages (Axelrod, 1968). south, and both suggest that active displacement

Geological Society of America Bulletin, January/February 2010 237 Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Doebbert et al.

47 1 234 Paleozoic quartzite in 48 ? DrainageLaramide Basins 49 Br Capture Boulder 50 Batholith 1: Gros Ventre Basin—Love, 1947 51 Wa7 Lowland Creek 52 3: Wind River Basin—Keefer, 1957 Volcanics 53 Wa6 46°N ~54–50 Ma Wa5 2: Bighorn Basin—Krause, 1985 54 P 4: NW GGRB—Dorr et al., 1977 K A: Harebell-Pinyon Conglomerate—Lindsey, 1973

~55–51 Ma MT WY Challis Bighorn Uplift Volcanic Idaho Batholith Field Absaroka ~51–48 Ma Volcanic Province 2 Bighorn ~50–45 Ma A ~48–44 Ma Basin 44°N Snake River Plain 1 N 3 Eocene Wind River Uplift Wind River Basins 4 Stratigraphic section Basin Sampled section Laney Unionid bivalves (C. Morrill, pers. comm. 2005) Low δ18O fluid inclusions (Seal and Rye, 1993) Paleoflows (Surdam and Stanley, 1980) ID Greater Green Inferred paleodrainage 42°N (after Janecke et al., 2000) UT River Basin Volcanic and plutonic rocks Sevier Fold and Thrust Belt Thorofare group (AVP) GRCC FC SB WB2 Sunlight group (AVP) SC WB1 Washburn group (AVP) CCR GD Challis Volcanic Province and other volcanics

40 39 ~50–46 Ma Ar/ Ar and K/Ar ages Pre - Cretaceous Uinta Uplift Cretaceous Bitterroot metamorphic core complex Plutonic rocks Probable quartzite source (Paleozoic) Precambrian rocks and Uinta Cambrian-Jurassic 100 km sediment Basin 40°N 114°W 112°W 110°W 108°W

Figure 1. Simplifi ed geologic map of the Greater Green River Basin and surrounding areas after Witkind and Grose (1972), Smedes and Prostka (1972), Wilson and Skipp (1994), Fisher et al. (1992), and House et al. (2002). The locations of measured sections shown include: GRCC—Green River Community College, SB—Sand Butte, GD—Grainery Draw, SC—Sage Creek, WB1—Arco Oil and Gas Washakie Basin #1 core, WB2—Arco Oil and Gas Washakie Basin #2 core, and CCR—U.S. Department of Energy Currant Creek Ridge core. Also shown is a measured section at Firehole Canyon (FC) from Rhodes (2002). Previously published paleocurrent indicators from the Sand Butte Bed (Surdam and Stanley, 1980) indicate drainage to the southwest, and documented conglomerates including Pinyon-type clasts (Love, 1947; Keefer, 1957; Dorr et al., 1977; Krause, 1985) demonstrate transport of water and sediment from central Idaho.

238 Geological Society of America Bulletin, January/February 2010 Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Geomorphic controls on lacustrine isotopic compositions maintained high relief during the Eocene. As Laney much as 13 km of extensional unroofi ng had Stratigraphy already taken place on the western footwall of the Bitterroot metamorphic core complex by Volcaniclastic-rich Br WB#1 Core deltaic sand ca. 49 Ma (House et al., 2002) and the meta- δ18O Depth 87Sr/86Sr morphic core of the Pioneer core complex was (ft) 20 25 30 0.711 0.712 Finely laminated also exposed prior to the initiation of Challis calcitic micrite: vol canism (O’Neill and Pavlis, 1988). 341 overfilled lake The only direct constraint on paleoelevation 350 for the Challis is the Eocene Thunder Moun- SBB 360 Dolomitic siltstone: abundant ripple tain fl ora, which estimates an altitude of only 370 marks ~1730 m (Axelrod, 1998). However, this paleo- 380 fl oral assemblage comes from deposits within 390 Dolomitic laminated a caldera and may greatly underestimate the Shale elevations of other parts of the Challis Volcanic Middle 400 ULB 410 Shallow facies Field, which likely included rapidly accumu- x including lated high relief which is no longer preserved. 420 stromatolites, oolite/ostracod It is clear that signifi cant volumes of material “Fill to Spill” grainstones, and Eocene mudcracks have been removed by erosion from the Challis Bridgerian xxx Analcite Tuff area, and high-elevation areas were likely (48.94 ± 0.12 Ma) Dolomitic and prefer en tially removed. Pebble to boulder con- BMB calcitic laminated glomerates affi liated with structural culmina- micrites alternating LLB 560 with shallow facies: tions in north-central Idaho are interbedded balance-filled lake with Challis volcanic material in middle Eocene 570 paleo valley fi ll (Janecke et al., 2000), and sup- 580 Underfilled lake port the idea of nearby high relief. 590 ~22‰ ~27‰ Fluvial sediments LANEY MEMBER STRATIGRAPHY ULC LLC Avg Avg Dessication Cracks Early WP x Tuff Stromatolite The lower LaClede Bed of the Laney Mem- xx Sixth Tuff Calcite Ostracod CB (49.62 ± 0.10 Ma) Mix Concretion ber, originally described by Roehler (1973), is Dolomite Rip-up clasts characterized by alternation between shallow and deeper-water lithologies (Fig. 2). Flooding Figure 2. Generalized Laney Member stratigraphy, with representative lithology, δ18O, and surfaces, sometimes marked by rip-up horizons, 87Sr/86Sr from the Washakie Basin #1 core (WB1). Units represented include the Cathedral are overlain by shallow stromatolite and ostra- Bluffs Tongue of the Wasatch Formation (CB), the Wilkins Peak Member of the Green cod, oolite, or pisolite grainstone facies which River Formation (WP), the lower LaClede Bed (LLC), Buff Marker Bed (BMB), upper grade into organic-rich laminated micrites (oil LaClede Bed (ULC), and Sand Butte Bed (SBB) of the Laney Member, Green River For- shales) representing deeper-water conditions. mation, and the Bridger Formation (Br). Average δ18O is high in the lower LaClede, with Shallow horizons are laterally extensive (Fig. 3), an average around 27‰ standard mean ocean water (SMOW) and maximum variation of and some stromatolites and grainstone beds ~3‰. Complete δ18O data are available in Table DR1 (see footnote 1). Shallow facies in the can be traced throughout the Washakie Basin lower LaClede correspond to radiogenic Sr, dolomitic mineralogy, and high δ18O. Sr, from (Shultz et al., 2002, 2004). In typical cycles, Rhodes (2002), exhibits a shift toward less radiogenic values coincident with the δ18O shift laminated micrites gradually transition back to across the lower LaClede–upper LaClede boundary. mud-cracked dolomicrite exhibiting casts of evaporite minerals such as nahcolite and trona (Rhodes, 2002). Fish and other fossils are abun- of intense desiccation (Rhodes et al., 2007). equivalent stromatolites, ostracod grainstones, dant in the lower LaClede but are often found Lake cycles resume for ~10–15 m above the and fl at-pebble conglomerates. At sites which concentrated in narrow horizons. These facies Buff Marker Bed, and in order to group sedi- lack these facies, the Analcite Tuff 6–8 m be- successions record fl uctuation of Lake Gosiute ments with similar overall lithofacies character- low the contact is the nearest clear marker. It is between the basin fl oor and its sill level (Surdam istics, this work includes lake cycles above the inferred that the lower-upper LaClede boundary and Stanley, 1979; Rhodes et al., 2002), likely Buff Marker Bed as part of the lower LaClede represents permanent fi lling of Lake Gosiute to with spillover into downstream Lake Uinta dur- (Fig. 2) after the manner of Carroll et al. (2008). a spill point located to the southeast (Carroll and ing highstands (Smith et al., 2008). The stratigraphic boundary between the lower Bohacs, 1999; Rhodes, 2002), and consequently Lacustrine deposition in the lower LaClede and upper LaClede Bed is represented by a re- the lower LaClede–upper LaClede boundary is Bed is interrupted by a distinct marker horizon duction in the dolomite content of micritic lacus- referred to as a “fi ll to spill” surface. known as the Buff Marker Bed (Roehler, 1973, trine carbonate (Fig. 3). The contact can also be The upper LaClede Bed is dominated by 1992, 1993). Marked by meter-scale desiccation identifi ed by the initiation of the isotopic shift fi nely laminated, organic-rich calcimicrite. The cracks and a 10–15 m siltstone in the Washakie addressed by this paper (Fig. 3). In the fi eld, the preservation of fi ne lamination demonstrates Basin (Figs. 2 and 3), the Buff Marker Bed has boundary can be recognized by an underlying an absence of burrowing organisms, and sug- previously been interpreted to represent a period bench-forming surface composed of laterally gests that deposition took place below wave

Geological Society of America Bulletin, January/February 2010 239 Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Doebbert et al.

Greater Green River Basin N a GRCC b FC SB WB2 FC SC WB1 CCR a' b' GD δ18O A 2520 CCR Upper LaClede Bed

Analcite Tuff

Isotopic shift GRCC 48.94 ± 0.12 Ma δ18O SC 25 30 Buff Marker Bed Lower LaClede Bed

6th Tuff 49.70 ± 0.10 Ma Wilkins Peak Member a a′

Fluvial Underfilled lacustrine δ18O 87Sr/86Sr Overfilled Dolomite Dolomite Measured Upper Lower lacustrine Volcaniclastic deltaic Section LaClede LaClede 5 m Mixed Mixed Average Average Balance-filled Fluvial deltaic Fischer 10 km lacustrine Calcite Calcite Assay (~22 permil) (~27 permil) Tuff Shallow Facies Base/Top of oxygen isotope shift

B GD

Sand Butte Bed δ18O 20 25 30 SB

Antelope sand (48.70 ± 0.19 Ma) ** 18 δ O WB2 δ18O 20 25 30 87 86 WB1 Sr/ Sr 20 25 30 18 0.711 0.712 0.713 δ O 87Sr/86Sr 20 25 30 Upper 0.711 0.712 LaClede Bed

Analcite Tuff 48.94 ± 0.12 Ma Isotopic shift Buff Marker Bed

Lower LaClede Bed b Cathedral Bluffs Tongue (Wasatch Formation) b′

Figure 3. Correlation of stratigraphy and isotopes (A) west of the Rock Springs Arch and (B) east of the Rock Springs Arch. δ18O data from this study are linked with Sr analyses of the same samples done by Rhodes (2002). 40Ar/39Ar ages are from Smith et al. (2003, 2008). A shift toward lower δ18O of ~6‰ takes place across the transition between the lower LaClede and upper LaClede Bed of the Laney Member, and reinforces correlations based on the Analcite tuff, Sixth tuff, the Buff Marker Bed, and stromatolite beds.

240 Geological Society of America Bulletin, January/February 2010 Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Geomorphic controls on lacustrine isotopic compositions base. Fish, freshwater ostracods, and leaves variation on δ18O, sample mineralogy was deter- applied to data to account for variations in ma- are abundant but irregularly distributed in the mined on a Sintag PAD V X-ray diffractometer chine performance during different runs. All upper LaClede, while shallow-water facies using a Cu Kα X-ray source (λ = 1.5418 Å) in oxygen data are reported relative to Vienna stan- and desiccation indicators are absent. Fresh- the S.W. Bailey X-ray Diffraction Laboratory at dard mean ocean water (VSMOW), and aver- water bivalves have also been reported in the the University of Wisconsin–Madison. ages weighted by standard deviation were used upper LaClede (Bohacs et al., 2001; Morrill Scans of powdered sample were run between for duplicate analyses. and Koch, 2002). The consistency of profundal 20° and 55° 2θ, a range which encompasses In order to test possible effects on δ18O of facies in the upper LaClede Bed implies that all signifi cant calcite and dolomite peaks. Step organic matter in Green River carbonates, eight lake level remained continually fi lled to its spill size for all scans was 0.02°, and step time was organic-rich subsamples and one aliquot of point during deposition. Above this surface, the either one second or two seconds, based on the MERCK internal calcite standard (0% organic volume of volcanic-derived sediment gradu- time required to obtain adequate count intensity. matter) were treated overnight with a 3% ally increases, and fi nal sedimentary infi lling of Rela tive areas of the sums of all dominant cal- H2O2 solution to remove organic material (cf. Lake Gosiute is represented by the deltaic dolo- cite peaks (peaks between 29.30° and 30.00° 2θ) Boiseau and Juillet-Leclerc, 1997). The possible mitic siltstones and volcaniclastic sandstones and all dominant dolomite peaks (peaks between infl uence of organic material was evaluated by that comprise the Sand Butte Bed. The Sand 30.40° and 31.10° 2θ) were used to estimate comparison of δ18O from treated and untreated Butte Bed is time-transgressive, and found both the calcite-dolomite proportion of carbonate for aliquots. Samples within 0.4‰ of each other laterally equivalent to and stratigraphically each sample according to the relationship are analytically indistinguishable, and seven of overlying LaClede Bed lacustrine facies (Figs. eight treated samples are within 0.3‰ of cor-

2 and 3). Associated sands progressively fi lled %calciteXRD = responding untreated values. This suggests that Σ Σ Σ (1) the Greater Green River Basin from the north [ Acalcite_peak/( Acalcite_peak + Adolomite_peak)]*100, the presence of organic matter has no measur- before spilling into the Piceance Creek Basin able effect on δ18O in these samples. One to the south. Sediment composition suggests where A represents the area under X-ray dif- sample, WB1-351, exhibits a measurable dif- that these deltaic sands are derived from lithic- fraction (XRD) peaks of the mineral phases ferential (0.7‰) between treated and untreated rich sediment sources, such as the Absaroka of interest. This relationship is a modifi ed ver- samples. However, data from this sample have Vol canic Province (e.g., Surdam and Stanley, sion of the method discussed by Royse et al. a high standard deviation, and the measured dif- 1980) or the Challis Volcanic Field (e.g., Smith (1971) using peak height. Although peak height ference could be due to a small signal and asso- et al., 2008), as shown in Figure 1. is shown to be more accurate than peak area ciated analytical issues. in that study, area is preferred in this case in METHODS order to better represent low, broad peaks gen- Mass-Balance Modeling erated by poorly crystalline material. Resulting

Sampling and Correlation %calciteXRD values are only semiquantitative A mass-balance model was designed to test but allow the defi nition of three generalized cat- the relative infl uence of environmental param- ≥ Samples for this study were collected from egories: calcite (%calciteXRD 80), mixed min- eters on isotopic composition in Lake Gosiute. three outcrop sections and three drill cores. eralogy (80 > %calciteXRD >20), and dolomite Boundary conditions for the model were estab- ≤ Section locations were chosen to represent (%calciteXRD 20). lished based on published estimates of tempera- depo sitional conditions over a wide area of the ture, precipitation, lake depth, and lake area. Greater Green River Basin (Fig. 1), but they Isotopic Analysis Other parameters, such as the δ18O and volume are also based on exposure quality. Initial cor- of infl uxes and outfl uxes, were defi ned to match relations were based on identifi able tuffs, the Isotopic analyses using a modifi cation of the lower LaClede δ18O data to model output in the Buff Marker Bed, and stromatolite beds. Cor- techniques of McCrea (1950) were performed initial balance-fi lled condition (Table 1). relations build on the work of Rhodes (2002) on a Finnigan Delta +XL mass spectrometer in The model equation (Table 1) modifi es an and use stratigraphic data from the Firehole the Stable Isotope Biogeochemistry Lab oratory initial isotopic composition with fi ve primary

Canyon section included in that work for sup- at Stanford University. CO2 for analysis was factors: precipitation within the immediate Lake port. The Green River Community College sec- evolved from reaction of carbonate powder with Gosiute drainage basin, infl ux of water from out- tion has uncertain relationships to the rest of the pure phosphoric acid at 72 °C in reaction ves- side the basin (i.e., river capture), evaporation measured sections because it lacks recognizable sels fl ushed with helium gas. Corrections have from the lake surface, evapotranspiration and stratigraphic markers (Fig. 3). not been made for variable acid fractionation evaporation from the basin catchment area, between calcite and dolomite, which may cause and out fl ow over a basin sill (Fig. 4). A lack of X-Ray Diffraction measured δ18O in samples containing dolomite data to constrain groundwater conditions and to be as much as 0.8‰ too high (Sharma and composition prevents the inclusion of a ground- Samples used in this study exhibit variable Clayton, 1965). Precision of analyses is ±0.2‰ water fl ux in this model. Likewise, values for in- carbonate mineralogy, ranging from pure cal- (1σ standard error) based on repeated analy- put parameters such as atmospheric vapor δ18O, cite to pure dolomite. Oxygen isotope fraction- ses of Stanford RHOM calcite standard and humidity, and wind are unconstrained, prevent- ation in calcite has been studied extensively MERCK calcium carbonate powder. Lower ing effective incorporation of back-condensation (O’Neil et al., 1969; Kim and O’Neil, 1997), precision (±0.3‰, 1σ) is assigned to one sam- in model calculations. However, these oversim- but fractionation during direct precipitation of ple run which produced abnormally variable plifi cations may be partly accounted for by the dolomite is poorly constrained at surface tem- standard values. Correction factors between initial choice of δ18O values for lake surface and peratures (Sharma and Clayton, 1965; Fritz and –0.65‰ and 0.70‰ (average 0.25%), based catchment precipitation. Because model fl ux val- Smith, 1970; Rosenbaum and Sheppard, 1986). on the deviation of the standard value aver age ues are rates, their effect is dependent upon the In order to evaluate the effect of mineralogical of each run from its expected value, have been time over which they are applied. For purposes

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TABLE 1. PREFERRED MODEL PARAMETER VALUES of these calculations, the observed isotopic shift (Vδ +Q δ t + Q δ t + Q δ t + 1000Q t(1+ (1/1.0092))) Final model δ = I PL PL PC PC R R E is constrained to within 200,000 years, so model equation LW (V + Q t + (Q t/1.0092)) O E calculations are performed using t = 200 ka. This retemaraPelbairaV V eula epyT timescale is long compared to residence time in T Temperature 19.6 ± 2.1 °C Boundary (Little Mountain fl ora: Wilf, 2000) the lake, so model output from both initial and P Annual rainfall 76.9 + 33.2/–23.2 cm/a Boundary modifi ed model conditions can effectively be (Little Mountain fl ora: Wilf, 2000) considered a steady-state value. 2 2 ≈ 2 AL Lake surface area 15,400 mi = 39,885 km 40,000 km Boundary (mapped Laney Area: Roehler, 1993) In the initial model condition, it is assumed d Lake depth 50 m, based on clinoform relief Boundary that Lake Gosiute exactly fi lled its basin, with (cf. Surdam and Stanley, 1980) no outfl ow (Fig. 4A). Precipitation on the lake α Water-calcite α = (106 (2.78)T −2 −3.39) 1000 Boundary W-C W −C e surface and runoff from the immediate drainage fractionation factor (O’Neil et al., 1969) area are balanced by evaporation from the lake δ Evaporation δ18O ⎡⎛ ⎤ Boundary EL 1 ⎞ ⎛ δ ⎞ surface, so evaporation rates are chosen to satisfy ⎢⎜ ⎟ ∗⎜ EL + 1⎟ − 1⎥ ∗1000 ⎝ α ⎠ ⎝ 1000 ⎠ ⎣ EV ⎦ QL + QC = QE. The initial isotopic composition of α = (2.0667+103 (0.4156)T −1+106 (−1.137)T −2 ) 1000 δ δ18 EV e lake water ( I) is based on average O of calcite (Majoube, 1971) in the lower LaClede Bed, and precipitation δ18O δ Outfl ow δ18O δ Boundary O LW values (δ and δ ) are defi ned to equalize the δ δ18 δ18 PC PL I Initial lacustrine O –2.38‰, based on lower LaClede O = 26.39‰ Boundary δ δ18 effects of intrabasinal infl uxes and evaporative PL Lake surface –12. 28‰, defi ned to balance initial ODefi ned precipitation δ18O 18O depletion so that the model output lake water δ Catchment δ –2‰ Defi ned δ18 δ PC PL O = I. Because catchment precipitation is δ18 precipitation O likely to fall at higher elevations and have lower δ River δ18O Velbaira fieD ned R δ18 δ t emiT ak002 fieD ned O than lake surface precipitation, PL is given a value 2‰ lower than δ . Evaporation δ18O V Lake volume 1/3 (AL ×)d fieD nedPC 2 AB Drainage basin area 110,000 km , based on modern basin-bounding uplifts Defi ned is a temperature-based function of lake-water E Catchment 7a/mc4 fieD ned18 C δ O (cf. Majoube, 1971) and does not account evaporation rate for the infl uence of other climatic factors such EL Lake surface 7a/mc5.9 fieD ned evaporation rate as wind speed or atmospheric humidity for the reasons discussed above with respect to back- QPL Lake surface rainfall P × AL Defi ned

QC Catchment infl ux (P–EC) × (AB –AL)fieD nedcondensation. Adjustment of parameters such as Q Extrabasinal infl elbairaVwo fieD ned R evaporation rate, precipitation rate, and tempera- Q Outfl ow (Q + Q + Q ) – Q Defi ned O PL C R F ture while holding other parameters constant in QE Lake surface EL × AL Defi ned evaporation rate this initial state allows testing of the sensitivity lake-water δ18O has to these factors. The second model condition differs from the initial state by inclusion of an extrabasinal river fl ux. Climatic and isotopic parameters from the A: Intrabasinal Inflow only initial condition are held constant at the pre- δ δ ferred values, and new infl ow (Q R) is exactly QPL PL, QE EL δ balanced by introduced outfl ow (QO) rather QC PC EC than adding measurable increases in lake depth or volume (Fig. 4B). The isotopic composition and volume of introduced infl ow are varied in δ δ δ −++ δ this second condition to produce a lake-water (V Q PLPLI Qt PCC Qt E EL t) = 0 δ18 δ18 LW =δ QO,QR O that matches the minimum calcite O ob- V served in the upper LaClede Bed. A version of B: Extrabasinal Inflow the model spreadsheet is included in GSA Data δ δ δ 1 QPL PL, QC PC, QE EL Repository Table DR2. δ δ QR R QO L EC 1GSA Data Repository item 2009190, DR1: Data repository table DR1 includes oxygen isotope data from Laney Member samples, with associated δ δ δ δ −+++ δ δ (V Q PLPLI Qt PCC Qt RR Qt E ELt Q– LO t) sample descriptions and mineralogy categorizations; LW =δ DR2: Data repository table DR2 shows the mass bal- V ance model spreadsheet tables in both values and equations formats; DR3: Data repository table DR3 Figure 4. Schematic representation of model fl uxes for (A) the initial records %calciteXRD values and the position and in- condition and (B) introduced extrabasinal infl ow. In the initial con- tensity of XRD peaks in Laney Member samples, dition, infl ow from the uplift-bounded basin is balanced by evapora- as well as indicating which peaks were included as tion keeping lake level at sill level. After an extrabasinal river infl ow calcite peaks (light gray) and dolomite peaks (dark gray) in %calciteXRD calculations, is available at is intro duced, the lake spills across the sill and outfl ow is added as a http://www.geosociety.org/pubs/ft2009.htm or by model parameter. request to [email protected].

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Apart from the initial lake-water δ18O value presence of preserved fi nely laminated facies formation regarding radiation balance, air tem- constrained by lower LaClede carbonate data, (cf. Olsen, 1990). Both catchment area and lake perature, and air moisture, modern lakes larger the best constraints on model conditions are depth estimates are highly uncertain but have than 500 km2 with precipitation between 70 and associated with paleobotanical annual tempera- virtually no effect on the model results (Fig. 5). 80 cm/a have actual evaporation ranges between ture and precipitation estimates from the Little Even in modern settings, evaporation is 26 and 73 cm/a (Fig. 6; Herdendorf, 1984). Al- Mountain fl ora (Wilf, 2000). Lake surface area is diffi cult to measure directly, and calculated though the reliability of this data is somewhat constrained by the maximum mapped area of the estimates can vary by as much as 8 cm/mo. uncertain, evaporation data from Williams Lake Laney Member (Roehler, 1993), and drainage depending on which equation is used (Winter in Minnesota, determined by energy budget cal- catchment area is based on the approximate area et al., 1995). However, two separate data sets culation methods (cf. Harbeck, 1958; Gunaji, surrounded by the basin-bounding drainage di- from modern lakes both suggest that model 1968), are likely to be more reliable and tell a vides in the modern Greater Green River Basin. evaporation rates are reasonable, and may even similar story. Seasonal daily evaporation rates Estimated lake depth is based on the preserved overestimate the effect of evaporation by using from this location calculated between 1982 and relief of delta-front clinoforms in the Sand Butte an initial precipitation-evaporation (P-E) ratio 1986 (Sturrock et al., 1992) would produce Bed (Surdam and Stanley, 1980) and on the <1. Based on compiled data calculated from in- between 79.9 and 102.9 cm/a evaporation, if they were applied to a 12-month open-water 12.00 season. This suggests slightly higher absolute evaporation rates than the Herdendorf data, P=175% but Williams Lake and nearby locations also E=25% Columbia receive >80 cm/a precipitation, and summer R. monthly P-E ratios suggested by these data are 10.00 25% P=150% E=50% also >1 (Sturrock et al., 1992). Due to interdependence between some model parameters within model calculations, model 50% 8.00 Max. Isotopic shift (7.9 ‰) output error is diffi cult to quantify. However, O 175% P=125% δ δ adjusting T, P, AL, AB, EC, EL, PC, and PL to

18 E=75% generate a range of Q versus δ18O curves for Missouri R. R Δ δ 150% δ Min. Isotopic shift (6.2 ‰) specifi c R values allows approximation of the 6.00 potential variability in model results. 75% Snake R. 125% P=110% RESULTS 175% E=90% 4.00

Lakewater Δ δ X-Ray Diffraction Mineralogy 150% 110% 90% Laney samples are dominated by calcite and 25% Colorado R. 2.00 dolomite mineral phases, but a strong quartz peak 125% is also present in almost all samples, and minor peaks from unidentifi ed subordinate phases are 175% 50% 110% 10% present in ~30% of the samples. Clays are not –10000% 0.00 identifi ed in XRD patterns from these samples. 50 110,000 40,000 19.6 76 75, P=76 Zero However, no attempts were made to prepare ori- 79.5 E=75,79.5 Inflow ented clay slides for XRD analysis (cf. Moore and Reynolds, 1997), and the range of 2θ angles represented by common clay peaks was not ana-

lyzed for most samples. The %calciteXRD deter- ) 2 minations show only weak correlation between 2 ) (cm/a) (cm/a) mineralogy and δ18O in these samples (Fig. 7A). (km Evaporation (km Precipitation In the lower LaClede Bed, only 32% of sam- Lake depth (m) Catchment area ples have %calciteXRD >80%, demonstrating Lake surfaceTemperature area (°C) inflow (cm/a) (cm/a) –17‰ extrabasinal that dolomitic or mixed-mineralogy carbonate dominates. This contrasts sharply with the upper LaClede Bed, where 73% of samples have >80% calcite based on XRD. See Table DR3 (footnote Evaporation + precipitation 1) for details of XRD results. Figure 5. Sensitivity of lacustrine carbonate δ18O to different model parameters. Preferred The most notable correlation between min- model values are given along the bottom axis, and percent values on the graph indicate the eralogy and δ18O (R2 = 0.7449) is restricted to percentage of those values that the labeled points represent. Horizontal lines showing a 6.2‰ samples below the Buff Marker Bed in the WB1 shift (the difference between lower LaClede average and upper LaClede average from cal- core (Fig. 7B). High δ18O in these samples cor- citic samples) and a 7.9‰ shift (the difference between the calcite lower LaClede average and responds with shallow-water lithology and low

the minimum observed upper LaClede value) are shown for reference. The effect of a –17‰ %calciteXRD (Fig. 2), suggesting that mineral- extrabasinal infl ow with discharge values similar to several modern rivers is also shown. ogical differences associated with facies change

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350 the assumption that this is generally true is made Fresh IP/E: model ratio in the context of in the interpretation of this data set. This change Brackish modern rainfall and evaporation coincides with a shift observed in 87Sr/86Sr re- ported by Rhodes (2002), where 87Sr/86Sr from Saline 300 the upper LaClede (μ = 0.71231, σ = 0.00041) Hypersaline is 0.00070 lower than 87Sr/86Sr from the lower μ σ Lake Gosiute model LaClede ( = 0.71161, = 0.00028). The lowest 87Sr/86Sr values observed in the upper LaClede 250 are 0.71091 and 0.71121, in the WB1 and WB2 cores, respectively, both minimums occurring within the δ18O shift (Fig. 3 and Table 2).

200 Model Results

Evaporation = precipitation Model results suggest that the isotopic com- position and volume of lake infl ows are signifi - 150 cant controls on lacustrine δ18O. In particular, the introduction of large extrabasinal infl ows with lower δ18O than catchment precipitation 18 Annual evaporation (cm) is capable of driving lacustrine δ O to lower 100 values equivalent to the shift observed in the LaClede Bed isotope record. If a captured river had δ18O ≤–24‰, a discharge ≥20 × 109 m3/a would be needed to produce the observed mag- 50 nitude of change. As noted in the discussion of model calculations, model output uncertainty is diffi cult to quantify. However, if the initial

assumption that intrabasinal infl ux QC + QPL = evaporative outfl ux (Q ) is incorrect, this esti- 0 E 0 50 100 150 200 250 300 350 mate may have uncertainty as large as +14 × 109 Annual precipitation (cm) or –12 × 109 m3/a. Larger volumes of water are needed to generate the isotopic shift with higher Figure 6. Comparison of model evaporation and precipitation parameters to annual actual riverine δ18O, and a discharge of ≥40 × 109 m3/a evaporation and annual precipitation from modern lakes greater than 500 km2 based on data (+11 × 109 or –16 × 109 m3/a) is required with from Herdendorf (1984). In most cases, precipitation is greater than measured evaporation an isotopic composition ≥–17‰ (Fig. 8A). even in saline lakes. This relationship indicates that it would be unreasonable to choose an These water volumes are consistent with a river initial model evaporation rate higher than the one used in this study (fi lled black box). between the size of the modern Colorado River and modern Missouri River, with low δ18O characteristic of a high-elevation water source. may account for small variations in δ18O in the positive excursion marked by dolomitic min- Model curves generated at 26 °C, rather than lower LaClede Bed. The slope of this correla- eralogy (Figs. 2 and 3). This isotopic event is the mean annual temperature of 19.6 °C, sug- tion line is –0.026, suggesting relative δ18O completed within less than 15 m above the fi ll gest smaller rivers would produce the observed fractionation between pure calcite and pure to spill surface. The highest observed oxygen isotopic change if carbonate formation took dolomite of ~2.6‰. Experimentally produced isotope values are found in the positive ex- place seasonally during warm summers. How- “proto dolomite” compositions show that under cursion at ~5 m above the boundary (30.3 ± ever, this effect is small for low δ18O infl ows, surfi cial conditions, oxygen isotopes of dolo- 0.2‰). Values of δ18O as low as 17.5 ± 0.3‰ so the minimum size of the river needed is only mite formed by disordered chemical precipi- are observed in the upper LaClede, and δ18O in reduced by a small volume (Fig. 8B). tation are expected to be 3‰–4‰ higher than the upper LaClede is consistently lower than in Changes in precipitation and/or evaporation calcite from the same waters (Fritz and Smith, the lower LaClede (Fig. 3). Calcite and dolo- are the only climatic parameters with capability 1970), which is consistent with this result. In- mite from the lower LaClede have an aver- in the model to independently produce change cluding WB1 lower LaClede samples above the age δ18O of 27.0‰ (Table 2), whereas upper equivalent to that observed in the Laney data Buff Marker Bed reduces the confi dence of this LaClede calcite and dolomite average 22.4‰ (Fig. 5). However, a 6.2‰ reduction in δ18O re- correlation (R2 = 0.3902), and no other section (4.6‰ more negative). quires at least a 33% reduction in lake-surface produces correlations with high confi dence. Calcite-only δ18O demonstrates an isotopic evaporation (~25 cm/a) under constant annual shift of –6.2‰, with a difference between aver- precipitation conditions. An increase in pre- Oxygen Isotope Composition age maximum and minimum δ18O of 7.0‰ cipitation of 50% or greater is needed in the ab- (Table 2). The consistency of δ18O trends at sence of reduced evaporation. In order to reach Oxygen isotope results show a distinct nega- different locations suggests that water in Lake the minimum measured lower LaClede value, a tive shift at the lower LaClede–upper LaClede Gosiute may have been homogenous with respect 50% or greater evaporation decrease (~40 cm/a) boundary, immediately preceded by a brief to δ18O during deposition of these sediments, and is needed, or an ~30% increase in precipitation

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35 ditions, and an absence of kerogen in the dolo- A y = –0.0435x + 27.901 micrite may refl ect oxidation due to subaerial 2 R = 0.3504 exposure (Surdam and Stanley, 1979). Corre- lation of dolomite content with facies change 30 also supports a playa-mudfl at mechanism for dolomite formation in the Green River Forma- tion (Wolfbauer and Surdam, 1974). If dolomite

O in the Laney Member had an allogenic playa 18 25 δ18 δ dolo micrite origin, its O composition may re- fl ect evaporative 18O enrichment in playa waters sepa rated from the main lake. 20 An alternative to the playa-lake model is biogenic Mg enrichment by blue-green algae in lake-bottom sediments, as has been sug- gested based on mineralogical data from the 15 Piceance Creek Basin (Desborough, 1978). 35 B y = –0.0292x + 29.089 According to this model, Mg concentrated R2 = 0.8301 into biomass by algae and released during or- ganic matter decay allows dolomite formation 30 in organic-rich carbonate sediments. Biogenic infl uence on lake-center dolomites is supported O

18 by cation substitution noted in profundal dolo- δ mite samples (Mason and Surdam, 1992), in 25 contrast to a lack of cation substitution in in- ferred mudfl at dolomites. Low-temperature formation of dolomite ce- 20 ments, fi lms, and travertines from fresh water, 0 102030405060708090100 2+ − attributed to interaction between Mg -HCO3 – % Calcite rich groundwaters and Ca2+-OH−–rich surface waters, have been documented in the Coast

Figure 7. Correlation of isotopic values with mineralogy, shown as %calciteXRD from X-ray Ranges of California (Barnes and O’Neil, 1971; diffraction. Plots shown are for (A) all δ18O data and (B) δ18O from the Washakie Basin O’Neil and Barnes, 1971). However, this model #1 core (WB1) below the Buff Marker Bed. The lower LaClede data from WB1 shown for dolomite formation is unlikely to explain here are the only data that demonstrate a notable covariant relationship between δ18O and widespread dolomicrite associated with the

%calcite0XRD. Green River Formation, because the Mg- and Ca-rich waters involved are associated with ultra mafi c rock types and serpentinization that combined with an ~30% decrease in evaporation. who argue that calcite precipitation resulting are not present in the Greater Green River Basin. Other combinations of reduced evaporation, in- from the periodic stratifi cation and/or mixing Because different formation mechanisms creased precipitation, increased temperature, and of fresh infl ows with more saline-alkaline lake with potentially different 18O/16O fractionations reduced lacustrine surface area are also capable water would explain laminated, kerogen-rich, may account for the formation of calcite and of causing isotopic change of this magnitude, but calcitic highstand deposits. dolomite observed in this record, it is important in all cases a dramatic increase in P-E (on the Dolomite formation in the Green River For- to note that the Laney δ18O shift is observed in order of 50% or more) is required. mation may be the result of different processes both dolomitic and calcitic samples. Despite the than calcite formation, and has been described generally similar values from stratigraphically DISCUSSION by two main competing models. In the Wilkins equivalent samples of different mineralogy, Peak Member, silt- to sand-sized particles trend assessments are focused on data obtained Origin of Laney Carbonate eroded from dolomitic surface crusts, caliches, from calcitic samples to minimize the infl uence and tufas have been proposed as the source of of potential mineralogy-based δ18O variation on Laminated, fi ne-grained calcitic carbonates dolomicrite in the lake, accounting for both the interpretations. such as those observed in the Laney Member are fi ne grain size of dolomitic carbonate and the texturally consistent with formation by primary presence of intraclasts in lacustrine sediment Late Diagenesis precipitation from surface waters. This mecha- (Smoot, 1983). Laney dolomite may form in a nism is one way that laminated carbonate sedi- similar way, where lowstand dolomicrites analo- Although partial silica replacement of car- ments in modern lakes are commonly formed gous to modern penecontemporaneous dolo mite bonate observed in stromatolitic and oolitic (e.g., Talbot, 1990), although laminations are formed in sabhka and supratidal settings could facies supports localized alteration in permeable only preserved in the absence of bioturbation. A have been transported and mixed with lacustrine facies of the Laney, it is not observed in micritic primary precipitation mechanism for the forma- calcite (Surdam and Stanley, 1979). Mud cracks sediments. Moreover, several points argue for tion of calcite in the Green River Formation is and intraclasts in the lower LaClede support this preservation of original or penecontemporane- also supported by Surdam and Stanley (1979), model by documenting periodic subaerial con- ous δ18O values in the Laney Member.

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TABLE 2. SUMMARY OF LANEY δ18O AND 87Sr/ 86Sr δ18O lower upper Δ Δ δ18O δ18O LaClede LaClede maximum- lower LaClede Section‡ maximum minimum μσ n μσ n minimum average-minimum Calcite data CCR 27.54 N/A 27.54 – 1 – – 0 N/A N/A GRCC 26.86 N/A 26.86 – 1 – – 0 N/A N/A GD 26.86 18.50 25.24 1.33 4 21.45 1.36 12 –8.37 –6.74 SB 26.25 19.93 26.25 – 1 22.08 1.04 12 –6.32 N/A WB1 27.07 20.97 26.63 0.55 14 22.24 1.01 15 –6.10 –5.66 WB2 27.61 20.87 26.38 1.07 9 22.56 1.09 20 –6.74 –5.50 Average 27.19 20.22 26.39 0.95 30 22.16 1.16 59 –6.96 –6.17 All data CCR 27.88 17.50 26.82 0.95 13 21.98 2.92 7 –10.38 –9.32 GRCC 29.13 26.52 27.56 1.23 5 – – 0 –2.61 N/A GD 28.39 18.50 25.72 2.1612 21.91 1.74 15 –9.89 –7.22 SB 30.11 19.93 27.38 1.5713 22.18 1.06 13 –10.18 –7.45 WB1 29.58 20.57 27.05 0.75 41 22.50 1.13 22 –9.01 –6.48 WB2 30.31 20.79 27.31 1.62 27 22.94 1.30 25 –9.52 –6.52 Average 29.49 19.94 27.00 1.39 111 22.43 1.52 82 –9.55 –7.06 87Sr/ 86Sr (data from Rhodes, 2002) lower upper Δ 87Sr/86Sr 87Sr/86Sr LaClede LaClede lower LaClede Δ Section‡ maximum minimum μσ n μσ n average-minimum average Calcite data WB1 0.71278 0.71091 0.71229 0.0003313 0.71159 0.00031 9 0.00138 0.00070 WB2 0.71244 0.71121 0.71189 0.000325 0.71158 0.00024 6 0.00068 0.00031 Average 0.71261 0.71106 0.71218 0.00037 18 0.71159 0.00028 15 0.00130 0.00059 All data WB1 0.71300 0.71091 0.71232 0.0003332 0.71164 0.00031 11 0.00141 0.00061 WB2 0.71373 0.71121 0.71227 0.0006210 0.71158 0.00022 7 0.00106 0.00068 Average 0.71336 0.71106 0.71231 0.00041 42 0.71161 0.00028 18 0.00130 0.00069 ‡Section abbreviations after Figure 1.

Because of the abundance of relatively Based on decreased δ18O of altered Eocene bi- creased temperature corresponds to higher rain- impermeable shales and mudstones in the valves, Morrill and Koch (2002) suggested that fall or snowfall δ18O (Rozanski et al., 1993) but Laney Member, it is unlikely that diagenetic diagenesis is a possible source of the very low can also lead to decreased fractionation during solutions penetrated pervasively after lithifi - micritic δ18O values observed by Norris et al. mineral precipitation and produce lower mineral cation (Wolfbauer and Surdam, 1974). Preser- (1996, 2000) in the Wilkins Peak Member. The δ18O (O’Neil et al., 1969). vation of 100-µ–scale lamination, visually best verifi cation of the primary nature of δ18O None of these factors, however, are adequate well-preserved fossil fi sh, and 87Sr/86Sr varia- in Laney micrites, however, comes from arago- to explain the shift observed in this record satis- tion with facies patterns in the lower LaClede nitic bivalve samples measured by Morrill and factorily. Although the facies change across the (Fig. 2; Rhodes et al., 2002) provide evidence Koch, collected north of Manila, Utah (Fig. 1). lower LaClede–upper LaClede boundary could that extensive postdepositional recrystallization These bivalves were taken from coquinas strati- be associated with increased precipitation, de- has probably not taken place. The fi ne-grained graphically equivalent to the upper LaClede creased evaporation, or both, the observed (clay and silt size) nature of calcite and dolo- Bed in the upper Laney (C. Morrill, 2005, iso topic shift would require a dramatic combi- mite particles, including iron-rich phases, is personal commun.), and preserved aragonitic nation of increased rainfall, reduced evapora- further evidence against diagenetic origin for mineralogy and annual growth layers demon- tion, and decreased temperature. δ18O values of the minerals (Surdam and Stanley, 1979). strate clearly that they refl ect primary deposi- calcite remain low through the entirety of the The δ18O data themselves also have several tional values. δ18O of the Manila bivalves falls upper LaClede Bed, suggesting that the isotope features that argue against diagenesis. Perme- between 19.54‰ and 25.56‰ (SMOW), with a shift was caused by a permanent change, and able stromatolite and grainstone facies, the mean of 23.13‰ and four samples below 22‰ no independent evidence has been reported for most likely carriers for diagenetic fl uids in the (Fig. 9). These low primary values suggest that concurrent rapid climate change in the western lower LaClede, tend toward high δ18O sug- a δ18O value of 22.16‰ (the observed upper United States. gesting that if diagenesis affects these rocks, LaClede average) is reasonable for upper Laney In contrast to the large δ18O shift observed in it causes high δ18O rather than the low dia- micritic carbonates. Laney micrites, an oxygen isotope record docu- genetic δ18O reported by Morrill and Koch menting the Paleocene–Eocene Thermal Maxi- (2002). The observed δ18O shift is progressive, Climate Change mum (PETM) in the Bighorn Basin shows less requiring that if diagenesis is responsible, it than 2‰ variation in oxygen isotopes from soil had a gradually increasing effect stratigraphi- Relative depletion in 18O can potentially carbonates (Koch et al., 2003). The PETM is cally upward. Furthermore, the same pattern be accounted for by many factors, such as associated with sea surface temperature warm- is observed in several sections across a wide cooler temperatures, increased rainfall and/or ing of as much as 8 °C (Zachos et al., 2001), area, with >100 km between the most widely decreased evaporation, increased elevation, while local warming estimates in the western separated (Fig. 1), indicating that diagenetic or change in dominant moisture source. This United States are 4–6 °C (Fricke et al., 1998; fl uids would have had to act uniformly over a multivariable system is further complicated by Fricke and Wing, 2004). If such a large climatic large lateral area. fractionation during mineral crystallization. In- change produced oxygen isotope variation of

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27 only ~2‰, climatic cooling or increase in rain- Pre-Capture Average A fall or snowfall required to produce a >5‰ shift Model curves calculated with a water would be expected to be extreme. Although temperature of 19.6 °C 25 the model does indicate that a large reduction δPL = –12.28, δPC = –14.28 in evaporation and/or increase in precipitation Model output values using δ18O and would be capable of producing the observed discharge values from modern rivers Missouri River 23 Columbia River shift, particularly in combination with de- Post-Capture Average (all) creased temperature (Fig. 5), this confl icts with the few applicable fl oral studies, which suggest either very slightly wetter conditions (Wing

(‰ SMOW) 21 and Greenwood, 1993) or even increasingly dry δ18 δ18 ORW = –12 conditions in the Middle Eocene (Leopold and

Calcite O RW = –14

O MacGinitie, 1972). 19 Post-Capture Minimum (GD) 18 Climatic explanations also fail to address the δ observed shift in Sr isotope values (Figs. 2 and δ18 3). Although δ18O is frequently infl uenced by 17 ORW = –16 climatic factors, Sr refl ects the crustal age and δ18 δ O 18 δ δ 18 RW composition of the rocks that interacted with O 18 = –18 O O 87 RW RW RW = –20 water entering the lake. Furthermore, Sr and = –24 = –22 15 86 Colorado River Yellowstone River Snake River Green River Sr do not undergo fractionation during evapo- ration, and therefore change in the Sr isotope 27 Pre-Capture Average ratio cannot be the result of decreased evapora- B Model curves calculated with a water tion from the lake surface. temperature of 26 °C Regional correlations based on dated tuffs 25 δPL = –10.39, δPC = –12.39 provide some of the strongest evidence against a climatic mechanism for hydrologic change in the Laney Member. Deposits of the Parachute 23 Creek Member in the Uinta Basin, contem- Post-Capture Average (all) poraneous with the overfi lled upper LaClede Bed, contain bedded evaporites suggestive of (‰ SMOW)

21 underfi lled conditions. This rules out a regional decrease in evaporative conditions at the time δ18O δ18 Calcite RW = –12 of the O shift (Smith et al., 2008), as might O 19 be expected from a regional climatic forcing

18 Post-Capture Minimum (GD) δ mechanism.

δ18O 17 RW = –14 Regional Uplift δ18 δ ORW 18 δ δ 18 δ 18 = –16 O 18 O O O RW Increasing the elevation of ranges surround- RW RW = –18 RW = –20 15 = –24 = –22 ing the Greater Green River Basin through rapid uplift would provide an explanation for decreased δ18O in the Green River Forma- 050100tion. However, the timing of the δ18O shift is Inflow (109 m3/a) constrained between 40Ar/39Ar ages of 48.94 ± 0.12 Ma from the underlying Analcite Tuff Figure 8. Model estimates of the isotopic composition and volume of extrabasinal infl ow (Smith et al., 2003) and 48.70 ± 0.19 Ma from required to produce the isotopic shift observed in the Laney calculated at 19.6 °C (A) and detrital feldspar of a sand in the overlying Sand 26 °C (B). Different lake-surface and catchment δ18O had to be used at different tempera- Butte Bed (Fig. 3; Smith et al., 2008). Detrital tures to preserve the initial δ18O value. Discharge values of modern rivers are plotted along grains were likely deposited rapidly after be- the x-axis for comparison, and show that capture of a river similar in size to the Snake or ing erupted, because the apparent age from the Missouri rivers would be capable of producing the shift observed in the Laney, if the iso- stratigraphically nearest overlying tuff, the Con- topic composition was low enough. A river the size of the Columbia River would be suffi cient tinental Peak Tuff (48.66 ± 0.28 Ma), is within to produce the isotopic shift even with relatively high δ18O. Where data were available, the error of this detrital age (Smith et al., 2008). δ18O and discharge of these modern rivers was used as model input, and the resulting values Based on an average modern global lapse rate of (open circles) are plotted for comparison to the model curves. Modern river δ18O values are 2.8‰/km (cf. Poage and Chamberlain, 2001), discharge-weighted average values of data from Coplen and Kendall (2000). River discharge ~2.2 km of uplift would be necessary to bring values are multiyear annual averages of streamfl ow data obtained from the U.S. Geological about a –6.2‰ change. Explaining the mag- Survey (USGS) Surface Water Daily Data database on August 29–30, 2007. These values nitude of change from maximum to minimum represent periods of record starting between 1910 and 1965 and ending in 2006 or early values (~–7‰) requires even greater surface up- 2007. Abbreviation: SMOW—standard mean ocean water. lift. To produce uplift of these magnitudes in the

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dence time in the reservoir increases evapora- tive enrichment of 18O in the water, this value might be even lower under natural conditions. Given the isotopic composition of the river, the discharge of the Columbia is more than double that needed to cause an isotopic shift such as the one in the upper Green River Formation (Fig. 8). The lack of detailed correlation between CCR 285.1 87Sr/86Sr and δ18O likely refl ects the drainage Unionid Bivalves of multiple high-elevation source terranes with (Morrill and Koch, 2002) variable 87Sr/86Sr. The Absarokas have char- 87 86 Average Average Laney Upper acteristically low whole-rock Sr/ Sr, with a Sr-concentration weighted average of 0.70549 Upper LaClede Unionid Bivalve LaClede Micrite (data from Peterman et al., 1970; Meen 15 17 19 21 23 25 27 29 and Eggler, 1987; Hiza, 1999; Feeley et al., 2002). Parts of the Idaho batholith west of the 18 δ O (SMOW) 87Sr/86Sr 0.706 isopleth also have low whole- rock 87Sr/86Sr (Criss and Fleck, 1987). These Figure 9. Comparison of calcite δ18O values from micrite in the upper LaClede Bed of the values would clearly contribute to decreased Laney (59 samples) to data collected by Morrill and Koch (2002) from aragonitic unionid 87Sr/86Sr as measured in Green River sedi- bivalves from the upper Laney Member north of Manila, Utah (15 samples). With the ex- ments (Rhodes et al., 2002). The contribution ception of one exceptionally light point from the Currant Creek Ridge core (CCR 285.1), of these 87Sr/86Sr lows is countered, however, all micrite values fall within the range of values shown by the unaltered aragonite. Average by the input of high and intermediate 87Sr/86Sr values for the two data sets (shown as vertical lines) are less than 1‰ apart. The low δ18O from other regional rocks. Precambrian base- aragonitic bivalves provide strong support for a primary origin of δ18O in calcitic Laney ment in the Wind River Mountains has an micrites. Abbreviation: SMOW—standard mean ocean water. average of 444 ppm Sr and a whole-rock Sr- concentration weighted average of 87Sr/86Sr = 0.72451 (data from Frost et al., 1998, n = 40), observed timeframe requires uplift rates on the Rb-poor, volcanics, or exposed rocks west of and Precambrian and plutonic rocks east of the order of 11 mm/a. For comparison, rapid uplift the 87Sr/86Sr 0.706 line in northwestern Idaho 87Sr/86Sr 0.706 line in central Idaho also have rates in the New Zealand Alps reach ~10 mm/a (e.g., Criss and Fleck, 1987). high 87Sr/86Sr (Criss and Fleck, 1987). Limited (Little et al., 2005). Such extreme rapid uplift Discharge-weighted average δ18O observed data from the Challis volcanics suggests that is inconsistent with waning Laramide activity in modern western U.S. rivers is rarely lower they have more intermediate Sr content and during deposition of the Laney, and also fails to than –17‰ (Fig. 8), implying that the seasonally 87Sr/86Sr signature (e.g., Norman and Leeman, explain the transition to overfi lled conditions in averaged composition of even high-elevation 1989). All of these locations are likely to have the lake. Increasing elevation by growth of vol- rivers is unlikely to be as low as –24‰. Con- contributed low δ18O water, but the Sr isotope canic topography could have been much more sequently, the captured river was probably signature expected from different relative con- rapid than tectonic uplift, and may have been larger than the minimum discharge that satis- tributions from these areas would vary widely. capable of producing isotopic change of large fi es the mass-balance model. Although large The eruption of large volumes of vol canic magnitude in the precipitation of the Challis or rivers commonly have higher δ18O than small material in the Absaroka and Challis areas Absaroka regions. tributaries due to the mixing of water from dif- undoubtedly affected existing regional drain- ferent parts of the catchment, rivers with dis- age networks. Cretaceous southeast-trending River Capture charges as great as 3000 m3/s (90 × 109 m3/a) paleovalleys within the Challis Volcanic Field are abundant in a survey of world rivers and (Janecke et al., 2000; Sears and Ryan, 2003) The capture of a low δ18O river provides a commonly carry water across distances of indicate eastward drainage from the Sevier fold mechanism for changes in both basin hydrology 1500 km or more (e.g., Petrere et al., 1998). The and thrust belt of Idaho. Sanidine 40Ar/39Ar ages and isotopic (δ18O and 87Sr/86Sr) composition. modern Columbia River is one such river that from quartzite-bearing ash-fl ow tuffs indicate Modeling results demonstrate the plausibility easily meets the high discharge–low δ18O re- that these valleys were infi lled between 49.51 ± of drainage capture as a mechanism for rapid quired by the model. At U.S. Geological Survey 0.14 Ma and 48.64 ± 0.12 Ma (M’Gonigle and isotopic change of this magnitude. The sug- (USGS) station 12472800 below Priest Rapids Dalrymple, 1996; Janecke et al., 1999; Janecke gested northern source of the Sand Butte Bed Dam, Washington, the Columbia had an annual et al., 2000). This implies that they were fi lled indicates that dominant drainage to the Greater average discharge of 106 × 109 m3/a between in at approximately the time of the inferred river Green River Basin during late deposition in 1917 and 2006 (Fig. 8). δ18O measured a few capture by Lake Gosiute. Water diverted from Lake Gosiute was from the north-northwest, miles downstream at station 12472900 ranges the drainages through these paleovalleys may where volcanic activity was resculpting the between –18.05‰ and –16.54‰ (Coplen and have contributed to a captured river ultimately landscape during the mid-Eocene. 87Sr/86Sr Kendall, 2000), with an average of –17.17‰ draining into Lake Gosiute. values decrease at the lower LaClede–upper given by weighting δ18O values according to Eocene conglomerates in the Bighorn, Wind LaClede boundary (Rhodes, 2002), also con- discharge from the upstream site on the date of River, and northern Green River basins also pro- sistent with introduction of drainage from measurement. Because these data are collected vide evidence of drainage sourced to the north areas with low 87Sr/86Sr bedrock such as young, downstream of a dam, where increased resi- and west, in the form of Paleozoic quartzite

248 Geological Society of America Bulletin, January/February 2010 Downloaded from gsabulletin.gsapubs.org on June 4, 2010 Geomorphic controls on lacustrine isotopic compositions clasts which have been tied to the Harebell- 87Sr/86Sr returns to slightly higher values (Figs. 2 Axelrod, D.I., 1998, The Eocene Thunder Mountain fl ora of Pinyon Conglomerates of the Jackson Hole area and 3). This may refl ect changing contributions central Idaho: University of California Publications in Geological Sciences 142, 61 p. (Love, 1939; Keefer, 1965; Dorr et al., 1977; from the low-moderate 87Sr/86Sr of places like Barnes, I., and O’Neil, J.R., 1971, Calcium-magnesium car- Krause, 1985). In the Wind River Formation the Challis and the Absarokas with more radio- bonate solid solutions from Holocene conglomerate cements and travertines in the Coast Range of Califor- near Du Noir, Wyoming, similar rounded Pre- genic Precambrian and plutonic rocks east of the nia: Geochimica et Cosmochimica Acta, v. 35, p. 699– cambrian quartzite clasts with pressure scars are 87Sr/86Sr = 0.706 line in Idaho. 718, doi: 10.1016/0016-7037(71)90068-8. associated with the earliest occurrence of vol- Beck, R.A., Vondra, C.F., Filkins, J.E., and Olander, J.D., 1988, Syntectonic sedimentation and Laramide base- canic pebbles (Keefer, 1957), suggesting asso- CONCLUSIONS ment thrusting, Cordilleran foreland: Timing of defor- ciation of Harebell-Pinyon type clasts with the ma tion, in Schmidt, C.J., and Perry, W.J., Jr., eds., early Absaroka volcanics. The most likely origi- Combined oxygen isotope geochemistry and Interaction of the Rocky Mountain foreland and the Cordilleran thrust belt: Geological Society of America nal source for these clasts is the Precambrian stratigraphic correlation support the conclusion Memoir 171, p. 465–487. Belt Supergroup in central Idaho (Lindsey, of Rhodes (2002) that a fundamental lake-type Blisniuk, P.M., and Stern, L.A., 2005, Stable isotope paleoaltimetry: A critical review: American Jour- 1972; Janecke et al., 2000), although conglom- change took place in the Laney Member of nal of Science, v. 305, p. 1033–1074, doi: 10.2475/ erates in the Bighorn, Green River, and Wind the Green River Formation between ~48.9 and ajs.305.10.1033. River basins may represent reworked Harebell- 48.7 Ma. This transition, from balance-fi lled to Bohacs, K.M., Grabowski, G.J., Jr., Carroll, A.R., and Miskell-Gerhardt, K.J., 2001, Non-marine sequence Pinyon sediments (e.g., Krause, 1985). Whether overfi lled conditions, is expressed by lithologic stratigraphy fi eld workshop; Guidebook to Wyoming directly deposited or reworked, these conglom- change and an isotopic shift toward lower δ18O and Colorado outcrops—2001, Rocky Mountain Sec- erates also indicate a history of drainage south values in lacustrine carbonate. The isotopic tion: SEPM (Society for Sedimentary Geology), 102 p. Boiseau, M., and Juillet-Leclerc, A., 1997, H2O2 treatment and eastward into the Laramide basins, consis- record of lake-type change provides a correla- of recent coral aragonite: Oxygen and carbon isotopic tent with the presence of a northwestern-sourced tion tool that can be applied where facies cycles implications: Chemical Geology, v. 143, p. 171–180, doi: 10.1016/S0009-2541(97)00112-5. river during the Eocene. are poorly expressed in the lower LaClede. Bradley, W.H., 1929, The varves and climate of the Green Rapidly building Absarokan topography dur- Capture of a low δ18O drainage network is River epoch: U.S. Geological Survey Professional ing Laney deposition would have presented a consistent with the isotopic record, lithologic Paper 158-E, 110 p. Bradley, W.H., 1964, Geology of the Green River and asso- considerable obstacle to east-fl owing drainage, record , and paleogeographic setting of the Laney ciated Eocene rocks in southwestern Wyoming and and an east-fl owing river would have been sus- Member, and provides the best explanation for adjacent parts of Colorado and Utah: U.S. Geological ceptible to capture by Lake Gosiute as volcanic the observed shift. Mass-balance modeling of Survey Professional Paper 496A, 86 p. Carroll, A.R., and Bohacs, K.M., 1999, Stratigraphic clas- debris blocked pathways to the Bighorn and Lake Gosiute also upholds this interpretation, sifi cation of ancient lakes: Balancing tectonic and cli- Wind River basins. Additionally, if some drain- demonstrating that capture of a large, low δ18O matic controls: Geology, v. 27, p. 99–102, doi: 10.1130/ 0091-7613(1999)027<0099:SCOALB>2.3.CO;2. age from central Idaho was already reaching the river could generate the observed change. Carroll, A.R., and Bohacs, K.M., 2001, Lake-type controls Greater Green River Basin during deposition of This study provides evidence that upstream on petroleum source rock potential in nonmarine ba- the lower LaClede, growing Challis topography geomorphic change can exert signifi cant control sins: American Association of Petroleum Geologists Bulletin, v. 85, p. 1033–1053. could have contributed to both increased water over the isotopic composition of downstream Carroll, A.R., Chetel, L.M., and Smith, M.E., 2006, Feast volume and decreased δ18O of this drainage. A waters, especially in settings with signifi cant to famine: Sediment supply control on Laramide ba- –7‰ to –10‰ δ18O shift is observed in paleo- topographic relief such as those where oxygen sin fi ll: Geology, v. 34, p. 197–200, doi: 10.1130/ G22148.1. sol, palustrine, and lacustrine carbonates from isotopes are used to infer orogenic histories. Carroll, A.R., Doebbert, A.C., Booth, A.L., Chamberlain, the Sage Creek Basin of southwestern Mon- Therefore, robust interpretations of other iso- C.P., Rhodes-Carson, M.K., Smith, M.E., Johnson, C.M., and Beard, B.L., 2008, Capture of high altitude tana and eastern Idaho between ~50 and 47 Ma topic records from continental settings must precipitation by a low altitude Eocene lake, west- (Kent-Corson et al., 2006), which may be evi- include an understanding of the associated ern U.S: Geology, v. 36, p. 791–794, doi: 10.1130/ dence for drainage reorganization northwest drainage histories of upstream catchments. G24783A.1. Chamberlain, C.P., Poage, M.A., Craw, D., and Reynolds, of the Absarokas. This implies that the major R.C., 1999, Topographic development of the Southern source of low δ18O water for both the Sage ACKNOWLEDGMENTS Alps recorded by the isotopic composition of authi- Creek and Greater Green River Basins was genic clay minerals, South Island, New Zealand: We are grateful to P. Brown, P. DeCelles, H. Fricke, Chemical Geology, v. 155, p. 279–294, doi: 10.1016/ likely in Idaho or western Montana. C. Kendall, M. Kent-Corson, C. Morrill, J. Valley, S0009-2541(98)00165-X. Water derived from high elevations in the P. Wilf, and H. Xu for discussions or data related Constenius, K.N., 1996, Late Paleogene extensional collapse to this paper. This manuscript benefi tted from help- of the Cordilleran foreland fold and thrust belt: Geo- Challis and Idaho Batholith regions provides logical Society of America Bulletin, v. 108, p. 20–39, δ18 ful review comments by K. Bohacs, E. Ito, and an explanation for the low O observed in M. Person. Reviews of an earlier related publication, doi: 10.1130/0016-7606(1996)108<0020:LPECOT> records from Laramide basins. Fluorite fl uid in- 2.3.CO;2. which were also helpful, were provided by S. Janecke, Coplen, T.B., and Kendall, C., 2000, Stable hydrogen and clusions from the Idaho mineral district, formed P. Koch, and E. Gierlowski-Kordesch. Financial sup- oxygen isotope ratios for selected sites of the U.S. Geo- at 51–50 Ma, trap meteoric water with δ18O of port was provided by ChevonTexaco, ConocoPhillips , logical Survey’s NASQAN and benchmark surface- –20‰ (Seal and Rye, 1993) providing direct the Donors of the Petroleum Research Fund of the water networks: U.S. Geological Survey Open-File American Chemical Society, American Association Report 00-160, 424 p. evidence of isotopically light water in central of Petroleum Geologists Grants-in-Aid, the Univer- Criss, R.E., and Fleck, R.J., 1987, Petrogenesis, geochronol- Idaho (Fig. 1). Criss and Taylor (1983) also ogy, and hydrothermal systems of the northern Idaho sity of Wisconsin Department of Geology and Geo- 18 16 δ18 physics, and National Science Foundation grants batholith and adjacent areas based on O/ O, D/H, suggest that extensive interaction with low O 87Sr/86Sr, K-Ar, and 40Ar/39Ar studies, in Vallier, T.L. (~–16‰) hydrothermal waters of meteoric ori- EAR-0230123 and EAR- 0114055 to A.R. Carroll and and Brooks, H.C., eds., Geology of the Blue Mountains EAR-0609649 to C.P. Chamberlain. gin took place within Idaho Batholith plutons region of Oregon, Idaho, and Washington: The Idaho batholith and its border zone: U.S. Geological Survey between 49 and 37 Ma. REFERENCES CITED Professional Paper 1436, p. 95–137. 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