Lacustrine Sedimentology, Stratigraphy and Stable Isotope Geochemistry of the Tipton Member of the Green River Formation

Abstract

The Tipton Member of the Green River Formation occupies much of the Greater Green River Basin (GGRB) of Wyoming and Colorado. Long hypothesized to record a single shift from open to partly closed hydrology, new detailed stratigraphy and stable isotope geochemistry indicates that its strata record open, then partly closed, then open, then partly closed hydrology, which are each recorded by distinct transitions in facies associations, geochemistry, carbonate mineralogy, and organic content. Intervals of open hydrology occur coincident with the progradation of deltaic sandstones that are absent during the partly closed intervals, suggesting that environmental transitions were controlled by avulsions of the Idaho River. The first of these transitions occurs at the contact between the Scheggs bed and overlying Rife bed, and is thought to reflect the initial impoundment of Lake Gosiute. The Scheggs bed ranges from 23.5 to 36.5 m, and is characterized by fluvial-lacustrine lithofacies, calcitic mineralogy, an average Fischer Assay content of 7.6 gal./ton, and low δ¹⁸O and δ¹³C values (25.3‰ and 0.7‰, respectively). These deposits transition over a five meter interval to the overlying 2–15 m-thick lower Rife bed. The lower Rife bed is characterized by fluctuating profundal lithofacies, dolomitic mineralogy, an average Fischer Assay content of 17.6 gal./ton, and high δ¹⁸O and δ¹³C values (29.3 and 5.3‰). The lower Rife bed transitions up-section over a two meter interval into fluvial-lacustrine lithofacies of the 2.5–20 m thick middle Rife bed, which exhibits calcitic mineralogy, an average Fischer Assay content of 9.7 gal./ton, and low δ¹⁸O and δ¹³C values (23.0 and 1.9‰). The third and final transition, from the middle Rife bed to the upper Rife bed, occurs gradationally over 6 m of section. The 6.5–22 m-thick upper Rife bed is characterized by fluctuating profundal deposits, dolomitic mineralogy, an average Fischer Assay content of 19.2 gal./ton, and high δ¹⁸O and δ¹³C values (29.8‰ and 8.5‰, respectfully). We interpret this succession of abrupt changes in lithofacies and isotope geochemistry within the Tipton Member to reflect the diversion, recapture, and ultimate diversion of a major source(s) of water and sediment into the basin.

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3.1 Introduction

The Green River Formation (GRF) in the Greater Green River Basin (GGRB) records the dynamic evolution of Eocene Lake Gosiute as it transitioned from originally open paleohydrology to closed and then finally back to open paleohydrology before its final infilling by alluvium (Fig. 3.1) (Carroll and Bohacs 1999; Smith et al. 2008). The Tipton Member, previously the Tipton Shale Member (Schultz 1920), contains a broad array of non-marine facies. It underlies the evaporative Wilkins Peak Member, and has been proposed to record a transition from a hydrologically open to hydrologically closed lake basin (Roehler 1993; Pietras et al. 2003). Despite a long history of field and core-based investigation and documentation (Pipiringos 1955; Schultz 1920; Roehler 1992; Oriel 1961), the detailed stratigraphy and depositional controls of Tipton Member remain incompletely understood. This study represents a basin-scale examination of the lithofacies, stratigraphic packaging, and stable isotope geochemistry of the Tipton Member along White Mountain in the eastern Bridger subbasin of the GGRB. We argue that the Tipton Member records two compositionally and isotopically distinct transitions from overfilled to balanced-fill conditions.

Fig. 3.1

Map showing the location of Eocene basins, basin-bounding uplifts, and measured sections (Base map modified from Witkind and Grose (1972). Abbreviations for Laramide uplifts: TG Teton-Gros Ventre, OC Owl Creek, MB Medicine Bow, SM Medicine Bow)

The Tipton Member was deposited within the GGRB, which is located in the foreland of the Sevier fold and thrust belt, and is bounded by Precambrian-cored, Laramide-style uplifts to the north, south and east (Fig. 3.1) (Dickinson et al. 1988; DeCelles 1994). Basement-involved arches divide the GGRB into the Washakie, Sand Wash, Great Divide, and Green River sub-basins, all of which contain Tipton Member strata. Fluvial deposits of the Luman Member of the GRF and Niland Tongue of the Wasatch Formation underlie the Scheggs bed, whereas evaporative facies of the Wilkins Peak Member overly the Rife bed (Fig. 3.2). The Tipton Member has been divided into two beds: the Scheggs bed and overlying Rife bed (Roehler 1991b), the contact between which is understood to reflect a transition from freshwater, overfilled conditions to more saline, balanced-filled conditions within the lacustrine system (Roehler 1993; Carroll and Bohacs 1999). The Farson Sandstone, an arkosic deltaic complex, laterally bounds and interfingers with the Tipton Member in the northern part of the basin (Roehler 1992), and is equivalent to coarse-grained alluvial strata of the Pass Peak Formation in the northwest GGRB (Smith et al. 2008; Steidtmann 1969). Principal Tipton Member lithologies include variably organic-rich calcareous mudstone and marlstone, fossil-bearing siltstone, ostracode and oolitic grainstone, stromatolite, and various sandstone lithofacies assigned to the Farson Sandstone. Its fossil assemblages include freshwater bivalves and gastropods, fish, plant fragments, and vertebrate and invertebrate trace fossils. The Tipton Member is Early Eocene in age, based on late Wasatchian faunas found within it and in equivalent strata, and sanidine ⁴⁰Ar/³⁹Ar ages for ash beds contained within it and the overlying Wilkins Peak Member (Smith et al. 2008). Assuming that the Scheggs bed represents 2/3 of the amount of time of the Lysitean mammalian subage, its deposition occurred over approximately 1.06 ± 0.43 m.y., (Smith et al. 2008, 2010). Based solely on their relative thicknesses, the Rife bed represents 0.60 ± 0.31 m.y. of deposition, whereas the Scheggs bed represents 0.46 ± 0.30 m.y.

Fig. 3.2

Stratigraphy and associated lake-types of the Green River Formation in the Greater Green River Basin, Wyoming

3.2 Methods

3.2.1 Stratigraphic Analysis

Stratigraphic sections were chosen along a 150 km north–south cross-section through the GRB to document Lake Gosiute’s evolution during deposition of the Tipton Member (Fig. 3.1). Five field sections were measured at decimeter-scale along the western and southwestern flanks of the Rock Springs Uplift along White Mountain and Flaming Gorge Reservoir. From the south to the north, they are Firehole Canyon (FC), Villa Lane (VL), Spring Mound (SM), White Mountain Petroglyphs (WMP), and Boar’s Tusk (BT). An additional field section, Whitehorse Creek (WC), was modified from Pietras (2003) and incorporated into this study. Three cores supplement field data: the U.S. DOE/LETC Currant Creek Ridge No. 1 (CCR), U.S. ERDA White Mountain No. 1 (WM) and the Union Pacific Railroad Blue Rim 44-19 (BR). Locations of core and field sections are shown in Fig. 3.1.

3.2.2 Sampling

Field samples were collected primarily for comparative petrophysical analysis, and were not used for XRD and stable isotopic analysis due to observable effects of weathering. Samples from two cores, CCR and WM, were collected for XRD and stable isotopic analysis. The sample spacing and sample density of each core was concentrated on the Scheggs-Rife contact, as determined by lithofacies assemblages.

3.2.3 Mineralogy

The carbonate mineralogy of profundal mudstone samples from two cores (CCR and WM) was assessed using XRD analysis. Each sample was ground into fine powder, and a split of which was analyzed at the University of Wisconsin-Madison using a Sintag PAD V X-ray diffractometer using a Cu Kα x-ray source (λ = 1.5418 Å). The scans of 54 samples were run between 20 and 55 degrees 2-θ, a range that captures all relevant calcite and dolomite peaks. Step size was set at 0.02 degrees, with a step time of 1 s. The calcite-dolomite proportion of each sample was determined using the relative areas of dominant calcite peaks (those between 29.30 and 30.00 degrees 2-θ) and dolomite peaks (those between 30.40 and 31.10 degrees 2-θ) according to the relationship:

$$ {\%}_{calcite}=\left[\frac{{\displaystyle \sum }{A}_{calcite\_ peak}}{\left({\displaystyle \sum }{A}_{calcite\_ peak}+{\displaystyle \sum }{A}_{dolomite\_ peak}\right)}\right]*100 $$

Based on this calculation, samples were categorized as dominantly calcitic (>80 % calcite), mixed (20–80 % calcite), or dominantly dolomitic (<20 % calcite).

3.2.4 Stable Isotopes

An additional split of each powdered sample was allocated for geochemical analysis. Samples were processed at the University of Michigan Stable Isotope Laboratory in Ann Arbor, Michigan where they were roasted in vacuo at 200 °C for 1 h to remove contaminants. Samples were then reacted at 77° ± 1 °C with 4 drops of anhydrous phosphoric acid for 8 min (12 min for dolomitic samples) in a Finnigan MAT Kiel IV preparation device coupled directly to the inlet of a Finnigan MAT 253 triple collector isotope ratio mass spectrometer. Maintained measured precision using this methodology is reported as better than 0.1 % for both carbon and oxygen isotope compositions. In this study, δ¹⁸O data is reported relative to VSMOW, while δ¹³C data is reported relative to VPDB. Where duplicate samples were run, averages weighted according to standard deviation were used.

3.3 Lithofacies

This study recognizes 13 lithofacies within the Tipton Member, as defined by lithology, organic content, sedimentary structures, biologic markers and paleo-flow indicators (Table 3.1).

Table 3.1 Tipton Member Lithofacies

Microlaminated kerogen-rich mudstone

Laminated, organic-rich mudstone is the dominant lithofacies of the Rife bed, though it is also found to a lesser extent in the Scheggs bed. The mudstone is generally dark-brown to black in color. Discrete, rhythmic, variegated laminations are μm to mm in scale, and planar parallel (Fig. 3.4a). Interlaminations (mm- to cm-scale) of tan siltstone, tuff, and chert are found infrequently within sections. Fischer Assay oil yields range between 20 and 28 gal./ton (Roehler 1991a), qualifying these rocks as high-quality oil shale (Culbertson et al. 1980; Dana and Smith 1972). Well-preserved, intact fish fossils are present though infrequent. Occasional phosphatic concretions are preserved as both thin (mm-scale), lenticular bodies along lamination planes as well as irregularly shaped nodules around which overlying laminations are deformed. When associated with thin interbeds of sand, stromatolite, and ostracode and oolitic grainstone, cm-scale mud cracks are observed. In outcrop, this lithofacies forms pronounced cliffs that can be traced laterally for tens of kilometers. Field sections weather blue in color, and are colloquially referred to as “blue beds”.

Interpretation: Preservation of fine, densely spaced laminations suggests deposition from suspension in an area beneath wave base and where bottom currents were continuously slow or non-existent. The high rate of organic preservation, reflected by high Fischer Assay oil yields, is interpreted to reflect deposition within low-oxygenated or entirely anoxic water conditions (e.g. Demaison and Moore 1980). However, it is unresolved whether permanent chemical and thermal stratification of lake waters is necessary for generation of kerogen-rich laminated mudstone (i.e., oil shale). Citing the presence of fossil catfish within oil shale lithofacies of the Laney Member of the GRF, Buchheim and Surdam (1977) suggested that only fluctuating or semi-permanent lake stratification is necessary for oil shale deposition within the GRF. An alternative interpretation of the presence of bottom-dwelling catfish within oil shale, however, is that the corpse bloated, floated out to deeper areas of the lake where it sunk and then came to rest along anoxic bottom waters.

The small phosphatic nodules around which overlying laminae conform are interpreted as coprolites, an observation also reported in profundal sediments of the Tipton Member (Castro 1962) and the Laney Member (Fischer and Roberts 1991). This facies is interpreted to represent profundal lacustrine deposition within a low-oxygenated, stratified lake conditions similar to those found within profundal zones of modern lakes Zurich (Bradley 1929; Kelts and Hsü 1978) and Tanganyika (Huc et al. 1990).

Laminated mudstone

This lithofacies is represented in both the Scheggs and Rife beds as grey to brown mudstone. In most examples, variegated laminations range between mm to cm in scale (Fig. 3.3b). Upper and lower contacts of distinct laminae are variable, occurring as discrete planar, diffuse/gradational planar, or wavy/irregular. Infrequent mm- to cm-scale interlaminations of grey-tan silt and tuff are also present. Within the mudstone and siltstone inter-laminations, well-preserved fish fossils, ostracode molds, and compacted, vertical burrows are present. Tan, spherical nodules of silt onto which overlying laminae conform are observed with variable frequency. Fischer Assay oil yields range between 9 and 22 gal./ton (Roehler 1991a), classifying this lithofacies as low-quality oil shale (Culbertson et al. 1980; Dana and Smith 1972). Outcrop expression ranges from covered, gradual slopes to moderately high-angle cliffs.

Fig. 3.3

Photographs of Tipton Member lithofacies: (a) Thinly laminated (μm to mm), organic-rich mudstone; (b) variably laminated (μm to cm) mudstone; (c) silicified Goniobasis tenera gastropods, which constitute the Goniobasis Marker bed; (d) freshwater bivalves preserved in the fossil-bearing siltstone lithofacies; (e) Gilbert-type foresets observed from the WMP section; (f) climbing ripples resulting from super-critical flow with both stoss- and lee-sides preserved; (g) injection feature where vF sand is injected into overlying F sand in the BR core; (h) concentric delamination of stromatolites which mark the Tipton-Wilkins Peak contact; and (i) brecciated stromatolites at the SM field section

Interpretation: The relatively diminished Fischer Assay oil yield of this lithofacies compared to that of the mm-laminated, organic-rich mudstone lithofacies suggests either diminished organic production, an increase in organic destruction and/or clastic dilution. The primary source of oil-shale kerogen in the GRF is autochthonous algae and bacteria (Tissot and Vandenbroucke 1983; Horsfield et al. 1994; Carroll and Bohacs 2001; Bohacs et al. 2000). During periods of sustained freshwater input, a corresponding increase in available oxygen may have increased degradation of these particulate organics (Horsfield et al. 1994), while increased inorganic sedimentation may have diluted preserved organic concentrations (cf. Carroll 1998). Alternatively, freshwater input has been attributed to the reduction of dissolved bicarbonate concentrations within a lake system, which ultimately decreases primary productivity (Horsfield et al. 1994). Collectively accounting for each possibility, this lithofacies is interpreted to reflect deposition during periods of freshwater input and low chemical and thermal stratification.

Massive kerogen-rich mudstone

Massive, organic-rich mudstone is present in the Rife bed and, less frequently, in the Scheggs bed. This mudstone is typically dark-brown to dark grey in color. Laminations are not visible in hand samples or thin section, but may exist cryptically. Silty interbeds are rare, and thin (mm-scale) interlaminations of bitumen-saturated tuff and horizontal fracture-fills of dolomite are frequently observed. Fischer Assay yields range between 22 and 26 gal./ton (Roehler 1991a), qualifying this mudstone as high-quality oil shale (Culbertson et al. 1980; Dana and Smith 1972). Both core and field samples have a bituminous odor and are absent of well-preserved burrows, ostracode molds and fish fossils. Outcrop expression of this facies is pronounced cliffs that can be traced laterally for several kilometers.

Interpretation: Like the laminated organic-rich mudstone, above, preserved organic matter suggests sub-mixolimnium deposition in a chemically and thermally stratified lacustrine system in which there was an insufficient alternation in the delivery of micrite to create lamination. Massive kerogen-rich mudstone beds could also have been deposited by hypopycnal plumes basinward of sites of fluvial input or created via entrainment of organic rich mud during storm events (cf. Renaut and Gierlowski-Kordesch 2010).

Massive mudstone

Massive mudstone is the dominant lithofacies of the Scheggs bed. It is almost exclusively light to medium grey in color, has no visible lamination, and generally preserves largely intact freshwater animals such as gastropods (2–4 cm) and bivalves (4–8 cm) (Fig. 3.3c and d), both of which commonly exhibit abrasion and removal of shell ornamentation. Ostracodes are abundant throughout, whereas vertical burrows occur infrequently. In the lower Scheggs bed, gastropods and bivalves are commonly silicified. Towards the upper Scheggs bed, however, gastropods become less frequent and articulated bivalves retain original aragonite mineralogy. The Fischer Assay oil yield of this lithofacies is low, ranging between 2 and 6 gal./ton (Roehler 1991a). Deep trenching (~0.5 m) did not reach un-weathered, intact section due to vegetated slopes. Abundant silicified float is, therefore, the primary facies identifier in the field. It should be noted that the observation of aragonitic fossils in the field was limited to the Villa Lane section, where a recent road cut had generated fresh exposure.

Interpretation: The absence of lamination within this lithofacies suggests either continuous sedimentation (Pasierbiewicz and Kotlarczyk 1997), or more likely the bioturbation of formerly laminated mud (Fischer and Roberts 1991; Demaison and Moore 1980) within a non-stratified lacustrine system (Bradley 1929, 1931; Carroll 1998). It may also reflect deposition by hyperpycnal plumes delivered to the lake center during periods of significant fluvial input or shoreface agitation by waves. The decreased kerogen content observed within this lithofacies is thought to reflect algal degradation resulting from increased oxygenation and decreased bicarbonate due to downstream outflow of lakewaters (Horsfield et al. 1994). This study interprets the massive mudstone lithofacies as littoral deposits within a freshwater system, which is consistent with the observation by Surdam and Stanley (1979) of bioturbated, mollusk and ostracode-bearing mudstone with freshwater deposits of the upper Laney Member of the GRF, as well as the observation of Cohen (1989) of abundant gastropod infauna within the littoral to sub-littoral zones of modern Lake Tanganyika.

Fossiliferous siltstone

Fossil-bearing siltstone is found in both profundal and marginal deposits. It is medium grey or tan in color and is either laminated or massive. Laminated intervals display mm- to cm-scale laminations of varied shades of brown silt and mudstone. Contacts among laminations occur as planar-parallel, wavy-parallel, and wavy non-parallel. Mollusc and ostracode fossils are most commonly associated with planar-parallel and massive interbeds. Where thicker than approximately 0.5 m in outcrop, this lithofacies typically constitutes low-angle, vegetated slopes that require trenching (<0.25 m) for stratigraphic observation. Exposure quality of thin interbeds of this lithofacies (<0.5 m) is highly variable and largely dependent upon the resistance of overlying lithologies.

Interpretation: This lithofacies is found within a variety of depositional and hydrodynamic environments. Where siltstone contains abundant gastropods that are continuously distributed throughout the bed, this lithofacies is interpreted to represent littoral deposition. Similarly dense gastropod concentrations are observed in the littoral margins of Lake Tanganika (Cohen 1989). Where this lithofacies occurs as thin, fining-upward interbeds within mudstone lithofacies, it is interpreted as turbidite deposits. In association with stacked, coarsening upward successions of sandstone and mudstone, the siltstone lithofacies is interpreted as pro-delta and distal bar deposits. Similar pro-delta deposits are observed in the Laney Member of the GRF (Stanley and Surdam 1978), in Late Pleistocene Lake Bonneville (Lemons and Chan 1999), and in Jurassic deposits within the East Gobi Basin, Mongolia (Johnson and Graham 2004).

Delta foreset sandstone

Within this study area, steeply dipping delta foreset lithofacies are limited to the Farson Sandstone Member at the WMP and WC (Pietras 2003) field sections. At WMP this lithofacies constitutes approximately 4.5 m of vertical section. Foresets dip 28 degrees (Fig. 3.3e) and are composed of micaceous, biotite-rich, vF-F, sub-angular sandstone that is well-cemented by calcite. Relatively large (up to 4 cm) rip-ups of underlying tan siltstone are observed along foreset planes, while twig and reed impressions are found less frequently. Irregular, 10–20 cm loading features typify the contact of this facies with underlying, convoluted siltstone. The upper portion of the steeply dipping foreset lithofacies is typically abruptly overlain by mudstone.

Interpretation: The steeply dipping foresets of this lithofacies are interpreted as delta front deposits. A high flow regime, capable of transporting terrigenous plant material basin-ward and eroding cohesive mud and siltstone, is thought to have existed during deposition. Soft-sediment deformation within underlying units suggests this lithofacies was deposited rapidly across a soft substrate, while bed thickness indicates water depth was at least 4.5 m. Accounting for compaction effects, foreset height and interpreted water depth may have been greater. Larger (25 m) Gilbert-type foreset beds are observed in Eocene Fossil Lake (Buchheim and Eugster 1998) and within Upper Laney Member of the Green River Formation (Stanley and Surdam 1978). Modern analogues include deltaic deposits in freshwater lakes Constance (Müller 1966), Pyramid (Born 1972) and Malawi (Scholz et al. 1993).

Trough cross-bedded sandstone

This lithofacies is a dominant component of the Farson Sandstone and is observed in both field and core sections. Beds are 10 cm–4 m thick and are composed of vertically aggraded sets that range between 5 cm and 25 cm in thickness and approximately 15–40 cm in width. Individual mm- to cm-scale bedding planes consist of grain-supported, angular to sub-angular, vF to M sand. Bounding surfaces between sets curve approximately 25 degrees at their steepest, and are tangential to underlying erosion surfaces. Laminae are often highlighted by concentrations of biotite and muscovite. No visibly identifiable floral or faunal remnants are preserved. Outcrop relief of this lithofacies is highly variable and dependent upon localized weathering effects.

Interpretation: Trough cross-beds occur in a variety of depositional environments, where flow is sufficient for the down-flow migration of dune structures (Allen 1962; Harms and Fahnestock 1965). In association with stacked sequences of laminated mudstone, siltstone and sandstone, this lithofacies is interpreted to represent delta slope and channel deposits. Similar trough cross-bedded deposits are documented in the Bitter Creek deltaic complex of the Laney Member (Stanley and Surdam 1978), as well in Jurassic-Cretaceous lacustrine deltas in the East Gobi Basin, Mongolia (Johnson and Graham 2004).

Horizontally-bedded sandstone

Horizontally bedded sand is exclusive to the Farson Sandstone and is observed in both core and field sections. Packages are between 20 cm and 2 m thick, and are composed of 2–5 mm thick, horizontal to near horizontal (less than 2 degree dip), vertically aggregated beds of vF sand, which are distinguished by muscovite- and biotite-rich interlaminations. Infrequently, vertical burrow structures, reed and twig imprints, and cm-scale interbedded tuffs are observed.

Interpretation: Horizontally laminated sandstone within the Farson Sandstone suggests deposition within high flow regimes (Harms and Fahnestock 1965; Bridge 1978; Paola et al. 1989; Arnott and Southard 1990; Cheel 1990; Arnott 1993). Typically found at the top of coarsening-upward packages of littoral mudstone, siltstone and sandstone, we study interprets horizontal laminations to have been deposited by shallow, high energy unidirectional flow and/or bidirectional wave-generated currents in a swash zone along upper delta-front and delta margins.

Hummocky cross-stratified sandstone

Micro-hummocky cross-bedded, vF sand is limited to Farson Sandstone deposits and is observed in both core (BR) and field (WMP) sections. Individual laminations are approximately 1–3 mm and dip approximately 15 degrees. Distinguished by mica-rich laminations, bed sets range between 3 and 6 cm thick, while aggregated packages of bed sets range between 10 and 25 cm in thickness. In the WMP field section, hummocks entrain abundant phosphatic fish “debris” composed of broken ribs, scales, platelets and vertebrae. Outcrop exposure of this lithofacies is highly variable and largely dependent upon the resistance of overlying lithofacies.

Interpretation: Hummocks result from a combination of oscillatory and weak unidirectional flow (Nøttvedt and Kreisa 1987; Leckie and Krystinik 1989; Arnott and Southard 1990; Dumas et al. 2005) in an environment where sedimentation rates are high and water depth low enough to facilitate large, fast waves, yet deep enough to maintain the oscillatory waves and unidirectional currents (Dumas and Arnott 2006). In association with stacked successions of mudstone, siltstone and sandstone lithofacies, hummocks are interpreted to have formed in the littoral zone under combined-flow between fairweather and storm wave base, likely in upper delta-front and lower delta-platform environments.

Climbing-rippled sandstone

This lithofacies is exclusive to the Farson Sandstone and is observed in both core and field section. Beds are between 0.3 and 1.5 m thick, and are composed of 2 mm to 4 cm thick vertically aggraded, variably undulating laminations of vF sand that are made distinct by intermittent silt-, muscovite- and biotite-rich laminations (Fig. 3.3f). Ripple limbs exhibit irregular sinuosity, dipping between 15 and 30 degrees on either side of the ripple crest. Unlike the planar parallel and hummocky sand lithofacies with which this lithofacies is commonly associated, organic matter and fossils are not observed. Outcrop expression is highly variable and largely dependent upon the resistance of overlying lithofacies.

Interpretation: Preservation of both the lee and stoss limbs of each ripple crest indicate deposition in an area having high rates of sedimentation from suspension and bed load (Jopling and Walker 1968; Allen 1982). Variability of inclination among the ripple limbs of this lithofacies is thought to be a function of small changes in the hydrodynamic environment and/or variable sediment discharge. As part of the Farson Sandstone, this lithofacies is interpreted to represent deposition along the delta-front, an area known for high-rates of sedimentation and wave-action modification thereof (McLane 1995).

Massive sandstone

The massive sandstone lithofacies is exclusive to the Farson Sandstone, and is observed in both core and field sections observed as beds of vF to mL, grain-supported sandstone that occur in 10 cm–4 m intervals. Silt and mud rip-ups, fish debris, burrows preserved by differential cementation, and terrestrial plant material such as reeds and root casts are commonly observed throughout the beds. Internal architecture, however, is absent. Underlying siltstone frequently exhibits scour and convolute bedding due to loading of overlying sandstone. Grain-size within the sandstone beds is variable, ranging from vF to mL. Transitions between sandstone grain-size are often, but not always, abrupt. Where contacts are sharp, load-induced injection, or flame structures, are observed (Fig. 3.3g).

Interpretation: Lack of internal architecture within this lithofacies is interpreted to result from either liquefaction related to delta slump processes or bioturbation along the delta-top or delta shore-face. Where underlying sediments exhibit deformation, slumping is thought to have occurred, while burrows, though limitedly preserved, indicate bioturbation along shore-face or delta top environments. Stanley and Surdam (1978) observed similar massive units along the shore-face of the Bitter Creek deltaic complex in the Laney Member.

Stromatolite

Though algal stromatolites are present in both the Rife and Scheggs beds, distinct morphological properties and depositional environments differentiate the two occurrences. Stromatolite beds within the Rife bed are 10–30 cm thick, weather orange-brown by concentric spalling (Fig. 3.3h), and are laterally extensive across several kilometers. Here, stromatolite beds are associated with a transition from organic-rich calcareous mudstone to green, mud-cracked, evaporite mineral-bearing siltstone. As such, this facies is interpreted as the base of the Tipton-Wilkins Peak contact. The stromatolite lithofacies in the Scheggs bed is exclusive to the Spring Mound (SM) field section and is laterally bound approximately 20 m to the north and south by two large (12 m high × 20 m wide) spring deposits with which it inter-fingers. The stromatolite package itself constitutes 40 cm of vertical section. The upper 30 cm of this is a breccia consisting of pebble- to gravel-sized clasts of broken algal material, shale and silt rip-ups (Fig. 3.3i).

Interpretation: Modern lacustrine stromatolite beds are observed along many lake margins and adjacent fluvial systems, such as Green Lake (Eggleston and Dean 1976), the Great Salt Lake (Halley 1976), and Lake Tanganyika (Cohen and Thouin 1987; Cassanova and Hillaire-Marcel 1992). Within the Rife bed, stromatolite beds are inferred to have formed near the shore, similar to those found in the Laney Member (Wolfbauer and Surdam 1974; Roehler 1993; Rhodes 2002). However, in the absence of an observable shore-parallel transect within the GGRB, neither this study nor Laney Member studies can confirm whether stromatolite beds are linear, shore-parallel features. Previous studies have not reported stromatolite beds in the Scheggs bed. It is likely that the high rates of hydrologic influx, corresponding shoreline transgressions and the chemical parameters of an overfilled lacustrine system prevented extensive growth. The structural high and local calcium ion-rich geochemistry of the spring mounds are interpreted to have enabled stromatolite growth and preservation at the Spring Mound field section. From the morphology and preservation characteristics, we infer that stromatolite beds within the Scheggs bed were initially protected by adjacent mound structures, but increased wave action resulting from lake expansion re-worked and brecciated the algal mats.

Ostracode and ooid grainstone

Ostracode and oolitic grainstone lithofacies are found in both the Scheggs and Rife beds. Medium to coarse grain-sized, undeformed ostracodes and oolites are preserved as horizontal laminations (2 mm–7 cm) and within vertical burrows. In both occurrences, ostracodes and oolites are associated with fossil-bearing siltstone, cm-laminated mudstone, and massive sandstone facies. Entrained within the grainstones are small (mm- to sub-mm) silt rip-ups and phosphatic fish debris, including scales, ribs, vertebrae and articulated jaws. The lithologic texture of the grainstone is generally friable in outcrop and loosely consolidated in core.

Interpretation: Though ostracodes are found throughout many lacustrine sub-environments (Cohen 2003), accumulation and formation of skeletal debris into a grainstone requires hydrologic sorting and/or minimal clastic dilution, suggestive of wave agitation along shorelines far from fluvial deltas. Oolites are similarly observed along wave-agitated lake margins (Cohen and Thouin 1987; Talbot and Allen 1996; Balch et al. 2005) where groundwater and lake water mixing leads to precipitation of inorganic carbonate (Wolfbauer and Surdam 1974; Kelts and Hsü 1978).

3.4 Lithofacies Assemblages and Associations

Two distinct lithofacies assemblages comprise the Tipton Member: fluvial-lacustrine and fluctuating profundal, both of which have distinct basin-ward and shoreward associations (Table 3.2; Fig. 3.4). Together, these lithofacies assemblages define four stratigraphic zones within the Tipton Member. The fluvial-lacustrine lithofacies assemblage defines the Scheggs bed (Zone A), while both fluctuating-profundal (Zones B and D) and fluvial-lacustrine assemblages (Zone C) characterize the overlying Rife bed (Fig. 3.2). These stratigraphic zones correspond to distinct zones of carbonate mineralogy as determined by XRD analysis. Profundal mudstone of fluvial-lacustrine Zones A and C is dominantly calcitic (69 % and 80 % calcite, respectively), while profundal mudstone of fluctuating-profundal Zones B and D is dominantly dolomitic (30 % and 12 % calcite, respectively).

Table 3.2 Tipton Member Lithofacies Associations

Fig. 3.4

Schematic interpretive cross-section of Lake Gosiute during (a) fluvial-lacustrine and (b) fluctuating profundal deposition

3.4.1 Fluvial-Lacustrine Lithofacies Assemblage

Basin-ward association

The dominant component of this association is the massive mudstone lithofacies, which bears abundant bivalves and gastropods (Fig. 3.5a). Interlaminations of fining upward beds (3–60 cm thick) of fossil-bearing siltstone lithofacies, which also preserve broken freshwater taxa, are frequent. Within the bottom-most cored meters of the CCR core, both massive and laminated organic-rich mudstone lithofacies contain interbeds of coquina. These interbeds range between 1 and 75 cm thick and contain bivalve and gastropod shell fragments within an organic-rich mudstone matrix. While the majority of these interbeds preserve primary aragonite, the largest interbed (75 cm) is has noticeably less mudstone matrix and is calcified throughout.

Fig. 3.5

Representative stratigraphic sections from the four principle lithofacies associations: The fluvial-lacustrine lithofacies assemblage has distinct (a) basin-ward and (b) shoreward associations. The basin-ward association is typified by the massive mudstone lithofacies and punctuated by fining upward interbeds of fossil-bearing siltstone lithofacies. Freshwater fauna, such as gastropods and bivalves are abundant throughout. The shoreward association is typified by stacked, coarsening upwards sequences of siltstone and sandstone lithofacies. The fluctuating profundal lithofacies assemblage also has distinct (c) basin-ward and (d) shoreward associations. The basin-ward association is typified by oscillations between thinly laminated, organic rich mudstone and variably laminated mudstone lithofacies. These oscillations are also apparent in Fischer Assay. The shoreward association is typified by a diverse array of fine-clastic, coarse-clastic, and biogenic lithofacies. Mud-cracked horizons are also observed. Well-preserved fish are common throughout both associations

The basin-ward association of this assemblage reflects sustained water and sediment input into a hydrologically open, non-stratified lake system. Abundantly preserved salt-sensitive fauna (e.g. bivalves) indicate well-oxygenated, freshwater conditions within the lake, while the absence of mudstone lamination and tuff interbeds suggest a combination of bioturbation, continuous sedimentation, and mixing within the suspended sediment column during deposition. Decreased Fischer Assay oil yield within fine-grained sediments is also evidence of sustained water and sediment input. Where fluvial systems enter a lacustrine basin, the concentration of particulate organics is reduced by clastic dilution, chemical degradation due to well oxygenated source waters, biologic consumption, or a combination thereof (Huc et al. 1990; Horsfield et al. 1994; Bohacs 1998). The effects of these processes are assumed to be greater in areas more proximal to source water input. The lateral expanse (>50 km of the study area) to which fine-grained profundal sediments of this association are depleted of kerogen is interpreted to reflect basinwide oxygenation and increased water and sediment input to the basin. Moreover, stacked 10–100 cm alternations between organic-rich and organic-depleted mudstone are not observed in this assemblage, which we interpret to reflect sustained hydrologic input and corresponding lake-level high-stand (Bohacs et al. 2000). Horsfield et al. (1994) observe similar lack of thin parasequences within fluvial-lacustrine deposits of the underlying Luman Tongue.

Organic-rich mudstone in Zone A of the CCR core is thought to reflect deeper, less-oxygenated waters beyond the effects of clastic dilution. At this location, Fischer Assay oil yield is greater (10.8 gal./ton) than that observed among fine-grained lithofacies located more proximal to the deltaic complex of the Farson Sandstone (e.g. 4.5 gal./ton at WM). Roehler (1992) for example interprets the CCR core location as one of the deepest areas of Lake Gosiute. Because the organic-rich mudstone lithofacies suggests anoxic or low-oxygenated bottom waters, interbeds of broken aerobic fauna are interpreted as turbidite deposits. Gravity currents, similar to those described in Kneller and Buckee (2000), are thought to have transported these fauna basin-ward from shoreline and littoral habitats. In this manner, finer-grained sediment entrained in the current, such as silt and littoral mud, would have been deposited further basin-ward. Up-section within CCR core, small beds (5–15 cm) of fining-upward silt are interpreted as distal deposits of gravity currents.

Shoreward association

Coarse-clastic deposits of the Farson Sandstone typify the shoreward association of the fluvial-lacustrine assemblage. Concentrated to the northern-most portion of the basin, stacked, coarsening-upwards successions of laminated mudstone, siltstone and sandstone lithofacies (Fig. 3.5b) are interpreted to represent a broad deltaic sequence, which grades basin-ward to laminated profundal mudstone. Mudstone lithofacies represent pro-delta deposits, siltstone lithofacies the distal bar, and sand lithofacies delta front and delta top deposits. Overall thickness of these successions decreases basin-ward, while the number of sequences and ratio of fine-grained sediments to sand increases. Load-induced contacts (e.g. Fig. 3.3g) are found between the gastropod-bearing siltstone and overlying trough cross-bedded, massive and Gilbert-type foreset lithofacies. Contacts between siltstone and high-flow regime planar parallel and wavy parallel lithofacies, however, are conformable. Similar lacustrine deltaic sequences have been described in the Laney Member (Stanley and Surdam 1978), Lake Bonneville (Lemons and Chan 1999) and in the East Gobi Basin (Johnson and Graham 2004). The progradational and aggradational geometries of the sandstones suggest sustained sediment influx, while rip-ups and irregular lithofacies contacts indicate infrequent storm generated high-flow and erosive hydrologic conditions.

3.4.2 Fluctuating Profundal Lithofacies Assemblage

Basin-ward association

The association is characterized predominantly by kerogen-rich, commonly fish fossil-bearing mudstone lithofacies (Fig. 3.5c), with minor interlaminations of fossiliferous siltstone and, less commonly, ostracode and ooid grainstone. The increased preservation of intact fish fossils and absence of benthos within this assemblage relative to that observed within the fluvial-lacustrine assemblage indicates increased chemical and thermal stratification and the presence of anoxic water bottoms. Moreover, the absence of bivalves suggests lake waters of this assemblage were more saline than those of the fluvial-lacustrine assemblage, where bivalves were abundant.

In detail, basin-ward association of the fluctuating profundal assemblage is characterized by alternating intervals of both massive and microlaminated kerogen-rich mudstone and less organic-rich mudstone lithofacies (Fig. 3.5c). These facies alternations are apparent in Fischer Assay logs of oil yield, and are interpreted to represent oscillations between low- and high-stand conditions within a hydrologically closed lake system. Specifically, laminated, organic-rich and massive organic-rich mudstone correspond to high-stand lake conditions, when lake level was at a maximum and clastic dilution at a minimum. During these conditions, cold, freshwater input flows over denser, more saline resident lake waters amplified the chemical and thermal stratification within the lake system. The corresponding increase in oxygenation and nutrient delivery results in increased production of particulate organics, which were preserved in the anoxic bottom waters. Less organic-rich mudstone intervals in turn correspond to low-stand lake conditions, when water input is decreased relative to high-stand conditions. Stratification, as well as oxygen and nutrient delivery was correspondingly suppressed by a reduction of freshwater input, resulting in a decrease in both the production and ultimate preservation of particulate organics.

Shoreward association

A diverse array of successive, interbedded fine- and coarse-clastic sediments within the Rife bed typifies the shoreward association of the fluctuating profundal assemblage (Fig. 3.5d). Laminated organic-rich mudstone, massive organic-rich mudstone and, to a lesser extent, variably laminated mudstone lithofacies are interlaminated by thin (<20 cm) packages of ostracode and oolitic grainstone, stromatolite, and massive sandstone lithofacies. Meter-scale intervals of massive sandstone and fossil-bearing siltstone are observed in the north central to northern-most field sections and are thought to represent closer proximity to lake margins. In both the WMP and BT field sections, stromatolites overlie large packages of fish and ostracode-bearing fossiliferous siltstone. At BT these stromatolites constitute a 2 m interval. Woody debris, burrows, ostracode molds, and fish debris are common throughout all facies within this association.

Like the basin-ward expression of this assemblage, the absence of freshwater fauna and progradational geometries suggests a relative reduction in sediment and water input when compared to the fluvial-lacustrine assemblage. Facies diversity among this association is also thought to reflect periods of high- and low-stands within the lake system, with fine-grained and bio-clastic sediments reflecting low-stand deposits and pulses of coarse-grained clastics representative of shifts towards high-stand conditions. These oscillations indicate rapid, basin-wide alteration in sediment deposition, organic production, and stratification within the lake-system

3.5 Basin-Scale Stratigraphy

Figure 3.6 illustrates the basin-scale packaging of the Tipton Member, which records three vertical oscillations between fluvial-lacustrine and fluctuating profundal lithofacies. Deltaic sandstone bodies (Farson Sandstone) accumulated at the northern edge of the basin during fluvial-lacustrine intervals, but are largely absent within fluctuating profundal intervals. Within the Tipton Member, transitions between lithofacies assemblages are gradational, particularly in basin-ward locations. Because the basin-ward expressions of both assemblages share multiple lithofacies, transitions lack distinct lithologic contrast. Basin-ward transitions are, therefore, constrained primarily on the basis of relative organic-content (fluctuating profundal lithofacies exhibit higher Fischer Assay oil yield) and the presence or absence of freshwater bivalves. The appearance of fossil fish at the base of the Laney Member indicates the transition from a hydrologically closed lake system to a partly open system with fresh surface water (Carroll and Bohacs 1999; Rhodes 2002). Fish remains are preserved within all lake-type assemblages of the Tipton Member, and similarly indicate that conditions, at least the uppermost lake water, were fresh. It should be noted, however, that intact fossils are more common in the fluctuating profundal assemblage, while fish “debris”, a slurry of bones, scales, and fins, is most common, though not restricted to, the fluvial-lacustrine assemblage. Though lithologic transitions also appear gradational at basin margins, they are marked by stronger lithologic contrasts. The contact from fluvial-lacustrine to fluctuating profundal assemblage, for example, is placed at the top of large bodies of prograding, coarse-clastic lithofacies and the initiation of profundal mudstone deposition.

Fig. 3.6

Basin-scale cross-section of the Tipton Member west of the Rock Springs Uplift in the Green River Basin, Wyoming. From south (A) to north (A′), these sections are Currant Creek Ridge Core No. 1 (CCR), Firehole Canyon (FC), White Mountain Core No. 1 (WM1), Villa Lane (VL), UPRR Blue Rim 44-19 core (BR), Spring Mound (SM), White Mountain Petroglyphs (WMP), Boar’s Tusk (BT), and Whitehorse Creek (WC). [Fischer Assay data modified from Roehler (1991a), and Whitehorse Creek (WC) field section modified from Pietras (2003)]. See Walker (2008) for complete stratigraphic dataset

3.5.1 Correlations

Correlation of the upper and lower limits of the Tipton Member was established using lithofacies assemblages, and is supported by Fischer Assay (Roehler 1991a), geochemical, and XRD analyses of CCR and WM cores (Fig. 3.6). Due to limited core and outcrop in the southern area of the basin, direct observation of the Wasatch Formation-Tipton Member contact was limited to field sections in the northern part of the basin (VL, SM, WMP, BT, and WC). Where observed, the contact is defined where oxidized (red) and reduced (green) fluvial plain siltstone and mudstone of the Wasatch Formation transition to gastropod-bearing mudstone and siltstone lithofacies of the Scheggs bed. This fossiliferous horizon is often referred to in literature as the Goniobasis Marker bed, as the horizon contains abundant Goniobasis tenera gastropods (Hanley 1976).

Previous analysis of the Tipton Member (Roehler 1991b) defined the Scheggs-Rife contact at the transition from freshwater, calcitic deposits to saline, dolomitic deposits. This study adheres to that previous terminology, placing the Scheggs-Rife contact at the first transition from a fluvial-lacustrine lithofacies assemblage (freshwater) to a fluctuating-profundal assemblage (more saline). This stratigraphic placement of the Scheggs-Rife contact coincides with a 56 % average decrease in calcite content and a gradual increase in δ¹⁸O and δ¹³C values (Fig. 3.6).

In the northern part of the basin (Fig. 3.6), the top of the Rife bed is characterized by stromatolite lithofacies (.25–1.5 m thick). These beds are conformably overlain by green, marcesite- and pyrite-bearing, mud-cracked mudstone and siltstone lithofacies of the Wilkins Peak Member. In the central and southern part of the GGRB, the Tipton Member-Wilkins Peak Member contact is defined where mm- and cm-laminated mudstone lithofacies of the Rife bed gradually transition to evaporative lithofacies of the Wilkins Peak Member. Within this study, the basin-ward and shore-ward expressions of the Tipton Member-Wilkins Peak Member contact are conformable, and coincide with a subtle isotopic trend towards lighter δ¹⁸O and δ¹³C values, as well as a sharp reduction in Fischer Assay oil yield, which is sustained at 0 gal./ton over a 1 m interval (Roehler 1991a). The conformable interpretation of the Tipton Member-Wilkins Peak Member contact presented by this study is somewhat discordant with that presented by Pietras et al. (2003), who propose that a major sequence boundary defines the contact. We argue instead that the Tipton-Wilkins Peak contact records a sharp change in water chemistry and brief subaeriel exposure of the edges of the basin, but not necessarily a basin-wide major lacuna.

3.6 Stable Isotope Analysis

Oxygen isotopes

Stable isotopic analysis of profundal mudstone from two cores, CCR and WM, distinguish four distinct δ¹⁸O zones within the Tipton Member (Table 3.3). These intervals directly correlate with zones defined by lithofacies assemblages. Fluvial-lacustrine Zones A and C have light δ¹⁸O values (25.3‰ and 23.0‰, respectively), while fluctuating profundal Zones B and D exhibit the heaviest δ¹⁸O signature within the Tipton Member (29.7‰ and 29.8‰, respectively). The δ¹⁸O transition from Zone A to Zone B (Shift 1) occurs gradually over 4 m within both CCR and WM cores. The transition from Zone B to Zone C (Shift 2), however, is abrupt, occurring over a 2 m interval. Because Zones C and D are under-sampled in both CCR and WM cores, the transition interval of Shift 3 cannot accurately be assessed.

Table 3.3 Average Zonal Tipton Member Geochemistry, mineralogy and oil yield

Carbon isotopes

The δ¹³C signature of the Tipton Member exhibits four distinct zones, all of which are coincident and directly correlative with those observed within the δ¹⁸O profiles of the same cores (Fig. 3.6; Table 3.3). Like the δ¹⁸O profile of the Tipton Member, δ¹³C is heaviest in fluctuating profundal Zones B and D (5.3‰ and 8.5‰, respectively). Fluvial-lacustrine Zones A and C have relatively lighter δ¹³C values at 0.6‰ and 1.9‰, respectively. Because carbon isotopic data was obtained from the same samples as δ¹⁸O values, the isotopic transition intervals and trends of the δ¹³C signature within the Tipton Member is the same as that of the δ¹⁸O profile.

δ¹⁸O and δ ¹³C covariance

Among profundal mudstone samples of the Tipton Member, correlation between δ¹⁸O and δ¹³C is generally moderate. Correlation is notably stronger within all samples of the CCR core (R² = 0.7292, n = 26) than it is among those within the WM core (Fig. 3.7a). δ¹⁸O and δ¹³C correlation among samples sharing distinct carbonate mineralogy is also strongest in the CCR core (Fig. 3.7b). Dolomitic samples within the CCR core, for example, exhibit a stronger correlation (R² = 0.9062, n = 6) than dolomitic samples within the WM core (R² = 0.0250, n = 10). Moreover, while dolomitic samples have the strongest δ¹⁸O and δ¹³C correlation compared to other carbonate mineralogies within CCR core, calcitic samples are the most strongly correlated within the WM core (R² = 0.3122, n = 8). This suggests that carbonate mineralogy does not influence δ¹⁸O and δ¹³C correlation in a predictable way across the basin. Among both core, fluctuating profundal zones exhibit stronger overall δ¹⁸O and δ¹³C correlation (R² = 0.2392, n = 18) than do fluvial-lacustrine zones (R² = 0.0835, n = 33) (Fig. 3.7c).

Fig. 3.7

The δ¹⁸O and δ¹³C covariance among litho-stratigraphic intervals of (a) CCR and WM cores, (b) carbonate mineralogy of samples within CCR and WM cores, and (c) within each litho-stratigraphic zone

3.7 Discussion

3.7.1 Stratigraphic Evolution

Major changes in Tipton Member lacustrine lithofacies were closely associated with initiations and cessations of Farson Sandstone deposition, suggesting that changes of fluvial influx into the basin were a primary driver. Within fluvial-lacustrine Zones A and C, progradational geometries along the lake margin indicate significant sediment and, thus, fluvial input into a hydrologically open (i.e. overflowing) lake basin. The Farson Sandstone is not observed within Zones B and D, indicating a relative reduction in sediment/fluvial input during fluctuating profundal deposition. With continuing tectonic subsidence in the basin and the absence of the high sediment input, accommodation was increased in fluctuating profundal Zones B and D. Reflecting this, lake margins transgressed and reach their maximum extent (Fig. 3.6). Additionally, prograding stratal geometries of fluvial-lacustrine intervals are replaced by aggradation of sediment in Zones B and D, indicating a transition to a hydrologically closed (i.e. impounded) lake basin. Carroll and Bohacs (1999) define three lake types, overfilled, balanced-filled, and under-filled, that are controlled by the amount of water and sediment input relative to basin accommodation. Using this lake-type characterization, fluvial-lacustrine intervals of the Tipton Member are classified as overfilled, while fluctuating profundal intervals are balanced-filled.

Distinct zones of carbonate mineralogy and Fischer Assay oil yield coincide with the lithologically defined intervals. Fluvial-lacustrine intervals A and C are relatively calcitic, exhibiting 75 % calcite on average. These intervals are generally poor in organic matter, with a weighted average Fischer Assay oil yield of 8.7 gal./ton, suggesting that well-oxygenated, unstratified lake conditions were present during deposition. Low oil yields can result from clastic dilution resulting from sustained fluvial input of siliciclastic detritus into an open lake system (Bohacs et al. 2000; Carroll and Bohacs 2001).

In contrast, fluctuating profundal intervals are relatively dolomitic, exhibiting 21 % calcite on average. The presence of dolomite within the Tipton Member can be explained via the biogenic model of Desborough (1978), whereby Mg is preferentially concentrated along lake bottoms as blue-green algae anaerobically decompose. Supporting this model is relatively high Fischer Assay oil yield (18 gal./ton) among profundal deposits of fluctuating profundal intervals. Alternatively, the playa lake model is commonly applied to the GRF to explain dolomite generation and distribution (Eugster and Surdam 1973; Mason and Surdam 1992). However, its application to the Tipton Member is inappropriate due to a lack of playa indicators (e.g. evaporite mineral casts, large scale desiccation features).

3.7.2 Isotopic Evolution

Diagenesis

Though post-lithification diagenesis can alter stable isotopic composition (Morrill and Koch 2002), our study observes several characteristics in both the lithology and geography of the study area that suggest the isotopic signature of the Tipton Member reflects the composition of Lake Gosiute during deposition.

Because aragonite is highly susceptible to diagenic alteration, its presence or absence can be used as a proxy for diagenic evaluation. Aragonitic bivalves were recovered from Zone A in the CCR core. Preserved growth bands within these shells, which would have been destroyed by diagenic alteration, suggest diagenesis had not occurred within surrounding sediments. Moreover, the mean δ¹⁸O and δ¹³C values of these bivalves (21.98‰ and −3.86‰, respectively) are lower than other light values recorded throughout the Tipton Member. Because aragonitic samples are interpreted to have been unaltered, similarly light δ¹⁸O and δ¹³C values within the Tipton Member are believed to reflect primary depositional conditions.

Profundal mudstone of the Tipton Member is relatively fine-grained and impermeable, a characteristic which inhibits the pervasiveness of diagenic solutions. Coquina interbeds within the CCR core best exhibit this impermeability. Preservation of primary aragonite is limited to matrix-supported interbeds, whereas clast- (i.e. shell-) supported interbeds are calcified. This suggests profundal mudstone is an effective shield against diagenic solutions.

The spatial distribution of isotopic trends themselves also suggests that alteration of δ¹⁸O and δ¹³C values is unlikely. Specifically, the same isotopic pattern is observed in CCR and WM cores, which are separated by over 65 km, a distance unlikely to be transversed by a homogenous diagenic solution. In order to create four, isotopically distinct zones, diagenic solutions would have had to occur with variable intensities along multiple, vertically spaced horizons. Furthermore, if diagenesis were responsible for each zone, the shift from isotopically heavy Zone B to isotopically light Zone C (Shift 2) would indicate more intense diagenic alteration up-section, an event that is highly unlikely.

Mineralogical effects on δ¹⁸O

Because carbonate minerals exhibit differential fractionation at surface temperatures (Sharma and Clayton 1965; Fritz and Smith 1970; Rosenbaum and Sheppard 1986), sample mineralogy can also compromise δ¹⁸O analysis. Dolomite, for example, is inferred to have a 3‰ heavier δ¹⁸O value than calcite precipitated from the same water (Fritz and Smith 1970). The isotopic shifts within the Tipton Member involve calcitic, mixed, and dolomitic mineralogy. As such it is possible that differential fractionation among distinct carbonate mineralogies may influence the δ¹⁸O signatures within the Tipton Member. However, in several shifts (i.e. Shift 1, Shift 2), the trends towards heavier or lighter isotopic values are preserved within the individual trend of each carbonate classification (e.g. calcite, mixed, and dolomitic). Therefore, while mineralogy may influence the magnitude of each isotopic shift, this observation suggests that shifts within the Tipton Member reflect major changes in lake water chemistry during deposition. Furthermore, covariance among mineralogy and δ¹⁸O results is insignificant among all carbonate mineralogies, with R² correlation values of 0.1043 (dominantly dolomitic samples), 0.2146 (dominantly calcitic samples), and 0.4263 (mixed samples) (Fig. 3.8).

Fig. 3.8

Comparison of mineralogy and δ¹⁸O composition within the Tipton Member. Calcitic samples exhibit the lowest δ¹⁸O values, while dolomitic samples exhibit the highest δ¹⁸O values

Effects on δ¹⁸O and δ¹³C and their implications

The variable δ¹⁸O signature observed within the Tipton Member is likely related to the combined effects of continued evaporation and variable residence time of lake waters. Evaporation preferentially removes δ¹⁶O from surface waters, thereby concentrating δ¹⁸O within a lake system. Increased residence time of lake waters within a lake basin compounds these evaporative effects, resulting in heavier δ¹⁸O values. Because lake water within a hydrologically closed lake system experiences longer residence time and, thus, exhibits heavier δ¹⁸O values, the heavy isotopic signature of fluctuating profundal zones B and D of the Tipton Member are interpreted to reflect closed lake conditions. Lighter δ¹⁸O values of Zones A and C indicate open lake conditions during fluvial-lacustrine deposition.

The δ¹³C signature of lacustrine deposits is generally more complex, as it is related to the rate of primary productivity (Kirby et al. 2002), the rate of organic decomposition within the lake system (Pitman 1996), and dissolved inorganic carbon (DIC) input from Phanerozoic limestone in Lake Gosiute’s catchment (δ¹³C ≈ 0). Photosynthetic algae preferentially remove δ¹²C throughout their lifetime, thereby enriching lake waters in δ¹³C if they are buried prior to decomposition. In this way, higher inorganic δ¹³C values reflect periods of abundant productivity, while lower δ¹³C values indicate depositional periods that are less conducive to organic production. Deposits within fluctuating profundal Zones B and D have higher δ¹³C (5.33‰ and 8.5‰, respectively) compared to fluvial-lacustrine Zones A and B (0.648 and 1.88‰) and are, therefore, interpreted to have supported higher rates of primary productivity.

Fischer Assay of balanced-filled and over-filled zones supports a primary productivity interpretation of δ¹³C results. In those zones having higher concentrations of δ¹³C, increased Fischer Assay values correspond. Zones B and D have an average Fischer Assay value of 17.6 Gal./ton and 19.2 Gal./ton, respectively. Comparatively, Zones A and C, which have lower δ¹³C values, exhibit significantly lower Fischer Assay measures of organic content at 7.6 Gal./ton and 9.7 Gal./ton, respectively.

Another possible influence on the δ¹³C profile across the Tipton Member is the rate at which particulate organics decompose within the lake system. As organic material decomposes, δ¹²C is released into the lake waters by methanogenesis, thereby decreasing the δ¹³C values recorded by sediments (Pitman 1996). Consequently, low δ¹³C may be used as a proxy to evaluate the extent to which lake waters were chemically and thermally stratified. Decomposition of organic matter is reduced if not stopped entirely within a stratified lake system. Within anoxic conditions of a stratified lake system, less δ¹²C would be available to dilute δ¹³C concentrations within the stratified lake. In this way, fluctuating profundal Zones B and D reflect periods of well-stratified lake waters. Low δ¹³C values within fluvial-lacustrine Zones A and C contrarily reflect less-stratified lake waters.

Carbon inputs from limestone bedrock (δ¹³C ≈ 0) can also lower the δ¹³C value of lake sediments. Lower Phanerozoic sections surrounding the GGRB are carbonate, and marginal conglomerates include limestone and dolostone clasts. Though these may have influenced the δ¹³C signature of the Tipton Member, the modeling of profundal deposits of the Laney Member (Doebbert 2006; Doebbert et al. 2010) indicates such an effect could not produce δ¹³C shifts observed in the GGRB.

Correlation between δ¹⁸O and δ¹³C values is often interpreted as a reflection of hydrologically closed lakes because of increased residence time of lake waters (Talbot 1990; Pitman et al. 1996) and successive, rapid lake volume oscillations (Li and Ku 1997). These observations coincide with correlation trends observed among intervals of the Tipton Member, where balanced-filled (i.e. closed lake-basin) deposits of fluctuating profundal intervals (Zones B and D) have a stronger correlation (R² = 0.2392) than overfilled (i.e. open lake-basin) deposits of fluvial-lacustrine Zones A and C (R² = 0.0835) (Fig. 3.7c).

3.7.3 Possible Origins of Isotopic and Lake Type Variation

Long-term variation in the rates of precipitation and evaporation are one possible explanation for variation in δ¹⁸O values of lake sediments, and would likely have varied following orbital changes to summer insolation (Morrill et al. 2001). The pace and magnitude of isotopic variation within the Tipton Member is difficult to relate to climate oscillations. In order to produce the dramatic δ¹⁸O shifts observed within the Tipton Member, temperatures would have had to have oscillated between extreme values during the Eocene. The Paleocene-Eocene Thermal Maximum (PETM) in the Bighorn Basin for example is marked by a 2‰ shift in δ¹⁸O (Koch et al. 2003), and has been associated with a 4–6 °C temperature increase within the western United States (Fricke et al. 1999; Fricke and Wing 2004), as well as a global sea surface warming of 8 °C (Zachos et al. 2001, 2008). To produce the 4.46‰, 6.68‰, and 7.64‰ δ¹⁸O shifts recorded within the Tipton Member (Shifts 1, 2, and 3, respectively), temperature change during Tipton deposition would have been pronounced, rapid in onset, and persistent, and have no clear corollary, in the marine record Westerhold and Röhl (2009). Existing geochronology (Smith et al. 2008, 2010) constrains Rife bed deposition to 0.60 ± 0.31 m.y. (2σ). Assuming all three litho-stratigraphic zones within the Rife bed (zones B, C, and D) were deposited in approximately the same amount of time, each would represent ca. 200 k.y., and may potentially coincide with orbital eccentricity variations.

The lithofacies, aerial extent, and δ¹⁸O characteristics of the two different lake types make is difficult, however to relate isotopic changes to climate parameters. Following the evaporation-driven hypothesis, during increased periods of evaporation, lake level should have diminished, subaerially exposing broad areas of the lake margin, which would have in turn led to evaporative concentration of ¹⁸O in lake waters, which is inconsistent with lithofacies-based observations of high lake level during deposition of δ18O-heavy zones B and D (Figs. 3.6 and 3.9). These fluctuating profundal zones consist largely of deep lake deposits that lack evidence for systematic desiccation and consistently overlie shallower-water lithofacies of underlying fluvial-lacustrine zones A and C.

Fig. 3.9

Interpreted four part synoptic lake type evolution of the Tipton Member. A major water and sediment source entered the GGRB during fluvial lacustrine deposition (zones A and C), and was diverted away from the basin during fluctuating profundal deposition (zones B and D)

Diversions of upstream drainage provides an appealing solution to the apparent paradox of high lake level and high δ¹⁸O composition observed in the Tipton Member. In such a scenario, abrupt changes to regional hydrology and isotopic composition would have been triggered by the inclusion or exclusion of a source of light δ¹⁸O-depleted waters and sediment delivered from an upland source (e.g., Carroll et al 2008). In our preferred model for the lake-type shifts observed within the Tipton Member (Fig. 3.9), Shift 1 is attributed to diversion of such a stream, followed by its recapture (Shift 2) and ultimate diversion (Shift 3). With each drainage diversion event, deposition of the Farson Sandstone ceased, accommodation increased in the absence of sediment input within the basin center, and lake margins correspondingly transgressed. Removal of a low δ¹⁸O hydrologic source and the associated increase in lake water residence time due to the lowered hydrologic throughput would both have acted to elevate the δ¹⁸O composition of Lake Gosiute. Doebbert et al. (2010) interpreted a similar magnitude δ¹⁸O and provenance shift in the Laney Member of the Green River Formation to have been triggered by the capture of a hinterland sourced stream. They interpreted it to be the terminus of a series of paleovalleys mapped by Janecke et al. (2000), the “Idaho River”. Though precise paleotopography of the upstream fluvial network(s) that triggered Tipton Member lake-type reorganizations is not certain, the location of its entry point in the northwest corner of the Greater Green River Basin, the areal distribution and progradation direction of the Farson Sandstone, and the presence of hinterland-sourced quartzite cobbles to the alluvial Pass Peak Formation (Smith et al. 2008) suggest strong similarities between the stream the fed Gosiute during zones A and C of the Tipton and the aforementioned Idaho River. The river in question would have likely have utilized a narrow pathway between the Wind River and Teton-Gros Ventre uplifts in the northern part of the GGRB, but would have delivered far less volcaniclastic detritus to Lake Gosiute because it predated the main phase of Challis volcanism (Smith et al. 2008). The ultimate cause for particular avulsions and captures remains enigmatic. Some could have been triggered by episodic faulting at upstream pathways between growing geologic structures within the drainage network or alternatively some could have occurred in an entirely autogenic fashion due to stream processes. In either case, any successful paleogeomorphic model for Tipton Member must account for the avulsion, return, and subsequent avulsion of this (or another) fluvial source to the basin during its deposition.

3.8 Conclusions

• 1. Thirteen distinct lithofacies occur within the Tipton Member and comprise two lithofacies assemblages: fluvial-lacustrine and fluctuating profundal, both of which have distinct basin-ward and shoreward expressions.

• 2. The Scheggs-Rife contact coincides with the first transition from fluvial-lacustrine (overfilled) to fluctuating profundal (balanced-filled) deposits.

• 3. Though lithostratigraphic transitions between profundal assemblages are subtle, distinct mineralogical, stable isotopic, and Fischer Assay oil yield values delineate abrupt lake system transitions.

• 4. Fluvial-lacustrine (overfilled) intervals exhibit prograding stratal geometries, dominantly calcitic carbonate mineralogy, low Fischer Assay oil yield, and light δ¹⁸O and δ¹³C values. Deposition is interpreted to have occurred within an open lake basin.

• 5. Fluctuating profundal (balanced-filled) intervals are defined by vertically aggrading stratigraphic geometries, dominantly dolomitic carbonate mineralogy, high Fischer Assay oil yield, and heavy δ¹⁸O and δ¹³C values. Deposition is interpreted to have occurred within an intermittently-closed lake basin.

• 6. Contrary to previous interpretations, the Rife bed contains both fluvial-lacustrine and fluctuating profundal intervals. The lower Rife bed is characterized by fluctuating profundal deposits, the middle Rife bed by fluvial-lacustrine deposits, while the upper Rife bed exhibits fluctuating profundal deposits.

• 7. In the northern GGRB, the Farson Sandstone is a lateral equivalent of both the Scheggs bed and zone C of the overlying Rife bed, and is thus constrained to fluvial-lacustrine intervals.

• 8. Oscillating lithologic and stable isotopic signatures within the Tipton Member are thought to reflect paleohydrologic reorganizations of the Farson Sandstone-sourcing fluvial system. Specifically, a diversion, recapture and ultimate diversion of this source are thought to have resulted in Shifts 1, 2 and 3.

• 9. Two possible mechanisms for paleohydrologic reorganization are proposed: episodic faulting or uplift upstream of the basin; and dynamic geomorphology of the Farson Sandstone-sourcing river itself. In both instances, a major fluvial source is diverted outside of the GGRB.