high-resolution stratigraphy of an underfilled lake … · a result, sedimentary facies...

Journal of Sedimentary Research, 2006, v. 76, 1197–1214

Research Article

DOI: 10.2110/jsr.2006.096

HIGH-RESOLUTION STRATIGRAPHY OF AN UNDERFILLED LAKE BASIN: WILKINS PEAK MEMBER,EOCENE GREEN RIVER FORMATION, WYOMING, U.S.A.

JEFFREY T. PIETRAS* AND ALAN R. CARROLL1University of Wisconsin-Madison, Department of Geology and Geophysics, 1215 West Dayton Street, Madison, Wisconsin 53706, U.S.A.

e-mail: [email protected]

ABSTRACT: Lakes tend to respond noticeably to minor changes in sediment and water balance driven by climatic, tectonic, orgeomorphic processes. This unique behavior of lacustrine basins can provide a high-resolution record of geologic processeswithin the continental setting, far from the globally averaged record of marine strata. The Wilkins Peak Member of the EoceneGreen River Formation, in Wyoming, USA, is dominated by aggradation of repetitive sedimentary facies successions recordingdistinct lacustrine expansions and contractions. These lacustrine ‘‘cycles’’ consist of up to four successive facies associations:littoral, profundal–sublittoral, palustrine, and salt pan. Because they comprise disparate facies that may never have beensimultaneously deposited in the basin, Wilkins Peak Member cycles are non-Waltherian successions that do not readily equateto any established sequence stratigraphic unit.

The completeness of the stratigraphic record in the Wilkins Peak Member varies continuously across the basin. At least 126cycles are present in the ERDA White Mountain #1 core near the basin depocenter, whereas only about one third as many arerecognizable 53 km north, nearer to the basin margin. Cycle boundaries terminate northward by gradual amalgamation intopalustrine facies, reflecting the interplay between varying lake levels and a south-dipping deposition gradient. Evidence forcomplete desiccation and hiatuses is commonplace even near the basin center; a truly complete record may not be presentanywhere.

Based on recent 40Ar/39Ar geochronology of tuffs, the average cycle duration in the White Mountain #1 core isapproximately 10,000 years. However, the true average duration is shorter due to the presence of lacunae, and the timerequired for deposition of the thinnest cycles may have been less than 1000 years. No external driving mechanism is presentlyknown for Eocene cycles of such short duration, raising the possibility that they are instead autogenic cycles related togeomorphic instability of the surrounding landscape.

Evaporite deposition corresponded closely in time and space to maximum differential subsidence, suggesting that tectonicinfluences on drainage patterns and basin accommodation exerted a primary control. Differential accommodation was mostpronounced during deposition of the lower Wilkins Peak Member and became progressively less so after that time. Based onstructural and stratigraphic observations, a period of increased tectonic uplift commenced concurrent with deposition of thelower Wilkins Peak Member.

INTRODUCTION

Much of our knowledge of continental climate and of terrestrial faunaland floral evolution comes from records preserved in lacustrine andrelated alluvial strata (e.g., Wolf 1978; McCune 1987; Stiassny and Meyer1999). Furthermore, because most long-lived lacustrine basins arefundamentally tectonic in origin (Kelts 1988), they potentially providehighly resolved archives of orogenic processes and related landscapemodification (e.g., Yuretich 1989; Lambiase 1990). However, thesensitivity of lakes to paleoenvironmental change can also complicatetheir stratigraphy, because of the potential rapidity of these changesrelative to the time required for sediment deposition (e.g., Owen et al.1990; Oviatt et al. 1992; Scholz and Finney 1994; Benson et al. 1995). Asa result, sedimentary facies representing vastly different depositional

environments may occur repetitively within relatively short verticalsuccessions (e.g., Gilbert 1882; Glenn and Kelts 1991; Talbot and Allen1996). This is especially true in underfilled basins, where stratal stackingpatterns record repeated extreme fluctuations in base level (Carroll andBohacs 1999).

A second problem common to many lacustrine deposits is the absenceof cosmopolitan marine fauna, which hinders biostratigraphic correlationof these strata to the geologic timescale. Milankovitch-period deposi-tional cyclicity has therefore often been proposed as an alternative meansof establishing high-resolution chronostratigraphy, on the basis thatorbitally controlled changes in solar insolation directly govern sedimen-tation rates and depositional environments in lakes (e.g., Bradley 1929;Van Houten 1962; Olsen 1986; Fischer and Roberts 1991; Roehler 1993;Fischer et al. 2004). However, the premise of externally regulated cyclicityhas seldom been tested against other independent measures of elapsed orabsolute geologic time (Steenbrink et al. 1999; Pietras et al. 2003a).Extreme Holocene lake-level changes are known to have occurred over

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sub-precessional time scales for reasons unrelated to orbital forcing (e.g.,Malde 1960; Johnson et al. 1996; Bouchard et al. 1998; Colman 1998;Reheis et al. 2002). Additionally, many previous proposals of Milanko-vitch cyclicity in lake deposits have been based either on single measuredsections or wells (e.g., Olsen 1986; Fischer and Roberts 1991; Roehler1993) or else on one-dimensional composites of overlapping butgeographically separated intervals (e.g., Steenbrink et al. 1999; Olsenand Kent 1996). It is therefore largely unknown to what extentgeographic variation in sedimentary facies or in cycle preservation mightaffect the outcome of such studies.

The Wilkins Peak Member of the Green River Formation ofsouthwestern Wyoming, U.S.A., is ideally suited for addressing theseproblems, due to its excellent exposures and to recent advances in its40Ar/39Ar geochronology (Smith et al. 2003). The general tectonic andclimatic setting of the Green River Formation has been well characterized(e.g., Love 1970; Wolf 1978; Gries 1983; Steidtmann et al. 1983; Roehler1991; Roehler 1993; Markwick 1994; Wilf 2000). Furthermore, previousstudies have established a sound fundamental understanding of thedepositional environments recorded by these lacustrine strata (Culbertson1961; Eugster and Hardie 1975; Smoot 1983; Roehler 1992). In this studywe take advantage of the nearly continuous exposures along WhiteMountain, a cuesta on the western flank of the Rocks Springs Uplift, to

construct the first detailed (centimeter- to decimeter-scale) regional crosssection of the entire Wilkins Peak Member. This cross section extendsfrom near the basin center in the south to the basin margin in the north(Fig. 1). The unique basin-scale perspective it provides allows reinterpre-tation of the controls on lacustrine evaporite deposition, of the nature ofthe repetitive facies associations that dominate Wilkins Peak Memberstratigraphy, and of the temporal significance of these ‘‘cycles.’’

REGIONAL SETTING AND STRATIGRAPHY

The greater Green River Basin (gGRB) formed between the Sevier foldand thrust belt to the west, and several Laramide-style Precambrian-coreduplifts on the north, east, and south (Fig. 1; Dickinson 1979; Lawton1985; Dickinson et al. 1988; Yuretich 1989). Eocene deposition of thelacustrine Green River Formation in Lake Gosiute roughly coincidedwith late pulses of shortening in the both the Sevier belt and Laramideuplifts (Love 1970; Steidtmann et al. 1983; Dickinson et al. 1988; DeCelles1994). The occurrence of palms, ferns, gingers, and crocodilians indicatea subtropical, humid climate with muted, equable seasonality andpredominantly frost-free winters (MacGinitie 1969; Markwick 1994).Leaf-margin and leaf-area analyses yield interpreted mean annualtemperatures (MAT) of approximately 15u to 23uC and mean annual

FIG. 1.— Generalized geologic map of the greater green River Basin and the surrounding highlands showing the location of measured sections and selected Eocenefaults (Witkind and Grose 1972; Love 1970; Steidtmann et al. 1983; Bradley 1995). Maximum extent of the Tipton Shale, Wilkins Peak, and Laney members is fromRoehler (1993). Extent of evaporites is from Wiig et al. (1995).

1198 J.T. PIETRAS AND A.R. CARROLL J S R

precipitation (MAP) of 75 to 110 cm/year during deposition of the GreenRiver Formation (Wing and Greenwood 1993; Wilf et al. 1998; Wilf2000). Eocene climate models for southwestern Wyoming estimate similarMAT and MAP values (Matthews and Perlmutter 1994; Morrill et al.2001).

The Wilkins Peak Member consists of carbonate-rich lacustrine andassociated facies that were deposited during an underfilled basin phase(Carroll and Bohacs 1999). It is enclosed by balanced-fill lake deposits ofthe Tipton Shale and Laney members and grades laterally into variegatedred and green alluvial mudstone and sandstone of the Cathedral BluffsTongue of the Wasatch Formation (Fig. 2). On the basis of recent40Ar/39Ar geochronology, the Wilkins Peak Member spans approximately1 My (Smith et al. 2003). The Wilkins Peak Member is 275 m thick at thetype locality near Rock Springs, Wyoming (Fig. 1), and thickenssouthward towards the Uinta Mountains, where it reaches 411 m(Roehler 1993). Main lithologies include tan to olive calcareous anddolomitic mudstone, kerogen-rich micro-laminated micrite, calcareoussandstone, arkosic sandstone, and bedded evaporites, including trona andhalite (Bradley 1959; Culbertson 1969; Smoot 1983). Insect remains, plantdebris, and trace fossils are the only commonly observed fossils.

White Mountain exposes a complete section of the Wilkins PeakMember where beds can be traced for over 50 km. We described sixoutcrop sections from the north end of White Mountain, south to thetown of Rock Springs. From north to south they are Boar’s Tusk (BT),Breathing Gulch (BG), Apache Lane (AL), Microwave Reflector (MR),Stagecoach Boulevard (SB), and Kanda (KA) (Fig. 1). An additionalsection was described from the ERDA White Mountain #1 core (WM).The WM core was collected from the fringe of the evaporite depocenter(Fig. 1). Together these seven sections extend over 53 km, providing anoblique-dip view of the northern part of the southward-thickeningWilkins Peak Member.

LITHOFACIES ASSOCIATIONS

The ten lithofacies below are defined on the basis of lithology, texture,and sedimentary and biogenic structures, and are grouped into 5lithofacies associations. This classification is broadly consistent withprevious descriptions and depositional-environment interpretations (e.g.,Eugster and Hardie 1975; Smoot 1983). However, significant advanceshave been made in understanding depositional processes in modern salinelakes and related fluvial systems in the two decades that have passed sincethe previous generation of studies (e.g., Smoot and Lowenstein 1991;Platt and Wright 1992; Tooth 2000a). Likewise, models for lake basinevolution have been refined (e.g., Lowenstein and Hardie 1985; Kelts1988; Carroll and Bohacs 1999), justifying a reexamination of the WilkinsPeak Member.

Littoral Association

Intraclastic Conglomerate Lithofacies.—The intraclastic conglomeratelithofacies is composed of granule to pebble-size mudstone intraclasts ina matrix of calcareous very fine to fine-grained sandstone (Fig. 3).Bedding is wavy, and individual beds are commonly less than 10 cmthick. Conglomerate beds unconformably overlie beds of calcareous anddolomitic mudstone. Intraclasts are typically flat with rounded corners,although some are more spherical. The composition of intraclasts issimilar to that of the underlying calcitic and dolomitic mudstone beds.Plant debris is common.

Intraclasts presumably were derived from the underlying mudstonebeds, which were at least partially lithified prior to being eroded (cf.Freytet and Plaziat 1982). Previous authors suggested that many of the‘‘poker chip’’-like intraclasts originally formed as the tops of desiccationpolygons on an exposed lake plain that were subsequently eroded during

FIG. 2.— North–south stratigraphic crosssection of the Green River Formation. Note theunique southward thickening of the Wilkins PeakMember compared to the Tipton Shale and Laneymembers, and the crosscutting relations betweenbasin bounding faults and the lower Wilkins PeakMember. Modified from Roehler (1989).

WILKINS PEAK MEMBER CYCLICITY AND BASIN EVOLUTION 1199J S R

flooding (Eugster and Hardie 1975; Smoot 1983). We adhere to theseprevious interpretations that the intraclastic conglomerate lithofaciesrepresents wave reworking of exposed lake-plain material into relativelyshallow water, through a combination of sheet floods and wave erosion.

Calcareous Sandstone Lithofacies.—The calcareous sandstone lithofa-cies is composed of interbedded very fine- to fine-grained calcareoussandstone and thin mudstone beds (Fig. 4A). Beds are typically 10 to50 cm thick, have sharp bases, and crop out as distinct white beds thatcan be traced for tens of kilometers. Grain types include carbonate,quartz, feldspar, mica, plant debris, and ooids (Fig. 4B). Mudcrack casts(Fig. 4C) and granule to pebble-size mudstone intraclasts are commonalong the base of beds. Internally, sandstone beds contain wavy cross-lamination and mudcracks, and preserve symmetrical ripple marks andrare bedding-plane trace fossils along bed tops. Interbedded mudstonelayers are typically less than 1 cm thick and tend to drape the tops ofsandstone beds.

Wave action is interpreted to be the dominant mode of sedimenttransport for the calcareous sandstone lithofacies, and is recorded by theubiquitous occurrence of nonparallel wavy lamination and symmetricalripple marks (cf. Allen 1981). The common occurrence of mudcrackssuggests frequent exposure and desiccation. Mudstone intraclasts werederived from underlying beds, similar to those found in the intraclasticconglomerate lithofacies. Interbedded mudstone drapes suggests periodsof slack water allowing finer-grained material to settle from suspension(cf. Reineck and Wunderlich 1968).

Summary of Littoral Association.—This association records shallow-water deposition during shoreline transgression over a previously exposedand partially cemented lake plain. These deposits are common to thelittoral zone, where wave action is the primary transport process. In thisdepositional environment relatively small falls in lake level can exposebroad areas, particularly when the depositional gradient is very low, aswas the case for the gGRB. These high-frequency fluctuations result inthe repetition of wave-transported facies associated with numerousdesiccation horizons.

Profundal–Sublittoral Association

Laminated Kerogen-Rich Calcareous Mudstone Lithofacies.—Sub-mil-limeter-scale laminated calcareous and dolomitic mudstone appears asblack or dark brown beds along freshly exposed surfaces and weather toa white or slightly blue-white color. Beds range in thickness from 0.1 to2 m and can be traced laterally for tens of kilometers. The preservation oforganic debris suitable for petroleum generation gives rise to the term ‘‘oilshale’’ that has been applied to this lithofacies. Oil yields from kerogen-rich strata in the Wilkins Peak Member reach up to 40 gallons per ton

(GPT) determined by Fisher Assay analysis, with TOC (total organiccarbon) ranging up to 19% (Bohacs 1998; Carroll and Bohacs 2001).

The occurrence of kerogen-rich strata reflects the combined affects ofprimary productivity, preservation, and lack of significant dilution byinorganic sediment (Bohacs 1998). Much of the organic materialpreserved in the Wilkins Peak Member was derived from algae andbacteria (Bradley 1962; Tissot and Vandenbroucke 1983; Bohacs 1998).The preservation of organic material and sub-millimeter-scale laminationsuggests inhibition of burrowing organisms. This commonly results fromlow oxygen levels in lake bottom waters (e.g., Demaison and Moore 1980;Talbot and Allen 1996). High salinity levels, inferred for the Wilkins PeakMember from the presence of evaporites, also tend to limit benthicorganisms. Biomarker analyses from organic-rich beds in the WilkinsPeak Member yields results consistent with deposition in low-oxygen,hypersaline waters (Carroll and Bohacs 2001).

Massive Calcareous Mudstone Lithofacies.—This lithofacies is similarto the laminated kerogen-rich calcareous mudstone lithofacies. However,millimeter-scale laminae are not as well preserved, beds are light brown toolive in color, and oil yields based on Fischer Assay analyses are lower(typically less than 10 GPT). The reduction of organic-richness and lossof well-developed lamination in this lithofacies presumably reflects theactivity of benthic organisms. This lithofacies suggests that bottom waterswere more oxygenated than during deposition of the kerogen-richlithofacies, implying shallower water depths.

Summary of Profundal–Sublittoral Association.—These strata representperiods of relatively high lake level and lateral lake expansion. The warm-equable climate inferred for the gGRB (see above) may have contributedto the persistence of density stratification in a meromictic lake withoxygen-depleted bottom waters (cf. Tucker and Wright 1990; Talbot andAllen 1996). In addition, a more laterally expansive lake would havetended to restrict the delivery of inorganic sediment to lake-marginalareas, further concentrating organic matter in areas nearer the lakecenter. Fischer Assay oil yields and percentage of total organic carbongenerally correlate with evidence from sedimentary facies for relativelydeep-water conditions (Carroll and Bohacs 2001).

Palustrine Association

Wavy-bedded Calcareous Siltstone–Sandstone Lithofacies.—The wavy-bedded calcareous siltstone and sandstone lithofacies is composed of lightgray calcareous siltstone and sandstone that occur as centimeter-thicklaminae interbedded with calcareous and dolomitic mudstone. Ripplecross-lamination and ripple marks are common. Many laminae are onlyone ripple height in thickness, but composite siltstone–sandstone–mudstone beds can be up to 2 m thick.

FIG. 3.— Intraclastic conglomerate lithofacies(polished slab). Note the abundance of mudstoneintraclasts at the base of this bed.


Silt and sand grains of this lithofacies were transported by wave actionas ripple trains, and mudstone interbeds were deposited by settling duringslack-water periods (cf. Reineck and Wunderlich 1968). Isolated rippletrains suggest a sand-starved setting where coarse-grained material isavailable only during floods (Allen 1981). This lithofacies thus representsdeposition in relatively shallow water, where silt- to sand-size sedimentswere transported by wave action during intermittent floods.

Brecciated Carbonate Mudstone–Siltstone Lithofacies.—This lithofaciescommonly occurs as recessive interbeds, 0.1 to 3 m thick, of olive andlight brown calcareous and dolomitic and olive to gray calcareoussiltstone. Beds are internally massive, but wavy laminae are preservedlocally. Mudcracks (Fig. 5A), angular mudstone intraclasts, and plant

debris are common. Calcareous nodules, trace fossils (Fig. 5A), and insectremains occur locally. Many beds are completely brecciated, and in somecases evaporite minerals filled the voids between clasts (Fig. 5B).

The local occurrence of wavy bedding suggests movement of grains bywave action. However, primary laminae are obscured by the formation ofdesiccation cracks, burrowing, and the growth of plant roots in mostbeds. These processes lead to brecciation and development of massivebedding. This lithofacies thus signifies an alternation between subaqueousdeposition of mudstone and siltstone with subaerial modification.

Previous studies argue for a primary or very early diageneticdolomitization (Eugster and Surdam 1973; Eugster and Hardie 1975;Surdam and Wolfbauer 1975; Smoot 1983; Mason and Surdam 1992;Norris et al. 1996; Pitman 1996). Deep Springs Lake in California (Jones1965) and several saline lakes of western Victoria, Australia (De Deckkerand Last 1988), offer potential modern analogues of primary dolomiteprecipitation. In these playa-lake settings calcite is precipitated on theouter fringe of a playa surrounding a central lake, increasing the Mg/Caratio of residual waters. Dolomite precipitation occurs from these Mg-rich and sulfate-depleted waters along the lake margin when salinity andalkalinity are elevated by evaporation (Tucker and Wright 1990). Calciteand dolomite clasts can then be transported into the lake by wind andwater.

Summary of Palustrine Association.—The intimate association oflacustrine carbonate accumulation and subaerial pedogenic modificationcommon to these units suggests deposition in the littoral to eulittoralzones (Tucker and Wright 1990). Together these zones form thepalustrine environment (Freytet and Plaziat 1982; Freytet and Verrecchia2002). In this environment pedogenic modification occurs on local highsor away from the shoreline, while carbonate precipitation occurs in lowerareas that alternate between wet and dry conditions (e.g., Platt andWright 1992). Associated lacustrine environments derive sediment fromthese areas during floods. Palustrine environments are most expansive inlow-gradient basins, like the gGRB, because minor changes in waterinflow can result in frequent wetting and drying of very broad areas.

Salt-Pan Association

Displacive Evaporite Lithofacies.—The displacive evaporite lithofaciesis composed of zones of evaporite crystals, commonly shortiteNa2Ca2(CO3)3 and trona Na3(HCO3)(CO3)?2(H2O), found within bedsof the previously described lithofacies. Shortite occurs as randomlyoriented euhedral grains ranging from microcrystalline to 1 mm to 3 cmacross, or as discrete layers (Fig. 6A). Trona typically occurs ascentimeter-scale radial blades (Fig. 6B) and microcrystalline aggregates.These evaporite crystals crosscut and distort primary lamination in thesurrounding host lithology, and lithofacies boundaries.

Because these evaporite minerals both truncate and distort primarylamination they are interpreted to have formed in the unlithified lake-bottom sediment from supersaturated brines (cf. Lowenstein and Hardie1985). Brines that form displacive evaporites are derived either directlyfrom lake water or from interstitial waters (Smoot and Lowenstein 1991).

Bedded Evaporite Lithofacies.—Evaporite minerals also form discretebeds throughout the Wilkins Peak Member. Culbertson (1971) reported42 evaporite beds, 25 of which are thicker than 1 m. These thicker bedsrange in area from 427 to 1,870 km2 (Bradley and Eugster 1969). Theirmaximum extent is shown in Figure 1. The most common mineralsinclude trona and halite, but numerous other saline minerals aredescribed in varied quantities (Bradley and Eugster 1969). Mudstoneinterbeds containing wavy laminae are common within evaporite beds(Fig. 6C). Evaporites are present as pseudomorphs in outcroppings(Fig. 6D).

FIG. 4.—Calcareous sandstone lithofacies. A) Resistant layer of the calcareoussandstone interbedded with deeply eroded mudstone lamina (outcrop). B)Photomicrograph of a calcareous sandstone bed; q, quartz; f, feldspar; o, ooid;c, calcareous matrix. C) Mudcrack casts on the base of a calcareous sandstone bed(hand sample).


Bradley and Eugster (1969) described both bladed and granulartextures of trona with inclusions of organic material, dolomicrite, silt, andpyrite that give beds a brownish color. They interpreted these beds, whichare intercalated with profundal mudstone, to represent primary tronadeposited subaqueously on the floor of Lake Gosiute. Similar tronadeposits have been described from modern lakes Magadi and Bogoria(Eugster 1970; Renaut and Tiercelin 1994) and from the NeogeneBaypazari Basin in Turkey (Helvaci 1998).

Summary of Salt-Pan Association.—This association represents thedevelopment of a salt pan (Lowenstein and Hardie 1985) during extendedperiods when evaporation exceeded influx of water to Lake Gosiute. Thecommon interbedding of evaporite and mudstone layers suggestsintermittent periods of flooding, freshening, and sediment advectionfrom the surrounding lake plain, followed by evaporative concentration.Similar mudstone interbeds have been described associated with tronadeposits of Searles Lake, California (Fairchild et al. 1998).

Fluvial Association

Siliciclastic Sandstone Lithofacies.—The siliciclastic sandstone lithofa-cies is distinct from other sandstone facies in the Wilkins Peak Memberbecause of its brown color and almost ubiquitous occurrence of climbingripples (Fig. 7A). Typical grain types include quartz, feldspar, and micathat range from silt to fine-grained sand. Individual beds range up to 2 min thickness; however, composite sandstone–siltstone–mudstone bedsetsare up to 25 m thick and can be traced across much of the basin(Culbertson 1961). Many beds consist of an amalgamation of severallenticular sandstone bodies stacked vertically and shingled laterally. Inaddition to climbing ripples, common sedimentary structures includedecimeter to meter-scale scour and fill (Fig. 7B), trough cross-bedding,and parting lineation. Both calcareous and noncalcareous mudstoneintraclasts are common in bed bases. Vertical burrows are commonwithin beds, and more exotic trace fossils such as bird trackways(Fig. 7C), insect nests (Bohacs et al. 2002), and mammal footprints occurlocally.

Lenticular sandstone bodies in this lithofacies are interpreted to besand-filled fluvial channels. These channels scoured into previouslydeposited mudstone and sandstone beds, creating laterally extensiveamalgamated sandstone units. Climbing ripples signify rapid depositionfrom sediment-laden unidirectional flows (Allen 1970, 1971). Climbingripples have been described in fluvial systems associated the waning stagesof floods (Sneh 1983) and from the deposits of density flows in a standingbody of water (Jopling and Walker 1968; Stanley 1974). Insect nestsassociated with some sandstone beds would require that that theseintervals were above the water table for significant periods of time(Bohacs et al. 2002). For this reason we favor a fluvial origin for climbingripples in this lithofacies.

Green to Brown Mudstone and Siltstone Lithofacies.—The brownmudstone and siltstone lithofacies is distinct from other mudstone faciesbecause of its color and its typically noncalcareous nature. Sheet-likearkosic sandstone interbeds 1 to 20 cm thick, plant debris, carbonaceousmaterial, and calcareous concretions are common. Beds are locallymottled red and green and weather to a blocky texture. Horizontal and

R

Fig. 5.—Brecciated calcareous mudstone–siltstone lithofacies (WM core). A)Burrows and mudcracks. B) Brecciated mudstone with shortite crystals betweensome clasts.


vertical surfaces marked by color change, differential cementation, and/ormodern fractures occur locally (Fig. 7D).

This lithofacies is interpreted to represent floodplain deposits. Fine-grained material was deposited from suspension in abandoned channels,in an overbank setting, or in small lakes formed during flood intervals(Miall 1996). Thin sandstone interbeds presumably represent sedimentthat was transported as bedforms during occasional floods such ascrevasse splays (Miall 1992). Calcareous concretions and associated redand green mottling suggest incipient pedogenesis, common to floodplainenvironments (Fenwick 1985; Retallack 1988). The vertical featuresobserved locally (Fig. 7D) also signify pedogenic alteration. They areinterpreted as pseudoanticlines that form by repeated swelling andshrinking of clays caused during seasonal wetting and drying in a soilprofile (Yaalon and Kalmar 1978; Paik and Lee 1998).

Summary of Fluvial Association.—This association corresponds to thenine sandstone–siltstone–mudstone marker beds (A through I) describedby Culbertson (1961). These marker beds are up to 25 m thick in theWilkins Peak Member and can be correlated across the basin (Fig. 2).These beds are thought to represent the deposits of braided streams andsheetfloods, and fine-grained overbank deposits (Smoot 1983). Duringflood intervals shallow lakes may have formed, depositing thin units ofsiltstone and mudstone strata towards the basin center and in local lows,signifying temporary reestablishment of lacustrine environments. How-ever, long-term discharge remained low, resulting in desiccation, return ofa floodplain environment, and subsequent pedogenic alteration. Theseinterpretations may point toward a ‘‘dryland river’’ model for the fluvialassociation (cf. Picard and High 1973; Tooth 2000a, 2000b). However,a full evaluation of this hypothesis is beyond the scope of the present

FIG. 6.— Displacive evaporite and bedded evaporite lithofacies. A) Photomicrograph of evaporite crystals (shortite) in a matrix of calcareous mudstone. Note howsome laminae are both deflected and truncated by the growth of evaporite crystals. B) Bladed trona (WM core). C) Bedded trona overlain by calcareous mudstone(sample collected from the FMC trona mine). Some displacive evaporite (DE) crystals can be seen in the mudstone layer. D) Outcrop photo of bladedtrona pseudomorphs.


study. The fluvial association is much better exposed to the south of thestudy area, offering an excellent opportunity for future work.

WILKINS PEAK CYCLES

The littoral, profundal–sublittoral, palustrine, and salt-pan associa-tions occur in repetitive facies successions that represent periods of lakeexpansion and contraction (Fig. 8; Eugster and Hardie 1975; Smoot1983). These repetitive successions are referred to as ‘‘cycles.’’ We do thisin order to be consistent with many past studies of carbonate-rich marinestrata and both carbonate and clastic lacustrine deposits (e.g., VanHouten 1962), but we do not intend our use of this term to imply anyspecific periodicity or driving mechanism.

Cycles in the Wilkins Peak Member are generally asymmetric in termsof facies thickness, grading upward from relatively thin intervals of

littoral and profundal–sublittoral facies associations, into thicker inter-vals of palustrine facies (Fig. 8). Overall thickness is highly variable,ranging between 0.14 and 5.70 m in this study. Cycles are bounded bywidely traceable surfaces that commonly are marked by desiccationcracks or centimeter-scale scour. These surfaces are usually overlain byintraclastic conglomerate and calcareous sandstone of the littoralassociation. However, in some cycles profundal lithofacies directly overliethe basal surface. In all cases the cycle boundary is abrupt and reflects anadvancement of the shoreline and an increase in water depth. Theprofundal–sublittoral association represents the maximum lake expansionduring the deposition of an individual cycle. The upper contact of theprofundal–sublittoral association is gradational with overlying mudstoneand siltstone beds of the palustrine association, suggesting gradualshoreline retreat and decreasing water depths. Most of the cycle thicknessrecords shoreline regression, in as much as palustrine facies typically

FIG. 7.—Siliciclastic sandstone and brown mudstone–siltstone lithofacies. A) Outcrop example of climbing ripples. B) Meter-scale scour and fill feature in a sandstoneoutcrop approximately 25 m in height. C) Casts of a bird trackway on a bedding plane of sandstone. D) Outcrop showing both vertical features (v) and horizontalfeatures (h) in brown mudstone–siltstone interpreted as soil features. The ledge at the top of the photo (ss) is a sandstone bed.


comprise two-thirds or more of the facies present. The profundal–sublittoral association is absent in some cycles, in which case thepalustrine association directly overlies the littoral association. Thisamalgamation most commonly occurs towards the basin margin(Fig. 9). Conversely, several cycles in our most basinal section (WM)are capped by bedded evaporites and/or zones of displacive evaporitemineralization, recording formation of a salt pan (Fig. 9).

Lateral Correlation of Cycles

Roehler (1993) documented a maximum of 77 cycles from the ERDAWhite Mountain #1 core (WM; Fig. 1) based on Fischer Assay analysisof retort oil yields. Peaks in oil yield were used as a proxy for lakehighstands. This approach, however, did not account for all episodes oflake expansion and contraction. Fischer Assay oil yields of the WM corewere measured from homogenized samples that averaged 30 cm to over1 m of core, whereas lacustrine cycles as thin as 14 cm can be recognizedvisually. Furthermore, some of the cycles are not sufficiently kerogen-richto be identified using Fischer Assay alone (Fig. 10). On the basis ofdetailed lithofacies descriptions, 126 cycles were identifiable in the WMcore.

Wilkins Peak Member cycles can be correlated from near the northernbasin margin towards the basin center, primarily using exposures on thewestern flank of the Rock Springs Uplift (Fig. 11). Correlations are aidedby the presence of several distinctive volcanic air-fall tuffs and by the ninecomposite bedsets of the fluvial association. Overall thickness patterns

and northward-terminating cycles indicate that accommodation wasgreater towards the south (Fig. 11). The number of cycles decreasessystematically northward, from 126 in the WM core to 46 at BG and 34 atBT. In contrast to the progressive basal onlap depicted by Roehler (1993),detailed correlations indicate that cycle boundaries gradually becomeunrecognizable northward, due to amalgamation within homogeneouspalustrine facies (Fig. 10; Pietras et al. 2003a). The maximum limit of anyparticular lake expansion can be roughly constrained between twospecified measured sections, but its precise position is difficult orimpossible to identify.

Figure 12 displays changes through time in differential basin sub-sidence, based on a comparison of cycle thickness near the basin marginto the thickness of the same interval near the basin center. Net sedimentaccumulation was essentially uniform across the study area duringdeposition of the Tipton Shale Member, indicating that the basin floorwas relatively flat. In the Wilkins Peak Member, three major episodes ofrapidly increased differential accommodation are evident, expressed byintensified southward thickening of cycles and increased number ofcycles. Following each episode, a gradual shift back to more uniformsediment accumulation is recorded in the overlying cycles. The lowestsuch interval begins at the base the Wilkins Peak Member and extendsroughly to the position of the Grey Tuff (Fig. 11). The next interval isroughly constrained between the Grey and Main tuffs, and the final unitextends from the Main Tuff to the contact with the Laney Member. Thesethree intervals are similar to, but not identical to, the lower, middle, andupper parts of the Wilkins Peak Member described by Roehler (1992).

FIG. 8.— Diagram of a typical lacustrine expansion–contraction cycle in the Wilkins Peak Member highlighting common sedimentary structures and the verticaldistribution of lithofacies associations. The photo to the right is one complete cycle in the ERDA White Mountain #1 core.


Differential sediment accumulation in these three intervals decreasedsystematically upward, returning to a relatively uniform configurationduring deposition of the Laney Member (Fig. 12).

The distribution of evaporite beds also reflects this pattern ofdifferential accommodation and indicates that the highest rates ofaccommodation occurred to the southwest of WM (Fig. 13). Threetrends appear from this data: (1) the number of evaporite beds and theirareal extent decreases upwards, (2) average evaporite bed thicknessdecreases upwards, and (3) the locus of deposition migrated northwardfollowing deposition of the lower Wilkins Peak Member. Collectively,these observations suggest that evaporite deposition generally coincidedwith periods of increased basin-floor slope and smaller lake surface areas,whereas more widespread and uniform lacustrine environments occurredwhen the basin floor was relatively flat.

DISCUSSION

Sequence Stratigraphic Interpretation of Lacustrine Cycles

The stratigraphy of the Wilkins Peak Member is dominated byrepetition of lake expansion–contraction cycles, revealing the extremesusceptibility of this system to high-frequency fluctuations in basinhydrology (Carroll and Bohacs 1999; Bohacs et al. 2000). A dilemmaexists regarding the proper sequence stratigraphic interpretation of thesecycles: they could potentially be classified as either parasequences or assequences, on the basis of the original definitions of those terms(Mitchum et al. 1977; Van Wagoner et al. 1988). In thickness, WilkinsPeak Member cycles are closer to typical shallow marine and fluvialparasequences than to sequences. The bulk of each cycle is represented byregressive palustrine facies, which is also consistent with a shallowing-

FIG. 9.—A) Model for the development of lacustrine expansion–contraction cycles in the Wilkins Peak Member. CB 5 cycle boundary, MES 5 maximum expansionsurface. B) Wheeler diagram showing how time is represented in Wilkins Peak Member expansion–contraction cycles.


upward parasequence. Bohacs (1998) noted that ‘‘lacustrine floodingsurfaces’’ might reasonably be substituted for the ‘‘marine floodingsurfaces’’ specified in the original parasequence definition (Van Wagoner1988). However, the relative magnitude of lake-level change associatedwith each cycle is extreme compared to the relative sea-level changesinferred from marine parasequences. Many Wilkins Peak Member cyclesbegin and end with complete or nearly complete desiccation, indicatingthat they record the entire base-level range of the lake. The absolute depthof Lake Gosiute during highstands is not known, but from comparisonwith Holocene lakes and the preservation of laminated mudstone itprobably reached tens or perhaps even hundreds of meters (e.g.,DeMaison and Moore 1980; Olsen 1990; Oviatt et al 1992; Scholz andFinney 1994).

Taking the alternative view, Wilkins Peak Member cycles appear tomeet the basic definition of a sequence as a ‘‘relatively conformablesuccession of genetically related strata bounded at its top and base byunconformities and their correlative conformities’’ (Mitchum et al. 1977).The pivotal question is whether the cycle boundaries are merely short-lived scour surfaces developed during transgression, or instead theyrepresent longer periods of basin-floor exposure and desiccation. Thisquestion is not easily answered, but because well-developed paleosols arerare it is possible that exposure times were relatively short. Furthermore,the cycles lack clearly defined lowstand facies. There are two possiblereasons for this. First, the flux of weathering products (both physical andchemical) into the basin was directly tied to the inflow of water, andtherefore would have decreased as lake level dropped. Much of the basinmay therefore have experienced greatly decreased depositional ratesduring dry periods. Second, the floor of Lake Gosiute had very low reliefand lacked any significant geomorphic shelf–slope break that could have

distinctly divided lowstand from transgressive deposits. Only thesurrounding bedrock surfaces provide any significant change in basin-floor slope, but presently there is no direct evidence for how highshorelines climbed on these basin walls.

In contrast to typical marine carbonate sequences, the transgressiveincrements of Wilkins Peak Member cycles are quite condensed. Most ofthe sediment deposited during transgression appears to have beensiliciclastic detritus and reworked carbonate lithologies rather than new,autochthonous production of carbonate. The enhanced biologic pro-ductivity that allows marine carbonate facies to ‘‘keep up’’ with rising sealevel (cf. Kerans and Tinker 1997) does not appear to have playeda significant role in Wilkins Peak Member deposition. Progradation ofcarbonate-rich palustrine deposits may have occurred during lakehighstands, but it is impossible to map such fine detail in thishomogeneous facies.

Wilkins Peak Member cycles may not correspond to any establishedsequence stratigraphic unit because they do not represent true ‘‘Walther-ian’’ facies successions. Walther’s law of facies succession states ‘‘… thatonly those facies and facies areas can be superimposed primarily whichcan be observed beside each other at the present time’’ (Middleton 1973).Sequences and parasequences are both defined as successions of relativelyconformable, genetically related strata, which are generally taken toimply that they adhere to Walther’s law. However, individual WilkinsPeak Member cycles contain facies that may never have laterallycoexisted within the gGRB. For example, evaporite beds lie directly overoil shale beds in some locations, but it is far from clear that both facieswere ever deposited simultaneously. Similar observations have beenreported from Holocene underfilled lacustrine basins such as the QaidamBasin in western China (Yang et al. 1995) and Death Valley, California

FIG. 10.—Detailed correlation of measured sections highlighting termination of cycle boundaries into lake-margin facies. Shale oil yield cycles from Roehler, 1993.GPT, gallon per ton.


FIG. 11.— High-resolution stratigraphic cross section of the Wilkins Peak Member at White Mountain. See Figure 1 for location of measured sections. Letters A–I correspond to the nine units of the fluvialassociation (Culbertson 1961). CBT, Cathedral Bluffs Tongue. Detailed versions of all measured sections appear in Pietras (2003).

1208J.T

.P

IET

RA

SA

ND

A.R

.C

AR

RO

LL

JS

R

(Lowenstein et al. 1998). In these basins perennial lake deposits liedirectly beneath modern salt pans. Wilkins Peak cycles probably are bestconsidered as event beds that reflect comprehensive paleoenvironmentalchange across the basin rather than representing gradual migration ofcoexisting environments. As such they are analogous to some marineevaporite deposits, for which a ‘‘non-Watherian’’ origin has also beenargued (e.g., Kendall 1988).

Cycle Duration vs. Orbital Periodicity

Previous authors have suggested that lacustrine expansion–contractioncycles in the Green River Formation were governed by changes inclimatic humidity, in turn driven by orbitally controlled changes in solar

insolation (e.g., Bradley 1929; Fischer and Roberts 1991; Roehler 1993).The cycles we describe in the Wilkins Peak Member have specifically beenattributed to 19–23 kyr precessional periodicities on the basis of FischerAssay analysis of organic matter content (Roehler 1993; Fischer et al.2004). Two direct observations challenge such interpretations. First,Fischer Assay data from the Wilkins Peak Member are not of highenough resolution to be used as a faithful proxy for lake level (see above).Secondly, the preservation of a variable number of cycles along a basintransect question the ability of any one section to represent a completestratigraphic record. In fact hiatuses appear to be plentiful even in themost basinal sections (Bohacs 1998), suggesting that a truly completestratigraphic record may not exist anywhere.

Recent high-resolution 40Ar/39Ar age determinations from four tuffs(Smith et al. 2003) in the Wilkins Peak Member provide an opportunityto directly constrain average apparent cycle duration at differentlocations by counting the number of cycles preserved between datedtuffs (Fig. 11). Based on the number of cycles present between each tuffpair, the average apparent cycle duration decreases by approximatelya factor of three from near the basin margin to near the center (Fig. 14;Table 1). All three intervals yield an apparent duration of approximately10,000 years for cycles in the WM core. It should be noted however thatthese durations are maximum values, because they do not take intoaccount hiatuses. These calculations also do not reflect the fact that cyclethicknesses in the subject interval vary by almost two orders ofmagnitude. For example, cycles in the middle Wilkins Peak Memberrange from 21 cm to 5.7 m (Fig. 15). Assuming that thinner cyclesrepresent shorter increments of time, the shortest true cycle durations aretherefore much less than 10,000 years, and in some cases may representlacustrine episodes that lasted less than 1000 years.

Durations of 10,000 years or less are too short to correspond to anyknown precessional period, even taking into account uncertainties inradioisotopic dating (Smith et al. 2003; Pietras et al. 2003a). Our studydoes not exclude the possibility that these short-period cycles combine toform longer cycles with precessional or other orbital periods. However, itis clear that alternative mechanisms are required for the genesis of sub-precessional cycles. Half-precessional frequencies observed in UpperCretaceous marine deposits from Brazil have been attributed to theexistence of an equatorial continent that was sensitive to both extremes ofthe precessional cycle (e.g., Short et al. 1991; Park et al. 1993). However,this mechanism is less plausible for an Eocene site in North America orfor cycles with durations nearer 1000 than 10,000 years. Ice-relatedmechanisms (such as Heinrich events) seem unlikely, given deposition ofthe Green River Formation during the warmest period of the Cenozoic(Smith et al. 2003) in a presumably ice-free world. However, Pekar et al.(2005) recently proposed the existence of early and middle Eoceneglacioeustatic sea-level changes, suggesting that this assumption mayrequire further examination. Other non-climatic mechanisms that havebeen implicated in causing expansion or contraction of lakes arenumerous, including tectonic modification of drainage divides (e.g., Saezet al. 1999; Pietras et al. 2003b), diversion of rivers by volcanic flows (e.g.,Bouchard et al. 1998), or vertical movements along basin outlets (e.g.,Malde 1960; Colman 1998; Reheis et al. 2002). These processes can occurrepetitively or as unique events, but they generally seem less than ideal forgenerating apparently ‘‘rhythmic’’ facies successions. Perhaps the mostpromising (and least explored) mechanism is geomorphic autogenesis,caused by migration of drainage divides or upstream propagation offluvial knickpoints (Hasbargen and Paola 2000). The temporal stability ofalluvial landscapes over periods of 103–104 years is largely unknownbut could potentially exert a first-order control on downstream basinsby modulating the flux of both water and sediment to downstreamlakes.

Interestingly, the variable updip-downdip preservation of cycles in theWilkins Peak Member acted as a natural frequency filter. Because fewer

FIG. 12.— Graphical correlation of timelines in the Wilkins Peak Memberbetween WM and BG. Inset figures show sediment accumulation across the basinfor each of the three portions of the Wilkins Peak Member. The nine sandstone–siltstone–mudstone beds of Culbertson (1961) are labeled A–I. All nine beds occurat WM, but only bed D is present at BG. WP, Wilkins Peak Member.


cycles are preserved in updip locations (nearer the basin margin), theirapparent frequency, based on cycle counts, is lower than cycles near thebasin center. Assuming that Wilkins Peak cycles actually record a rangeof different frequencies, it is likely that shorter ‘‘true’’ frequencies areprogressively excluded from the record in sections nearer to the basinmargins. At some geographic positions within the basin, the duration ofthe shortest preserved cycles may in fact correspond to the range ofknown precessional periods. However, the facies expression of WilkinsPeak cycles is similar throughout much of the study area, regardless oftheir thickness or true duration. It would therefore be impossible to

identify which cycles correspond to precession solely on the basis of theirexpression at one location.

Evaporite Deposition in Laramide Basins

Previous studies have implied that deposition of bedded evaporite inthe Wilkins Peak Member was caused by a period of markedly dryclimate (e.g., Bradley and Eugster 1969; Roehler 1993), but presentlythere are no independent data that directly support this conclusion.Paleofloral collections have been reported from the greater Green River

FIG. 13.—Vertical, lateral, and areal extent of25 potential economic trona beds (Culbertson1971) in the lower, middle, and upper Wilkins PeakMember. Darker colors on the map representstacked trona beds. Trona data from Wiig et al.(1995). Extent of Wilkins Peak Member is fromRoehler (1992). WR, Wind River Mountains; RSUplift, Rock Springs Uplift.


Basin (MacGinitie 1969; Wolfe 1978; Wilf 2000), but these liestratigraphically below and above the principal evaporite-bearingintervals. On the basis of leaf-margin and leaf-area analysis of leavescollected from the Niland Tongue (below the Wilkins Peak Member),Wilf (2000) interpreted a mean annual temperature of 23.0 6 3.7uC andmean annual precipitation of 100 cm/yr (standard error + 60.5,242.2 cm/y). For a second locality that spans the uppermost WilkinsPeak to lowermost Laney members (Little Mountain), he interpretedmean average values of 19.6 6 2.1uC and 76.9 cm/yr (standard error+ 33.2, 223.2 cm/yr). He further observed that changes in floralcomposition between these two sites were consistent with a climate thathad become cooler and drier, with more seasonal rainfall. However, Wilfalso pointed out that the Little Mountain flora was collected near thesouthern margin of the gGRB from an interval that records generallyrising average lake levels and that this collection therefore may not berepresentative of the conditions that prevailed during periods of evaporitedeposition. Fossil leaves have not been reported from the more evaporite-rich lower to middle Wilkins Peak Member, and therefore inferences ofa relatively dry paleoclimate are based on the occurrence of evaporitebeds themselves rather than on floral evidence.

In contrast to its somewhat ambiguous paleoclimatic setting, the spatialand temporal relationship of the Wilkins Peak Member to changing basintectonics is relatively clear. Thickness patterns (Fig. 12; Roehler 1992)

and evaporite distributions (Fig. 13) reveal a marked increase indifferential accommodation during deposition of the lowermost WilkinsPeak Member, consistent with tectonic loading of the southern basinmargin by renewed Uinta Mountain uplift (Hansen 1986; Roehler 1993;Bradley 1995). The detailed record presented in this study shows thatdifferential accommodation was greatest during deposition of the oldestWilkins Peak Member strata and gradually decreased through time(Fig. 12). Structural observations along the Uinta Mountain frontreinforce this interpretation. For example, Roehler (1993) estimatedroughly 150 m of offset of the Sparks Ranch thrust fault (Fig. 1) duringearly Wilkins Peak Member deposition, on the basis of cross-cuttingrelations with the laterally equivalent lower Cathedral Bluffs Tongue.Farther west, along the center of the Uinta Mountains, crosscuttingrelationships demonstrate that the Henry’s Fork thrust fault (Fig. 1) wasalso active during the late early to early middle Eocene (Bradley 1995). Aboulder conglomerate related to this fault has been correlated to theTipton Shale Member and the overlying Cathedral Bluffs Tongue(Hansen 1986). On the northern basin margin, McGee (1983), Steidtmannet al. (1983), and Pietras et al. (2003b) documented that movement on theContinental Fault occurred simultaneously with the onset of WilkinsPeak Member deposition. This reverse fault movement blocked drainagesthat had previously provided sediment and water to Lake Gosiute(Pietras et al. 2003b). Deposition of alluvial-fan conglomerate at the

FIG. 14.—Estimate of apparent cyclicity determined for the lower, middle, andupper Wilkins Peak Member plotted against the position in the basin. Error barsreflect 2s uncertainty in 40Ar/39Ar age determinations (Smith et al. 2003), anddecrease towards the basin center because the uncertainty is distributed betweenmore cycles.

FIG. 15.—Histogram of cycle thicknesses for the middle Wilkins Peak Member.

TABLE 1.—Number of cycles present and average cycle duration for portions of the lower, middle, and upper Wilkins Peak Member constrained by datedtuffs (Smith et al. 2003).

Unit Duration Upper WP 260 6190 kyr Middle WP 430 6150 kyr Lower WP 310 6190 kyr

Section Num. Cycles Ave. Duration (kyrs) Num. Cycles Ave. Duration (kyrs) Num. Cycles Ave. Duration (kyrs)

WM 27 9.6 6 7.0 45 9.6 6 3.3 41 7.6 6 4.6KA nm nm 32 13.4 6 4.7 nm nmSB nm nm 29 14.8 6 5.2 25 12.4 6 7.6MR nm nm 22-28 18.2 6 8.2 19 16.3 6 10.0AL nm nm 19-25 20.9 6 9.7 15 20.7 6 12.7BG 16 16.3 6 11.9 16 26.9 6 9.4 12 25.8 6 15.8BT 10 26 6 19.0 13 33.1 6 11.5 9 34.4 6 21.1

See Figure 11 for the stratigraphic position of dated tuffs. (nm – not measured).


eastern basin margin in response to uplift of the Granite Mountains mayalso have coincided with Wilkins Peak Member deposition, butcorrelations between the conglomerate and the Green River Formationare presently not well-defined (e.g., Love 1970).

On the basis of the above structural histories, we infer that the long-term shift from balanced-fill conditions during deposition of the Rife Bedof the Tipton Shale Member to underfilled conditions of the Wilkins PeakMember was principally a response to tectonic isolation of the basin.40Ar/39Ar age determinations (Smith et al. 2003) on tuffs in the GreenRiver Formation date this transition to between 51.25 6 0.31 and50.70 6 0.14 Ma (2s uncertainties). Differential subsidence of thegreater Green River Basin mostly ended by 49.96 6 0.08 Ma (Fig. 11).Wilkins Peak Member cycles deposited after that time (upper WilkinsPeak Member) are relatively uniform in thickness across the basin, andevaporite and fluvial sandstone beds become rare. From the timing ofreturn to balanced-fill conditions in the basin (i.e., the Laney Member), itappears that approximately 1 My or less was required to restore thesurrounding landscape to a state similar to that which existed before theonset of enhanced tectonism.

CONCLUSIONS

1. The Wilkins Peak Member is dominated by repetitive, asymmetricfacies successions that represent expansion and contraction of LakeGosiute. The lower part of each cycle is relatively thin and recordslacustrine transgression that often culminated with deposition oflaminated oil shale. The upper regressive part of each cycle isthicker and most commonly consists of palustrine facies. Evaporitefacies deposited in a salt pan commonly cap cycles near the basindepocenter. Wilkins Peak Member cycles do not readily conform toany existing sequence stratigraphic nomenclature, and insteadshould be considered as aggradational, non-Waltherian event bedsthat record rapid and comprehensive paleoenvironmental change.

2. At least 126 Wilkins Peak Member cycles are present in the WMcore near the basin center, in contrast to 77 cycles previouslyidentified using Fischer Assay data alone. The number of cyclespreserved within a specific chronostratigraphic window decreasessystematically northward (toward the basin margin), reflecting aninterplay between varying magnitudes of lake expansion anda south-dipping depositional gradient. This interplay resulted increation of a natural stratigraphic ‘‘frequency filter’’ and impliesthat the completeness of the resultant stratigraphic record variescontinuously across the basin. Evidence for complete desiccationand hiatuses is commonplace even near the basin center, suggestingthat a complete or nearly complete chronostratigraphic record maynot exist anywhere. This study indicates the need for extremecaution when interpreting periodic signals from other lake deposits;especially when based on one-dimensional analysis of cycles orwhere independent chronostratigraphic control is absent.

3. The average apparent duration of cycles near the basin center isapproximately 10,000 years, judging from dated tuff horizons(Smith et al. 2003). The true duration of these cycles is shorter,due to the ubiquity of hiatuses. Because of the wide range in cyclethickness (0.14–5.70 m), the shortest cycles most likely representsubstantially less time than the average, possibly less than1,000 years. No external driving mechanism is presently knownfor Eocene cycles of such short duration. It is therefore proposedthat some or all of the Wilkins Peak Member may be autogenic,reflecting geomorphic instability in the drainage basin thatsurrounded Lake Gosiute.

4. Evaporite deposition in the Wilkins Peak Member correspondsclosely to periods of maximum differential basin subsidence and isconcentrated in the greater Green River Basin depocenter. On the

basis of structural and stratigraphic observations, a period ofincreased tectonic uplift commenced concurrent with the lowerWilkins Peak Member. We infer that tectonic modification ofregional drainage and basin subsidence patterns exerted a primarylong-term control on evaporite deposition.

ACKNOWLEDGMENTS

We thank Meredith Rhodes-Carson, Mike Smith, Kevin Bohacs, ReubenJohnson, Brooke Swanson, Brad Singer, and Jack Neal for stimulatingdiscussions on the outcrop. Mike Smith provided insightful comments ona prior version of the manuscript. John Dyni, Cari Johnson, and RichardYuretich provided helpful reviews. This study was funded by donations fromConoco, Texaco, the American Association of Petroleum Geologists Grants-in-aid program, the J. David Love Fellowship, the Department of Geologyand Geophysics at the University of Wisconsin-Madison, by a grant from theDonors of The Petroleum Research Fund, administered by the AmericanChemical Society, and by National Science Foundation grants EAR-9972851(to Singer) and ATM-0081852 (to Singer and Carroll).

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Received 3 October 2005; accepted 9 March 2006.


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