Paleohydrological change during the Early-Middle Eocene: Insights from Long-lived Green River Formation

doi:10.21203/rs.3.rs-1503204/v1

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Paleohydrological change during the Early-Middle Eocene: Insights from Long-lived Green River Formation

https://doi.org/10.21203/rs.3.rs-1503204/v1

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The evolution of Earth’s climate and oceans during the Early Eocene Climatic Optimum (EECO) has been a focus of extensive research, but the continental response to global climate change in this warmer than modern world is poorly understood. This study quantitatively reconstructs terrestrial paleoenvironment and paleohydrological changes as recorded in the early–middle Eocene (52.7‒43.8 Ma) lacustrine sedimentary succession of the Green River Formation exposed in Indian Canyon, Uinta Basin, northern Utah, USA. Detailed lithofacies variations and whole rock elemental and mineralogical data reflect lake level changes over the 1000 m-thick succession. By comparing the lithofacies and the major element compositions and clay mineralogies, we show that Ca/Al ratio can be utilized as a proxy of lake level, Mn/Fe ratio as a proxy of lake bottom-water redox conditions, and K/Al ratio and kaolinite abundance as proxies of chemical weathering rates. Based on principal component analysis of the major elements and mineralogy, the first principal component is interpreted to represent paleo-hydrological changes, involving variations in the lake level and degree of chemical weathering. Comparison of this record of paleohydrological change for the mid-latitude region of North America with stable δ¹⁸O and δ¹³C data for benthic foraminifera, which reflect global climatic change, indicates wetter periods correspond to intervals of lighter δ¹⁸O and heavier δ¹³C values, and vice versa. These results show that the extremely organic-rich “Mahogany Zone” in this sequence corresponds to the transitional period between the EECO and cooler middle Eocene (ca. 48.7 Ma). This implies that climatic instability and resultant environmental stress during this transitional climate state led to massive blooms of Botryococcus green algae in Lake Uinta, the remains of which are the source of organics preserved in the Mahogany Zone.

Eocene

hothouse

lacustrine

oil–shale

mid-latitude

hydrological change

The latest Paleocene and early Eocene spanned the hottest interval (> 10°C warmer than pre-industrial times) of the Cenozoic, with episodic warming events such as the Paleocene Eocene Thermal Maximum (PETM; 56.0 Ma; Zachos et al. 2001; 2008) and Early Eocene Climatic Optimum, which hosted multiple hyperthermal events (EECO; 53.3‒49.1 Ma; Westerhold et al. 2018a). This “hothouse” period was characterized by elevated atmospheric CO₂ levels (> 1000 ppm; Beerling and Royer 2011; Anagnostou et al. 2020) and a shallower carbonate compensation depth (CCD; Palike et al. 2012; Harper et al. 2020), and was followed by the transition into “icehouse” conditions and development of Antarctic ice sheets during the late Eocene (Zachos et al. 2001, 2008; Burke et al. 2018; Hollis et al. 2019; Westerhold et al. 2020). Under the most extreme projections (RCP8.5) by the Intergovernmental Panel on Climate Change (IPCC Sixth Assessment Report), by AD 2200 atmospheric pCO₂ will approach 2000 ppm with an associated increase in global surface temperatures of 3–10.5°C. Since this predicted temperature is comparable to that of the PETM and EECO, numerous general circulation models (GCM) have reconstructed past environmental change, such as precipitation and temperature, during these intervals, considered analogous to our warming climate over the next 200 years (Winguth et al. 2010; Carmichael et al. 2017; Hutchinson et al. 2018; Inglis et al. 2020; Lunt et al. 2021; Shields et al. 2021). General circulation models simulated the increased CO₂ concentrations of the PETM and suggest decreased precipitation in the northern mid-latitude (30°‒50°N) regions, especially in North America (e.g., Winguth et al. 2010; Carmichael et al. 2017) but with increased rainfall observed in high-latitude regions (> 50°N) (Carmichael et al. 2017). Other high resolution climate simulations reflect changes in the location and intensity of mid-latitude storm tracks during the extreme transition from the late Paleocene to PETM (Shields et al. 2021), including the northward movement of the jet stream and atmospheric river precipitation intensity in western North America (40°‒42°N). During the PETM, regions of enhanced terrestrial warming were closely associated with geological evidence for increasingly dry conditions, whereas regions with less terrestrial warming generally became wetter (Hutchinson et al. 2021). Thus, the simulations are in conflict and empirical evidence from geological archives such as lake sequences are needed to resolve the discrepancies and test the models.

Widespread Early–middle Eocene age lacustrine oil–shale deposits comprise proxy records of continental climate change during this critical interval. Famous deposits of this kind include: the Green River Formation in the western United States; the Messel oil shale in Germany; and the Fushun, Huadian, and Xining oil–shale deposits in China (Birdwell et al. 2016; Birgenheier et al. 2019; Chen et al. 2017; Cumming et al. 2012; Hu et al. 2019; Lenz et al. 2007; Lenz and Wilde 2018; Meijer et al. 2019; Sayem et al. 2018; Smith et al. 2008, 2010; Strobl et al. 2015; Sun et al. 2013; Yang et al. 2019). The widespread occurrence of these petroleum source rocks indicates that at least transient high humidity characterized mid-latitude regions of the Northern Hemisphere during this period, which is inconsistent with some of the climate simulations (e.g., Winguth et al. 2010; Carmichael et al. 2017). Among these lacustrine records, the Green River Formation hosts the longest sedimentary record, preserves a sensitive continental environmental response to the EECO and post-EECO (Gall et al. 2017; Birgenheier et al. 2019; Elson et al. 2022), and benefits from detailed age constraints from ⁴⁰Ar/³⁹Ar ages on intercalated tuffs (Smith et al. 2008, 2010). Thus, the combinations of long depositional duration and many dated levels, makes the Green River Formation an ideal venue to explore the continental proxy record of the warmest interval of the Cenozoic.

In this study, we present detailed lithofacies variations and whole-rock elemental and mineralogical data for a ~ 9 Myr record (53‒44 Ma; dates from Smith et al. 2008) over the early-middle Eocene from the Green River Formation in the western Uinta Basin, Utah, USA. Our data quantitatively reconstruct the paleoenvironmental and paleohydrological changes during this greenhouse interval. We also extracted the periodicities in lithofacies variations, which were plausibly paced by astronomical cycles, and established an orbitally tuned age model. Finally, we discuss the relationship among our regional changes to global archives recorded in other lacustrine records and by δ¹⁸O and δ¹³C of marine benthic foraminifera data, and the calcium carbonate compensation depth (CCD).

Geological Setting

The Green River Formation was deposited in a network of continental interior lakes that developed in a fragmented foreland province laterally spread eastward of the Cordilleran thrust and fold belt. The early-middle Eocene deposits are widely distributed across the western US, namely the Uinta Basin in northern Utah, the Greater Green River Basin in Wyoming, and the Piceance Creek Basin in Colorado (Carroll et al. 2006; Whiteside and Van Keuren 2009; Tänavsuu-Milkeviciene and Sarg 2012; Birdwell et al. 2019). The basins are divided by volcanic chains of topographical anticline highs associated with Laramide Orogeny uplift, and were intermittently active during the Cretaceous and Eocene (Dickinson et al. 1988; Smith et al. 2008). Paleomagnetic studies indicate that the paleolatitude of the United States throughout the Cenozoic was similar (40°‒45°N) to the present-day (Kent and Irving 2020).

⁴⁰Ar/³⁹Ar dating of several intercalated tuffs constrains deposition of the Green River Formation in the Greater Green River Basin to ~ 53.0‒48.3 Ma (Smith et al. 2008; 2010). In the Uinta and Piceance Creek basins, the Carbonate Marker Unit, a brackish interval deposited in the lower Green River Formation, is radioisotopically dated to ~ 54.0 Ma in a tuff layer near the base of the unit (Remy 1992). The age of uppermost lacustrine sedimentation in the Uinta Basin is ~ 44.0 Ma, based on ⁴⁰Ar/³⁹Ar dating of the Strawberry Tuff (Smith et al. 2008) as the basin experienced infilling from the east to west (Bohacs et al. 2000). The age of the youngest strata in the Piceance Creek Basin is less resolved, however Smith et al. (2008) have suggested ca. 47.5 Ma.

Prior sedimentological studies of the Green River Formation, in the Greater Green River Basin, have identified widespread river inflow during ca. 53.0‒51.5 Ma, after which Lake Gosiute transitioned into an evaporative environment until 50.5 Ma. Lake Gosiute freshened into a fluctuating profundal lake at ca. 49.5 Ma, before the capture of the Idaho River system resulted in widespread deposition of volcaniclastic rocks at ca. 48.0 Ma (Pietras et al. 2003; Smith et al. 2008). In the Piceance Creek Basin, a lacustrine environment established at ~ 51.5 Ma before an evaporative environment prevailed at ~ 50.5 Ma. The lake expanded again at ca. 48 Ma before finally infilling with volcaniclastic material (Smith et al. 2008).

Towards the south, the long-lived Lake Uinta occupied the Uinta and Piceance Creek basins, which were partially separated by a structural high named the Douglas Creek Arch (Cashion, 1967; Carroll and Bohacs, 1999). The Uinta Basin contains the longest sedimentation record of the Green River Formation-hosting basins and is the focus of this study. The Indian Canyon section of the Uinta Basin is ~ 1000 m thick and encompassed a depositional period of ~ 9 Myr (Smith et al. 2008). The Green River Formation in this section is located above the fluvial Colton Formation (late Paleocene to early Eocene age) and below the alluvial-fluvial Uinta Formation (late Eocene age) (Cashion, 1967; Carroll and Bohacs, 1999; Smith et al. 2008). Several ⁴⁰Ar/³⁹Ar ages have been obtained for intercalated tuffs in industry cores and field sites by Smith et al. (2008), which can be correlated to the Indian Canyon section: Curly tuff = 49.02 ± 0.89 Ma; Wavy tuff = 48.37 ± 0.86 Ma; Fat tuff = 46.34 ± 0.80 Ma; Portly tuff = 45.58 ± 0.79 Ma; Oily Tuff = 45.14 ± 0.78 Ma; Strawberry tuff = 44.00 ± 1.19 Ma (Fig. 1). Multiple sandstone dominant levels containing lighter organic C isotope ratios found above and below the D Marker within the Sunnyside Delta Interval, have been proposed to represent the expression of hyperthermal events at ca. 52 Ma (Gall et al. 2017; Birgenheier et al. 2019; Westerhold et al. 2018a). Based on these absolute ages and carbon isotopic data, deposition of the lacustrine sediments in Indian Canyon occurred between ~ 53 and ~ 44 Ma (Smith et al. 2008, 2010; Tsukui and Clyde 2012; Birgenheier et al. 2019), and the average sedimentation rate was ~ 10–15 cm/kyr (Smith et al. 2008).

The lower Green River Formation of the Uinta Basin (ca. 53‒49 Ma) fluctuates between fluvial and lacustrine environments, and alluvial fans may have extended into some parts of the basin (Smith et al. 2008; Birgenheier et al. 2019). At ca. 48.7 Ma, an extremely organic-rich unit called the “Mahogany Zone” developed in the Uinta and Piceance Creek basins during the lake highstand. This key unit is thought to have formed as a result of increasing salinity of the lake water in both basins (Smith et al. 2008; Vanden Berg and Birgenheier, 2017). In the upper Uinta Basin (~ 49‒45 Ma), increased volcanic activity and inflow of volcaniclastic debris resulted in the lake level to shallow (Smith et al. 2008). In this low lake level stage, Lake Uinta was an evaporative-dominated environment, in which exotic, high atmospheric pCO₂ indicative minerals such as nahcolite formed (Jagniecki et al. 2015). The final stages of Green River Formation deposition in the Uinta Basin (~ 45.0‒43.5 Ma) are the most sparsely documented. Preserved strata indicate that a shallow lake occupying the Uinta Basin had become fluvial-lacustrine by this time before infilling from the east to west with alluvial material from the Uinta Formation (Bohacs et al. 2000; Smith et al. 2008; Vanden Berg and Birgenheier, 2017).

The lakes of the Green River Formation have been argued to record a range of temporal variability from orbital pacing (20 − 2,400 kyr) and sub-orbital periodicities at the quasi-decadal (8–18 year) and sub-decadal (2–8 year) timescales (Sayles 1922; Bradley 1929; Fischer and Roberts 1991; Ripepe et al. 1991; Cole 1998; Machlus et al. 2008; Meyers 2008; Whiteside and Van Keuren, 2009; Aswasereelert et al. 2013) despite the tectonic influence that usually persist through long-lived lacustrine basins, including the Green River Formation (Kelts, 1988; Pietras and Carroll, 2006). The extensive sedimentary record available at Indian Canyon affords the ability to reconstruct past lake-level change and to interrogate the presence of orbital signals in a continental interior site over the early-middle Eocene.

Lithofacies descriptions

To semi-quantitatively reconstruct the lake level changes, we classified the lithofacies according to grain size, color (related to the organic matter and carbonate contents), mineralogy, and sedimentary structures (Fig. 2), modified after “Depth Ranks” used by Olsen (1984, 1986, 1990) for the lacustrine Newark-Hartford basin successions of early Mesozoic age. Lithofacies types were set to reflect differences in the lake levels and degree of reduction (i.e., stratification) and oxidation (i.e., oxygenation of water column/desiccation) of the lake. These types were each assigned a numerical facies code from 0 to 6, arranged in order of inferred increasing water depth. Facies 0 is a massive reddish and dark greenish siltstone with a mottled structure, representing a paleosol deposited during the lowest lake level. Facies 0.5 is sandstone with ripple marks formed by longshore currents or fluvial channel inflows, with occasional sedimentary beds containing hummocky structures caused by high-energy wave action. Facies 1 is sandy siltstone with evaporite minerals and facies 1.5 is dolomitic sandstone. Facies 2 is dolomite, and facies 2.5 is dolomitic siltstone. Some of facies 1.5‒2.5 have poorly developed laminations and yield fish fragments and plant fossils. Facies 3 is carbonaceous gray siltstone, and facies 3.5 is carbonaceous brown siltstone. Facies 4 is brown claystone, and facies 4.5 is olive claystone. Facies 5 is dark gray claystone, and facies 5.5 is dark gray claystone with weak laminations. Facies 6 is organic-rich dark brown-black mudstone with fine laminations.

Major elements, mineralogy, and principal component analysis

A total of 662 samples of each lithology was collected from the 1000-m-thick sedimentary succession at 1‒2 m intervals. X-ray fluorescence (XRF) spectrometry analyses were conducted to determine the major element contents using a Rigaku ZSX Primus II instrument equipped with a Rh X-ray tube operated at 50 kV and 60 mA, in the Graduate School of Environmental Studies, Nagoya University, Nagoya, Japan. Glass beads were prepared by mixing each sample, which was ignited to 950°C to decompose carbonates, with anhydrous lithium tetraborate flux and subsequent fusion. The XRF data were calibrated by analyses of rock reference standards issued by the Geological Survey of Japan (GSJ; Geochemical Reference Sample Data Base). Data are reported on a volatile-free basis.

The mineral compositions of 468 powdered samples were determined by X-ray diffraction analysis using a PANalytical, X'Pert PRO MPD operated at 45 kV and 40 mA with Cu Kα radiation, at the Center for Advanced Marine Core Research, Kochi University, Kochi, Japan. The elemental and mineralogical data were subjected to principal component and paleothermometry analysis. For the principal component analysis, the element contents were divided by the Al contents and the mineral abundances were divided by the illite contents to calculate the Z-score normalization:

Z = [x – mean(x)/stdev(x)].

The obtained Z-scores were subjected to principal component analysis using the software package PAST (Hammer et al. 2001). Mean annual temperatures (MATs) were estimated from the paleosol weathering index (PWI = 100 × [(4.20 × Na) + (1.66 × Mg) + (5.54 × K) + (2.05 × Ca)]) and calculated from the equation (standard error of ± 2.1℃):

T (℃) = -2.74 × ln(PWI) + 21.39 (Gallagher and Sheldon 2013).

Timeseries analysis

Our timeseries analysis followed the following five steps to convert thickness to time.

1) We divided the section into four intervals based on lake level stages. Each of these was independently subjected to Fourier analysis in which we derived the main spectral frequencies in thickness. Dominant thickness periodicities in the lithofacies for each interval (Fig. 3; 4) were determined with the software Acycle (Li et al. 2019). Variations in the lithofacies types were subjected to wavelet power spectrum analysis in Acycle, and band-pass filtering analysis using a Gaussian window using AnalySeries (Paillard et al. 1996), to extract the relative frequencies of the dominant spectral peaks. 2) Next, we used published ages (Smith et al. 2008) on intercalated tuffs to obtain durations and thus accumulation rates for each of these intervals. 3) Using these accumulation rates, we picked the thickness frequencies closest to the 405 kyr, g5-g2 (Jupiter and Venus) eccentricity cycle for each of the intervals, and filtered to these frequencies. 4) We then tuned each of the four sections to a pure 405 kyr cycle. 5) Finally, we concatenated the four sections, now in time, and registered the composite to the dated tuffs, thus placing our lithological section into a time framework and producing our age model (Fig. 5). These steps are described in more detail as follows.

The duration of the obtained spectral peaks was estimated based on the average sedimentation rate calculated from the radioisotopic ages of intercalated tuffs (Smith et al. 2008). When the obtained periodicity corresponded to an orbital component, we used the 405 kyr cycle as the tuning target, and converted the periodicity of the lithofacies changes into a constant duration of 405 kyr in the time domain (Fig. 6). Filtering results were adjusted to better conform to the mid points of the interpreted 405 ky cycles.

Stage subdivision based on the lithofacies changes

Based on the lithofacies, we subdivided the Indian Canyon section into the following nine lake stages: (1) fluvio-lacustrine stage; (2) deep lake stage; (3) fluvio-lacustrine stage; (4) highly fluctuating lake stage; (5) shallow lake stage; (6) evaporative lake stage; (7) shallow lake stage; (8) fluctuating lake stage; (9) fluvio-lacustrine stage (in ascending order; Fig. 1). A comparison between our subdivisions and the stratigraphic classifications of previous studies (Smith et al. 2008; Birgenheier et al. 2019) are shown in Fig. 1.

Stage 1 is the first fluvio-lacustrine stage (0‒210 m stratigraphic level). We assigned two thick (> 6 m) sandstone beds to be the basal part of this stage, which correspond to the Carbonate Marker Unit (Birgenheier et al. 2019). The top of this stage is defined by a 3-m-thick sandstone bed at a stratigraphic level of 210 m, which corresponds to the D Marker (Birgenheier et al. 2019). This stage consists mainly of alternating olive claystone, light gray dolomite, and reddish brown or greenish gray paleosols associated with a fluvial channel sandstone. Based on the regular occurrence of fluvial channel sandstones and paleosols, the paleoenvironment during this stage was a repetition of fluvial-lacustrine facies. Given the study site’s southwestern basinal margin location, the depositional setting of this stage may have been a lakeshore environment.

Stage 2 is the deep lake stage (210‒310 m stratigraphic level). We assigned a reddish brown paleosol layer at a stratigraphic level of 310 m to be the top of this stage, which corresponds to the C Marker (Birgenheier et al. 2019). This stage consists mainly of alternating beds of light gray dolomite, olive claystone, and organic-rich brown-black mudstone. The periodic alternation of olive claystone and dolomite is interpreted to reflect fluctuating lake levels. The presence of organic-rich brown-black mudstone and absence of sandstone suggest deposition during high lake level. Stages 1 and 2 correspond to the Sunnyside Delta Interval (Birgenheier et al. 2019; French et al. 2020).

Stage 3 is the second fluvio-lacustrine stage (310‒420 m stratigraphic level). We assigned the 1-m-thick sandstone beds at a stratigraphic level of 420 m to be the top part of this stage, which correspond to the S1 Sandstone (Birgenheier et al. 2019). This stage also consists mainly of alternating olive claystone, light gray dolomite, and reddish brown or greenish gray paleosols associated with a fluvial channel sandstone. Due to the frequent occurrence of fluvial channel sandstone, hummocky sandstone, and paleosols, the paleoenvironment during this stage was fluvial-lacustrine with associated storm-related waves. In this stage, alternations of dolomite and mudstone have periodicities of ∼1.7 m (Fig. 3a). This stage corresponds to the Transitional Interval of the Douglas Creek Member (Birgenheier et al. 2019; French et al. 2020).

Stage 4 is the highly fluctuating lake stage (420‒520 m stratigraphic level). We assigned the organic-rich mudstones appearing at a stratigraphic level of 520 m (at the summit of the Indian Canyon section) as the top of this stage. This consists mainly of alternating beds of light gray dolomite, olive gray claystone, and organic-rich mudstones. In this stage, alternations of dolomite and mudstone shows clear periodicities of ∼1.1 m (Fig. 3b). The lower part of this stage includes the Mahogany Zone, which is the most organic-rich unit in the Green River Formation and consists of dark brown-black mudstone with thin laminations (Fig. 2l) with up ∼45 wt.% organic C (Bradley 1931; Tuttle and Goldhaber 1993; Whiteside and Van Keuren, 2009). The Mahogany Zone is thus considered to have been deposited in a deep and stratified lake with reducing bottom-water conditions (Bradley 1931; Tänavsuu-Milkeviciene and Sarg 2012; Vanden Berg and Birgenheier, 2017). Clear periodic alternations of dolomite and mudstone indicate a highly fluctuating lake level, which was likely due to precipitation changes. The lake system was likely fluctuating profundal in nature and had a balanced inflow and outflow of lake water and sediments. Thin beds of chert are occasionally interbedded in this stage, suggesting that the lake water was gradually becoming alkaline. This stage corresponds to the lower part of the R-8 and Mahogany Zone of the Parachute Creek Member (Birgenheier et al. 2019; French et al. 2020).

Stage 5 is the shallow lake stage (520‒740 m stratigraphic level). We assigned the carbonaceous siltstone and overlying orange siltstone bed appearing at a stratigraphic level of 740 m as the top of this stage. This stage consists mainly of alternations of light gray dolomite and grayish carbonaceous siltstone, suggestive of a low lake level with limited fluctuations. The presence of hummocky cross stratification in a dolomitic sandstone layer also indicates a low lake level. This stage contains nahcolite minerals generally formed in alkaline lake environments (Buchheim 1994; Smith et al. 2008; Jagniecki et al. 2015). In this stage, alternations of dolomite and mudstone shows clear periodicities of ∼5 m (Fig. 4a). As noted by Smith et al. (2008), the numerous tuff layers in this stage suggest that nearby volcanic activity was occurring. This stage corresponds to the upper part of the R-8 to Hypersaline Zone phase 2 of Birgenheier et al. (2019) and French et al. (2020).

Stage 6 is the evaporative lake stage (740‒810 m stratigraphic level). We assigned a 10-m-thick orange siltstone bed at a stratigraphic level of 810 m as the top part of this stage. This stage consists mainly of interbeds of dolomite, carbonaceous gray siltstone, and sandy siltstone. In particular, it contains a large amount of evaporite minerals, such as gypsum, nahcolite, and trona (Eugster and Surdam 1973; Lundell and Surdam 1975; Buccheim 1994). The occurrence of mud crack structures suggests that the lake was in an evaporative environment. The sandy siltstone found in this stage was likely deposited in this evaporative setting near the lakeshore. This stage corresponds to the lower part of the saline facies (Hypersaline Zone phase 3; Birgenheier et al. 2019; French et al. 2020).

Stage 7 is the shallow lake stage (810‒900 m stratigraphic level). We assigned the carbonaceous siltstone and overlying brownish claystone bed appearing at a stratigraphic level of 900 m as the top part of this stage. This stage consists mainly of alternations of light gray dolomite and grayish carbonaceous siltstone, indicative of a low lake level environment. Due to the abundant occurrence of chert layers, this stage was characterized by the presence of alkaline-rich water. This stage corresponds to the upper part of the saline facies of the Parachute Creek Member (hypersaline Zone phase 3; Birgenheier et al. 2019; French et al. 2020).

Stage 8 is a fluctuating lake stage (900‒975 m stratigraphic level). We assigned the brownish claystone and overlying carbonaceous siltstone beds, including channel sandstone, appearing at a stratigraphic level of 900 m as the top part of this stage. This stage consists mainly of alternating dolomite and brown claystone, and the paleoenvironment was characterized by fluctuating lake levels. In this stage, the alternations of dolomite and mudstone show clear periodicities of ∼3 m (Fig. 4b). The occurrence of brown mudstone, but no organic-rich brown-black mudstone, suggests that the lake level during this stage was deeper than in the shallow lake stage and shallower than the deep lake stage. Rocks formed in this stage tend to be silicified, and bedded chert and silica concretions occur frequently throughout the section. The paleoenvironment during this stage is interpreted to have been the most alkaline and oxidizing lake setting, in which organic decomposition occurred (Kuma et al. 2019; Yoshida et al. 2021). This stage corresponds to the lower part of the sandstone–limestone facies of the Parachute Creek Member (Birgenheier et al. 2019; French et al. 2020).

Stage 9 is the third fluvio-lacustrine stage at the top of the section (975‒1000 m stratigraphic level). It is overlain by the Uinta Formation at the end of the Indian Canyon section near Duchesne. This stage consists of fluvial channel sandstone intercalated with some coal seams. In this stage, only a few chert layers occur, suggesting that the alkalinity of the lake water had decreased. Smith et al. (2008) also noted that the uppermost part of the section may have been formed in freshwater lakes associated with river inflow. Due to the occurrence of fluvial channel sandstone and thin coal seams, the paleoenvironment during this stage was fluvial-lacustrine. This stage corresponds to the upper part of the sandstone–limestone facies of the Parachute Creek Member (Birgenheier et al. 2019; French et al. 2020).

Changes in major element chemistry and mineralogy

To quantitatively reconstruct the paleoenvironmental changes, we focused on Ca/Al, Mn/Fe, K/Al, Si/Al, and Ti/Al ratios. The Ca/Al ratio is used as an indicator of lake level (Hou et al. 2017; Sun et al. 2019). The dolomite has the highest CaO content, followed by the siltstone and organic-rich mudstone. Microscopic observations revealed that the dolomite was primarily precipitated in a shallow lake setting (Last 1990; Wright and Wacey 2005). The Mn/Fe ratio is commonly used as a proxy of bottom-water redox conditions (Naeher et al. 2013). Dolomite samples have high Mn/Fe ratios, and the organic-rich mudstone samples have low Mn/Fe ratios. Given that minerals formed by weathering, such as kaolinite and smectite, are Al-rich, and that clay minerals, such as illite, are K-rich, the K/Al ratio are used as proxies of chemical weathering (Beckmann et al. 2005). TiO₂ and Al₂O₃ are considered to be stable during chemical weathering and diagenesis, because of their low solubility and lack of involvement in biogeochemical processes. As such, the Ti/Al ratio is commonly used as a proxy of provenance change (Hayashi et al. 1997; Sun et al. 2019). In the Indian Canyon section, TiO₂ and Al₂O₃ are positively correlated, suggesting that the source rock did not change significantly.

Logratio-transformed data tend to follow a multivariate normal distribution (e.g., Aitchison and Shen 1980; Ohta and Arai 2007); we thus examined the relationships among ln(Ca/Al), ln(Mn/Fe), ln(K/Al), and ln(Si/Al) (Fig. 7). Ln(Ca/Al) and ln(Mn/Fe) exhibit a clear positive correlation, with high ln(Ca/Al) and ln(Mn/Fe) values for the dolomite samples and low ln(Ca/Al) and ln(Mn/Fe) values for the organic-rich brown-black mudstone samples. This observation indicates that variations in lake level and bottom-water redox conditions were the primary controls on the changes in lithofacies. ln(Ca/Al) and ln(K/Al) also exhibit a general positive correlation, with high ln(Ca/Al) and ln(K/Al) values for the dolomite samples and low ln(Ca/Al) and ln(K/Al) values for the organic-rich brown-black mudstone samples. Some dolomite samples show anomalously high ln(Ca/Al) but low ln(K/Al) values. These low-ln(K/Al) dolomite samples also have high ln(Si/Al) values. Given that the sandstone samples are different, with lower ln(Ca/Al) and slightly higher ln(Si/Al) values, excess enrichment of Si in some dolomite samples was caused by silica precipitation due to organic matter decomposition and the resultant pH decrease of the lake water (Kuma et al. 2019).

Stratigraphic changes in Ca/Al (lake level proxy), Mn/Fe (bottom-water redox proxy), K/Al (chemical weathering proxy), Si/Al (grain size and Si enrichment), and Ti/Al ratios are shown in Fig. 7. Stratigraphic changes in the mineralogy are shown in Fig. 8. Ca/Al, K/Al, and Si/Al ratios are all higher between stages 5 and 8 (i.e., the shallow and evaporative lake). In these stages, smectite, calcite, and dolomite were the dominant minerals. Higher Mn/Fe ratios and the presence of gypsum only occurred during stage 6 (i.e., evaporative and oxic conditions). In contrast, lower K/Al and Si/Al ratios and abundant kaolinite characterized stage 1 to the lower part of stage 4, when the lake level was higher. In stages 1 to 3, in which there are interbedded paleosols, the paleo-temperature estimated from the PWI is 10°C‒12°C ± 2.1°C.

The results of the Z-score principal component analysis of the Al-standardized major element data and illite-standardized mineralogical data are shown in Supplementary Fig. S2. Principal component (PC) 1 represents components with high values in a dry environment, such as Si/Al, Ca/Al, K/Al, calcite/illite, and dolomite/illite, whereas negative values are associated with a wet environment (e.g., kaolinite/illite). As such, changes in PC1 (28% of the total variance) are interpreted to primarily reflect the paleo-hydrological variations. PC1 tends to decrease slightly from stage 1 to 2, and increase slightly from stage 2 to 4. Values tend to be high from the upper part of stage 5 to stage 7, and decrease again from stage 7 to 9 (Fig. 9).

Time series

The main finding for the spectral analysis is that the lithological and geochemical variations are consistent with orbits pacing. The power spectrum of the composite depth rank series shows a number of well-defined peaks (Supplementary Fig. S3‒S6) around 7.7, 25, and 50 m/cycle in stages 1 and 2. The average sedimentation rate of ~ 7.0‒14.1 cm/kyr was calculated from the age difference of 2.91 ‒5.87 Myr in the 410-m-thick succession between the basal age of the section (54‒53 Ma; Remy 1992) and Curly tuff (49.02 ± 0.89 Ma; Smith et al. 2008). Based on this average sedimentation rate, the obtained periodicities (7.7, 25, and 50 m/cycle) are equivalent to 55‒110, 177‒357, and 355‒714 kyr/cycle, respectively. The averaged ~ 50 m/cycle of mudstone to sandstone cycles (355‒714 kyr/cycle) corresponds to the 405 kyr eccentricity cycle (Supplementary Fig. S3).

There are strong bands in spectral power consistently present around 7, 15, 23, and 35 m/cycle in stages 3 and 4. The average sedimentation rate of 3.5‒6.1 cm/kyr was calculated from the age difference of ~ 650 kyr of the 40-m-thick succession between the Curly tuff (49.02 ± 0.89 Ma) and Wavy tuff (48.37 ± 0.86 Ma; Smith et al. 2008) and lamination thickness counting of ~ 35-µm-thick annual couplets (Fig. 2l). Based on this average sedimentation rate, the obtained periodicities (7, 15, 23, and 35 m/cycle) are equivalent to 114‒200, 246–429, 377‒657, and 574‒1000 kyr/cycle, respectively. We assumed that the averaged ~ 23 m/cycle containing fine-to-coarser-back to fine cycles (377‒657 kyr/cycle) correspond to the 405 kyr eccentricity cycle (Supplementary Fig. S4). In this case, the 1.7 m periodicity shown in Fig. 3a corresponds to ~ 28‒49 kyr, and the 1.1 m periodicity shown in Fig. 3b relates to a ~ 18‒31 kyr cycle, which likely corresponds to the precession cycle.

Strong bands in spectral power are consistently present around 28 and 63 m/cycle in stages 5 and 6. The average sedimentation rate of 8.8‒27 cm/kyr was calculated from the age difference of ~ 1.59‒4.87 Myr of the 430-m-thick succession between the Wavy tuff (48.37 ± 0.86 Ma) and Oily tuff (45.14 ± 0.78 Ma; Smith et al. 2008). Based on this average sedimentation rate, the obtained periodicities (28 and 63 m/cycle) are equivalent to 104‒318 and 230‒716 kyr/cycle, respectively. The ~ 63 m periodicity (230‒716 kyr/cycle) related to siltstone-carbonate cycles likely corresponds to the 405 kyr eccentricity cycle (Supplementary Fig. S5). The 5-m periodicity shown in Fig. 4a corresponds to ~ 18‒56 kyr, which is likely the precession or obliquity cycle.

There are strong bands in spectral power consistently present around 30 and 51 m/cycle in stages 7, 8, and 9. The average sedimentation rate of ~ 5.1‒14 cm/kyr was calculated from the age difference of 1.14‒3.11 Myr for the 160-m-thick succession between the Oily tuff (45.14 ± 0.78 Ma) and Strawberry tuff (44.00 ± 1.19 Ma; Smith et al. 2008). Based on this average sedimentation rate, the obtained periodicities (30 and 51 m/cycle) are equivalent to 214‒588 and 364‒1000 kyr/cycle, respectively. The ~ 30 m periodicity (214‒588 kyr/cycle) and the ~ 51 m periodicity (364‒1000 kyr/cycle)of brown claystone-grey siltstone- brown claystone cycles, correspond to the 405 kyr eccentricity cycle (Supplementary Fig. S6). The ~ 3 m periodicity shown in Fig. 4b corresponds to ~ 21‒59 kyr, which is likely the precession or obliquity cycle.

The estimated sedimentation rate was lower in stages 1 to 4, higher in stages 5 and 6, and then lower again in stages 7 to 9, as expected for higher/lower sedimentation rates within shallower/deeper lake stages. Wavelet analysis of lithological data of facies variations reveals amplitude modulation of the 405 kyr cycle by ~ 1.8 Myr periodicity, congruent with marine records for the early–middle Eocene (Westerhold et al. 2012; 2013).

Periodicities and accumulation rates obtained in this study are in line with previous studies from the Uinta Basin (Whiteside and Van Keuren, 2009) that obtained periodicities of ~ 33 and 40.5 m with sedimentation rates of 9.9‒12.0 cm/kyr for drillcores in the eastern Uinta Basin that correspond to the eccentricity cycle. Findings also corroborate periodicities observed in the Greater Green River Basin (Meyers 2008) at ca. 52‒49 Ma. Based on an estimated average sedimentation rate of 16.95 cm/kyr, these periodicities are equivalent to 18.8, 22.6, 39.9, 52.1, 94.8, 123.8, and 404.2 kyr, which correspond to precession, obliquity, and eccentricity cycles, respectively (Meyers, 2008). Oscillations in lacustrine highstand, lowstand, and alluvial settings correspond to 20‒30-m-thick precession-modulated (100 kyr) summer insolation changes (Smith et al. 2014). The lake was an alluvial setting during the eccentricity-driven insolation minima, and in lowstand and highstand settings during the insolation maxima (Smith et al. 2014), which is consistent with the interpretations of the present study.

Temporal changes in paleoenvironmental setting and relationship to global climate

Based on our time-series analysis, stages 1‒4 correspond to ca. 52.8 to 47.4 Ma, while stages 5‒9 correspond to ca. 47.4 to 43.7 Ma (Fig. 9). A comparison with global climatic records, including carbon and oxygen isotopic data for benthic foraminifera, reconstructions of precipitation and chemical weathering in China, and variations in the CCD are shown in Fig. 10. The higher lake level inferred from the low Ca/Al and low Mn/Fe ratios, and enhanced chemical weathering inferred from the low K/Al ratios and kaolinite abundances between stage 1 and the lower part of stage 4 (i.e., the Mahogany Zone) all correspond to the warm EECO (e.g., lighter δ¹⁸O values; Zachos et al. 2008; Westerhold et al. 2018a). Reconstructed mean annual precipitation (MAP) from a terrestrial sedimentary sequence in the Fushun Basin (Chen et al. 2017) and the chemical index of weathering (CIW) from a terrestrial sedimentary sequence in the Xining Basin (Sayem et al. 2018) also indicate higher precipitation and enhanced weathering occurred at this time. Interestingly, the disappearance of kaolinite in the Green River Formation occurred at the same time as the decrease in MAP in the Fushun Basin (Chen et al. 2017) (Fig. 10). In addition, the highest CIW value in the Xining Basin occurred at ca. 51 Ma (Sayem et al. 2018) and corresponds to the deep lake stage (stage 2) of the Green River Formation. The reconstructed paleotemperatures (10°C‒12°C) obtained from paleosol mudstone samples in stages 1‒3 (52‒49 Ma) are slightly higher than the reconstructed paleotemperatures (9.5°C‒12.0°C) obtained from the Greater Green River Basin, Wyoming (Supplementary Fig. S7; Hyland et al. 2016). Furthermore, heavier δ¹³C values of benthic foraminifera (Westerhold et al. 2020) and shallowing of the CCD (Palike et al. 2012) also appear to correspond to the deeper lake stages (i.e., stage 2 and the lower part of stage 4). Increased pCO₂ and climatic warmth resulted in ocean acidification and a decrease in pH values of seawater, which led to a shallower CCD. Increased precipitation due to climatic warming likely resulted in enhanced oceanic productivity through increased terrestrially derived nutrient input, which thus lowered the δ¹³C values of benthic foraminifera. These lines of evidence suggest there was a possible link between increased terrestrial humidity, climatic warmth, enhanced ocean productivity, and shallowing of the CCD during the EECO.

The lower lake level inferred from the high Ca/Al and Mn/Fe ratios, and weaker chemical weathering inferred from the high K/Al ratios and occurrence of smectite between the upper part of stage 4 to stage 8 correspond to cooler post-EECO conditions. Given that kaolinite and smectite were not detected, but gypsum occurs and the highest K/Al ratio is present in stage 6 (46.2‒45.7 Ma), chemical weathering was weakest in this stage, and an arid environment prevailed (Fig. 10). The high Si/Al ratios likely reflect deposition of coarser sediment in an evaporative lake setting. In addition, the abundant chert layers and silica enrichment of dolomite samples in stages 5‒8 imply increased alkalinity of the lake water due to the evaporative conditions (Kuma et al. 2019). High and irregular spikes of the Ti/Al ratio reflect input of detrital grains with a different provenance. The subsequent occurrence of several coal seams in the uppermost part of stage 9 (44.2–43.7 Ma) suggests that the paleoenvironment changed back to a wetter fluvio-lacustrine system, corresponding to the time when global climate warmed again prior to the Middle Eocene Climatic Optimum.

Variations in PC1 obtained from the principal component analysis also show a similar trend of a wetter environment during the EECO and drier environment after the EECO. Therefore, PC1 represents paleo-hydrological changes recorded by the Green River Formation in the mid-latitude region of North America (Fig. 10). Available terrestrial records in both North America and China indicate a wetter climate during the EECO, which is consistent with some model simulations (Hutchinson et al. 2018; Shields et al. 2021), although other models have predicted drier conditions (Winguth et al. 2010; Carmichael et al. 2017). Widespread lacustrine deposition and formation of oil shale deposits during the “greenhouse” interval is also consistent with evidence for a drastic reduction in Hadley cell circulation during the mid-Cretaceous (Hasegawa et al. 2012).

The organic-rich Mahogany Zone is located within the interval of rapidly fluctuating lake levels at ca. 48.5 Ma, which corresponds to the transition from the early Eocene “hothouse” to middle Eocene “greenhouse” and spans the termination of the EECO. In addition, rhythmically alternating beds of chert and dolomite occur where the δ¹⁸O values of benthic foraminifera transition from lighter to heavier values at ca. 48 Ma and heavier to lighter values at 45‒44 Ma. These alternating beds of chert‒dolomite are interpreted to form by the cyclical blooming and decomposition of the green algae Botryococcus (Kuma et al. 2019). Environmental stress (e.g., desiccation and/or an increase in temperature) during this transitional climate state led to massive Botryococcus blooms, enhancing hydrocarbon production (e.g., Demura et al. 2014) in this mid-latitude region of the US continental interior.

Our reconstruction of the paleoenvironmental conditions and paleohydrological changes recorded in Green River Formation sediments indicate that the interval of early Eocene warmth, including the EECO, was associated with a wetter environment and increased chemical weathering in the lower part of the section, whereas the cooler interval following the EECO was dominated by a drier environment and decreased chemical weathering. The deep lake stage and the highly fluctuating lake stages correspond to periods when oceanic sediments record heavier carbon isotopic compositions in benthic foraminifera and a shallowing of the CCD, providing useful insights into how the paleoenvironment and paleohydrology of the mid-latitude continental regions of North America responded to global climatic changes. Our data indicate that warming during the EECO likely resulted in widespread wetter conditions in mid-latitude continental areas of the Northern Hemisphere, which is consistent with some climate model simulations. Our results also show that the exceptionally organic-rich Mahogany Zone corresponds to the descent from the “hothouse” EECO into the cooler middle Eocene (ca. 48.7 Ma). This implies that climatic instability and resultant environmental stress during this transitional climate state led to massive blooms of Botryococcus algae in Lake Uinta. The remains of these algae became the source of the vast quantity of organic matter preserved in the Mahogany Zone.

Availability of data and material

All the data reported in this article are available from the corresponding author.

Competing interests

The authors declare that they have no competing interests.

Funding

This study was supported financially by JSPS Grant-in-Aid for Scientific Research (B) (No. 19H04256) and Grant-in-Aid for Young Scientists (B) (No. 16K21059) provided to H.H.

Authors’ contributions

R.K. and H.H. designed the research. R.K., H.H., and J.H.W. wrote the manuscript. R.K. and H.H. surveyed sections and collected samples. R.K. conducted all geochemical analysis and provided all the photographs. All authors edited and revised the paper.

Acknowledgements

We are very grateful for helpful discussions with many people, in particular Y. Asahara, H. Yoshida, M. Minami, M. Humblet of Nagoya University, R. Tada of University of Tokyo, T. Irino of Hokkaido University, and S. Nishimoto of Aichi University. We also thank to N. Katsuta of Gifu University, M. Ikeda of Tokyo University, H. Shozaki of Tokyo Institute of Technology, M. Sasaoka, K. Ishikawa of Kochi University, and Y. Muramiya of Fukada Geological Institute for collecting samples, T. Ohta of Waseda University and M. Civel-Mazens for supporting principal component analysis, T. Matsuzaki and Y. Shimizu of Kochi University for supporting XRD analysis, Y. Mizutani, T. Kimura and Z. Zhang of Nagoya University for supporting XRF analysis. This study was supported by cooperative research program at Center for Advanced Marine Core Research (CMCR), Kochi University (No. 20A027, 20B024, 21A037, 21B035)

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Paleohydrological change during the Early-Middle Eocene: Insights from Long-lived Green River Formation

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