Middle to late Cenozoic cooling and high topography in the central Rocky Mountains: Constraints from clumped isotope geochemistry

Earth and Planetary Science Letters 408 (2014) 35–47

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Earth and Planetary Science Letters www.elsevier.com/locate/epsl

Middle to late Cenozoic cooling and high topography in the central Rocky Mountains: Constraints from clumped isotope geochemistry Majie Fan a,∗ , Brian G. Hough b , Benjamin H. Passey c a b c

University of Texas at Arlington, Department of Earth and Environmental Sciences, Arlington, TX 76019, USA Hamilton College, Department of Geosciences, Clinton, NY 13323, USA Johns Hopkins University, Department of Earth and Planetary Sciences, Baltimore, MD 21218, USA

a r t i c l e

i n f o

Article history: Received 24 May 2014 Received in revised form 26 September 2014 Accepted 30 September 2014 Available online xxxx Editor: T.M. Harrison Keywords: Rocky Mountains paleorelief clumped isotopes oxygen isotopes temperature

a b s t r a c t The timing and mechanism of the formation of the high elevations of the central Rocky Mountains (Rockies) in the western interior of North America remain controversial. We examine the middle and late Cenozoic paleorelief between the Great Plains and the central Rockies using clumped isotope geothermometry of carbonate cements of fluvial and eolian sandstones. Our petrographic observations suggest that the studied carbonate cements were formed primarily during eodiagenesis in near-surface conditions and retain primary paleoclimate and paleoelevation signals. Carbonate clumped isotope temperatures decrease ∼10 ◦ C, and calculated water δ 18 O values decrease ∼4h from ∼35 Ma to ∼5 Ma in both the central Rockies and the western Great Plains, following middle and late Cenozoic global cooling trend. A persistent longitudinal gradient in temperature and δ 18 O value existed between the central Rockies and the Great Plains during the Oligocene and Miocene, with temperatures 4–8 ◦ C higher, and calculated water δ 18 O values 3–6h higher in the Great Plains. These observations suggest the mean elevation of the central Rockies was ∼1 km higher than the western Great Plains during the middle and late Cenozoic, similar to the present-day regional relief. When placed in the context of other paleoaltimetry studies and geological observations, our findings support the hypothesis that the high mean topography of the Rockies was developed during the late Eocene, possibly related to isostatic adjustment, or dynamic uplift caused by foundering of lower mantle lithosphere or the Farallon slab, or both. © 2014 Elsevier B.V. All rights reserved.

1. Introduction The central Rocky Mountains (Rockies) in Wyoming and its adjacent areas are an extensive region of high mountains and intermontane basins bounding the east side of the North American Cordillera (Fig. 1A). The region was near sea level in the Western Interior Seaway at ∼80 Ma based on thick Cretaceous marine sedimentation (Roberts and Kirschbaum, 1995). At present, the mean elevation of the central Rockies is ∼2.2 km with mountain peaks up to 4 km high and basin floors at 2–1.5 km above sea level. The high topography and high relief of the central Rockies gradually decreases eastward to the low, flat Great Plains (∼0.9 km high) in western Nebraska (Fig. 1B). It is generally accepted that much of the current elevation of the central Rockies was created during the late Cretaceous–early Eocene Laramide orogeny, and that the orogeny disrupted marine sedimentation and partitioned the

*

Corresponding author. E-mail address: [email protected] (M. Fan).

http://dx.doi.org/10.1016/j.epsl.2014.09.050 0012-821X/© 2014 Elsevier B.V. All rights reserved.

region into intervening basement-cored uplifts and sedimentary basins as a result of shallow subduction of the Farallon oceanic plate underneath the western U.S.A. (e.g., Dickinson and Snyder, 1978; DeCelles, 2004; Liu et al., 2010a). However, recent paleoaltimetry studies suggest the Laramide orogeny only formed the high mountain ranges with mean elevations similar to or higher than today (Gregory and Chase, 1992, 1994; Wolfe et al., 1998; Fan and Dettman, 2009; Fan et al., 2011), whereas basin floors remained low (∼500 m) during the earliest Eocene (MacGinitie, 1969; Fan et al., 2011). Therefore, the central Rockies must have experienced post-early Eocene surface uplift of the basin floors. Opposing datasets have led to different interpretations of the timing and mechanism of post-early Eocene surface elevation changes in the central Rockies. Mapping the distribution of basin fill remnants within the Rocky Mountain basins and interpretation of tilting history of the adjacent western Great Plains have led to the view that slow, regional subsidence continued to the late Miocene (Bradley, 1987; McMillan et al., 2002, 2006; Leonard, 2002), and differential uplift of the central Rockies occurred during the latest Miocene–Pliocene with rivers eroding out much

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M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

Fig. 1. (A) Map of the study area in the western U.S.A. showing the distribution of Oligocene–Miocene sedimentary rocks (gray area), and the measured T (47 ) values of carbonate cement. (B) Mean (black line) and variation (gray zone) of elevation along a transect (dashed line in A) across the central Rockies and adjacent Great Plains showing an eastward decrease in surface elevation. GM: Granite Mountains, BH: Bighorn Mountains, BHs: Black Hills; WR: Wind River Range, OM: Owl Creek Mountains. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

of the basin fill (McMillan et al., 2006; Duller et al., 2012). It is suggested that this uplift was the crustal response to adding upper mantle buoyancy beneath the Rockies (Angevine and Flanagan, 1987; McMillan et al., 2002; Moucha et al., 2008; Karlstrom et al., 2011). Conversely, widespread late–middle Eocene erosion in the central Rockies and the disconformity between the upper Cretaceous and middle Cenozoic strata in western Nebraska have led to a different view that regional subsidence only continued to the early–late Eocene, and differential uplift of the central Rockies occurred at ∼42–37 Ma (Cather et al., 2012). This uplift may reflect the isostatic adjustment and/or thermal perturbation of the lithosphere of the overriding North American plate when the Farallon shallow slab was removed or when the mantle lithosphere was delaminated (Roy et al., 2009; Cather et al., 2012; Roberts et al., 2012). Evaluation of these two hypotheses of postearly Eocene uplift requires new data to elucidate the surface uplift history of the central Rockies during the middle and late Cenozoic. While stable isotope geochemistry remains an important tool for constraining surface elevation and climate change, the accuracy of such constraints is confounded by the observation that elevation and climate changes may cause similar stable isotope signatures of surface water (Jeffery et al., 2011; Poulsen and Jeffery, 2011; Feng et al., 2013). Stable isotope paleoaltimetry is further complicated by changes of vapor condensation temperature and atmospheric circulation, moisture recycling, intensity of evaporation and sublimation, and drainage reorganization (e.g., Davis et al., 2009; Lechler and Niemi, 2011; Chamberlain et al., 2012). Carbonate clumped isotope geothermometry provides a new opportunity to reconstruct paleoelevation with less error (Huntington et al., 2010; Quade et al., 2013; Hough et al., 2014; Snell et al., 2014). Carbonate clumped isotope geothermometry determines the carbonate mineralization temperature using the temperature-dependent enrichment of the “clumped” isotopologue 13 C18 O16 O(47 ) in CO2 derived from phosphoric acid digestion of carbonate (e.g., Ghosh et al., 2006; Eiler, 2007; Huntington et al., 2010; Dennis et al., 2011). Clumped isotope geothermometry studies of Quaternary lake and soil carbonates have shown that carbonate mineraliza-

tion temperatures primarily reflect lake and soil temperatures and the temperatures are sensitive to elevation changes in the western U.S.A., with a lapse rate of ∼−4 ◦ C/km (Huntington et al., 2010; Hough et al., 2014). This rate is generally consistent with the lapse rates of modern air and lake water temperature in the western U.S.A., which vary between 3 ◦ C/km and 5 ◦ C/km (Meyer, 1992; Wolfe, 1992; Huntington et al., 2010; Hough et al., 2014). Additionally, clumped isotope geothermometry improves carbonate oxygen isotope paleoaltimetry by providing the temperature information needed for unambiguous calculation of water δ 18 O values based on measured carbonate δ 18 O values (Passey et al., 2010; Peters et al., 2013; Hough et al., 2014). The accuracy of elevation estimates based on clumped isotope geothermometry is influenced by several factors, including carbonate diagenesis (Huntington et al., 2011), seasonality of carbonate formation (Peters et al., 2013; Hough et al., 2014), and lapse rate of temperature change to elevation (e.g., Meyer, 1992; Wolfe, 1992). An emerging approach in clumped isotope paleoaltimetry is to compare the clumped isotope temperature difference of the same type of carbonate material between the high-elevation site and a nearly synchronous low-elevation site, and applying the lapse rate to derive the paleorelief (e.g., Lechler et al., 2013; Snell et al., 2014; Garzione et al., 2014). This method assumes that climate change produces equal temperature changes at both high and low elevation sites when atmospheric pCO2 is stable (Poulsen and Jeffery, 2011). Once diagenesis of carbonate samples can be ruled out, the major uncertainty of clumped isotope paleoaltimetry is from the variation of temperature lapse rate, and uncertainty of the clumped isotope-derived temperature estimates. Here we report 47 -derived temperatures (T (47 )) from carbonate cements of fine-grained sandstones in Wyoming and western Nebraska, and undertake a rigorous examination using isotopic, petrographic and chemical techniques to assess the presence and extent of diagenetic influences on the mineralogy and 47 values of these samples. Our objectives are to test whether carbonate cement T (47 ) values can primarily record near-surface temperature changes, and, if so, to provide insight into the topographic evolution of the central Rockies with respect to the adjacent Great Plains during the middle and late Cenozoic. 2. Geologic and climate settings Our study area is located in the central Rocky Mountains in Wyoming and the adjacent Great Plains in western Nebraska (Fig. 1). Thick early Paleogene fluvial and lacustrine strata record synchronous exhumation of the Laramide mountain ranges and subsidence of the intermontane basins during the Laramide orogeny (e.g., Dickinson et al., 1988; DeCelles, 2004). The tectonic processes during the middle Eocene–Miocene are debated, as summarized in the introduction of this paper. Although late Cenozoic normal faults are relatively common in Wyoming, they record only a few percent extension (Snoke, 1993). The major normal faults bounding the southern ends of the Granite Mountains and Wind River Range in central Wyoming caused ∼0.6 km of collapse of the two mountains during the latest Miocene–Pliocene (Love, 1970; Steidtmann and Middleton, 1991; Snoke, 1993), probably due to the west-to-east migration of thermal cooling of the Yellowstone hotspot during the last 10 Ma (Anders et al., 2009). Modern climate of the study area is arid to semi-arid, with annual precipitation amounts increasing from ∼15 cm in central Wyoming to ∼60 cm in western Nebraska, and mean monthly air temperatures ranging from −10 ◦ C to 22 ◦ C (NCDC, 2012). The hydroclimate is dominated by the competing influences of the dry polar air mass, the Westerlies, and the moist tropical–subtropical air masses from the Gulf of Mexico, the Atlantic, and the Pacific Ocean (Bryson and Hare, 1974). The juxtaposition of the relatively

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

37

Table 1 Sample location information and carbonate petrography and XRD results. Sample ID

Lat. (◦ N)

Long. (◦ W)

Elev. (m)

Estimated age (Ma)

OB1-25

42.25

108.85

2395

49 ± 1

OB2-11 WS230 SP2-0 SP3-99 SP3-104 CB146 CB49 MPm40 MPm10+3 LTG740 LTG156 EF10-04 EF15-90 EF2-04 snail LaG30-3 LaG30+7 LaG6-34 MC-H-10m-top MC-H-top H35-4540 RF4660 ACH 004 AH2m+base AH1m+C001

42.32 42.61 42.69 42.44 42.44 42.37 42.37 42.55 42.55 42.65 42.65 42.63 42.61 42.71 42.86 42.86 42.82 41.60 41.60 42.42 42.56 41.47 41.60 41.47

108.86 108.27 107.26 107.56 107.56 107.06 107.06 107.49 107.49 106.74 106.74 105.30 105.25 105.36 103.58 103.58 103.59 102.78 102.78 103.73 103.98 103.07 102.78 103.07

2301 2068 2022 1935 1940 1909 1905 2059 2050 1826 1750 1484 1578 1489 1148 1152 1291 1180 1190 1385 1428 1252 1199 1275

24 ± 1 28.2 ± 1 25 ± 2 16 ± 0.3 16 ± 0.3 10.5 ± 0.8 10.5 ± 0.8 4±1 5±1 32.5 ± 1 36 ± 1 33 ± 0.5 32.5 ± 0.5 33 ± 1.5 34 ± 0.5 34 ± 0.5 32 ± 1 26 ± 3 26 ± 3 19 ± 0.5 18 ± 1 8±1 9±2 6±1

Carbonate type

CL categorya

lacustrine oolitic cement cement cement cement cement pedogenic cement cement cement cement cement cement cement diagenetic cement cement cement cement cement cement cement pedogenic cement cement

3 2 – 3 1 – 2 2 2 3 3 – 3 3 Spar 3 3 3 – – 3 3 2 – 1

Spar (%)b

TIC (%)

Carbonate (%)c

37.6

11.5

96

4 .4

5 .4 8 .7 3 .6 3 .2 4 .0 6 .8 4 .6 5 .2 5 .9 3 .9 5 .7 5 .1 3 .2 12.0 3 .3 4 .5 5 .7 4 .7 7 .6 3 .4 5 .0 9 .0 2 .5 6 .7

45 73 30 27 33 57 38 43 49 33 48 43 27 100 28 38 48 39 63 28 42 75 43 27

– 4 .3 82.0 – 0 .0 1 .2 6 .7 20.0 0 .6 – 2 .6 1 .9 100.0 7 .6 6 .1 16.6 – – 23.0 13.1 0 .0 – 42.0

Carbonate typed

Formation/ group

–

Green River Fm.

LMC LMC LMC LMC LMC LMC LMC, D LMC LMC LMC LMC LMC LMC – LMC LMC, D LMC – LMC LMC LMC, HMC LMC – LMC

Arikaree Fm. White River Grp. Split Rock Fm. Split Rock Fm. Split Rock Fm. Split Rock Fm. Split Rock Fm. Moonstone Fm. Moonstone Fm. White River Fm. White River Fm. White River Fm. White River Fm. White River Fm. White River Grp. White River Grp. White River Grp. Arikaree Grp. Arikaree Grp. Arikaree Grp. Arikaree Grp. Ogallala Fm. Ogallala Fm. Ogallala Fm.

a 1. Samples showing micrite-spar microtexture. 2. Samples showing dominantly dull or no luminescence with minor amount of bright orange luminescence. 3. Samples showing two different intensities of luminescence that are spatially intermixed, or bright orange luminescence occurring as isolated, vugular pore fillings in the form of blocky spars; dash (–): no analysis. b c d

Spar percentage is calculated using formula: S a /( T a ∗ C%), where S a and T a are the surface area of spar and whole thin section, and C% is carbonate percentage. Carbonate percentage is calculated from TIC% by assuming all carbonate is calcite. LMC: low-magnesium calcite; HMC: high-magnesium calcite; D: dolomite.

18

O-enriched air masses from the Gulf of Mexico and the relatively O-depleted air masses from the Pacific and their temporal contribution cause an eastward increase of modern precipitation isotope values associated with the decrease of elevation in our study area (Liu et al., 2010b). 18

3. Stratigraphy The latest Eocene–Miocene strata in the central Rockies and western Great Plains are primarily composed of eolian and fluvial fine-grained sandstone and siltstone, with thicknesses ranging between <0.1 km and ∼1.5 km (Fig. 1A). In the central Rockies, we sampled the middle Eocene Green River Formation in southwest Wyoming (Zeller and Stephens, 1969), latest Eocene–Oligocene White River Formation distributed in central and eastern Wyoming (Emry, 1973, 1975; Evanoff, 1990), the early Miocene Arikaree Formation in western Wyoming (Zeller and Stephens, 1969), the middle–late Miocene Split Rock Formation (Love, 1961, 1970) and the Pliocene Moonstone Formation (Love, 1970; Anders et al., 2009) in central Wyoming. In the western Great Plains, we sampled the latest Eocene–Oligocene White River Group (LaGarry, 1998), the latest Oligocene–early Miocene Arikaree Group (MacFadden and Hunt, 1998), and the late Miocene Ash Hollow Formation (Diffendal, 1987) in western Nebraska. The estimated ages of samples are based on published magnetostratigraphy, radiometric age constraints, and North America Land Mammal Ages (Table 1). Burial depths of the sampled strata in the central Rockies vary between 1.0 km and 2.8 km based on hydrocarbon thermal maturation modeling conducted in the Green River and Powder River basins (Roberts et al., 2005; Anna, 2005), preserved stratal thicknesses in other locations (e.g., Love, 1961, 1970), and the postlatest Miocene regional erosion amount of ∼1.5 km (Pelletier, 2009). Burial depths of the sampled strata in the western Great

Plains vary between 0.6 km and 1.7 km based on the preserved overlying strata thickness in western Nebraska (Swinehart et al., 1985), and a post-latest Miocene regional erosion of ∼0.5 km (Cather et al., 2012). By assuming a normal geothermal gradient of 25 ◦ C/km, burial temperatures of the strata vary between 25 and 70 ◦ C, significantly lower than the critical temperature (∼100 ◦ C) above which solid-state C–O bond reordering is thought to alter carbonate clumped isotope compositions (e.g., Dennis and Schrag, 2010; Henkes et al., 2014). We collected 24 fine-grained sandstones and one altered land snail fossil from outcrops along a regional transect from central Wyoming to western Nebraska (Fig. 1, Table 1). Fifteen of the samples were collected from the central Rockies to ensure a long and continuous temperature record, and the nine other samples were collected from western Nebraska to establish the reference temperature at a low-elevation site. 4. Analytical methods Fine-grained sandstone samples were broken into small pieces with the longest axis less than 1 cm and evaluated under magnification to avoid sampling visible spar. The samples were then ground using a ceramic mortar and pestle. Because this approach does not exclude small spar crystals in the powdered samples, the amount of spar in each sample is estimated based on petrographic analysis. The altered land snail was the only sample with sufficient sparry calcite separable for 47 analysis. This sample was carefully drilled using a hand-held drill under magnification to avoid possible contamination from the host rock carbonate cement. 4.1. Mineralogical, TIC%, and petrographic analysis X-ray diffraction (XRD) was performed on 21 samples. Diffraction patterns were obtained using a Bruker D8 Advance Diffractometer equipped with Cu(Kα ) radiation at 40 keV and 40 mA.

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M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

Samples were scanned from 15◦ to 55◦ 2θ at 0.02◦ steps. Carbonate mineralogy, including high-magnesium calcite (HMC), lowmagnesium calcite (LMC), and dolomite were determined based on the position of the d(104) peak. If more than one type of carbonate occurred, the ratios of different types of carbonate were estimated based on the peak heights of each carbonate, which is a semiquantitative approach. Percentage of total inorganic carbon (TIC) of all samples was determined using a UIC CO2 Coulometer, which has an analytical precision better than 0.8% based on repeated analysis of a pure calcite lab standard with weight of 5–35 mg. Polished thin sections were produced from 16 sandstone samples and two pedogenic carbonate samples. Thin sections were examined utilizing both polarizing and cathodoluminescence microscopes to characterize their primary and diagenetic fabrics. A Reliotron cold cathodoluminoscope was operated at an acceleration voltage of 7–9 kV and a beam current intensity of 0.3–0.5 mA to examine the luminescence of carbonate cement. Surface areas of spars were measured using Image J software. 4.2. Isotope analysis Carbonate clumped-isotope analysis was conducted at Johns Hopkins University following the methods of Passey et al. (2010). Two or three aliquots of each sample were reacted at 90 ◦ C with 100% phosphoric acid, with the resultant CO2 being cryogenically purified before analysis on a Thermo MAT 253 mass spectrometer. Equilibrium CO2 samples prepared at 1000 ◦ C and 30 ◦ C were analyzed concurrently with the sample unknowns to normalize sample 47 values (Dennis et al., 2011). The 47 values were also normalized to 25 ◦ C acid reaction temperatures using a correction factor of 0.092h (Henkes et al., 2013). Clumped-isotope data of calcite cements are reported relative to the absolute reference frame (ARF) of Dennis et al. (2011), and temperatures are calculated from 47 values using the inorganic calibration of Ghosh et al. (2006), as adapted to the ARF scale by Dennis et al. (2011). Because there is currently no published dolomite-specific experimental clumped isotope calibration, and no clear experimental evidence that the clumped isotope calibrations of dolomite and calcite are different (Bonifacie et al., 2011; Ferry et al., 2011), we use the calcite calibration to calculate the clumped isotope temperatures of samples containing dolomite and HMC. During the course of the study, two internal carbonate standards (Carrara Marble and 102-GC-AZ01) and one international carbonate standard (NBS-19) were analyzed multiple times. The mean 47 values and standard deviation (SD) of these standards are 0.399 ± 0.013h (n = 10), 0.713 ± 0.016h (n = 14), and 0.391 ± 0.010h (n = 3), respectively. The standard √ errors (SE) for 47 reported in Table 2 are calculated as SE = SD/ N, where SD is the observed standard deviation for repeated analyses of samples, and N is the number of repeated analyses. If the observed SD is lower than the known laboratory precision (= 0.013h, the mean precision we observed for repeated analyses of standards), SE is calculated using SD = 0.013h. Error in T (47 ) is calculated by propagating the SE for 47 and an estimated error in 47 of 0.0028h for the acid temperature correction (Passey et al., 2010), through the relevant paleotemperature equations. The δ 13 C and δ 18 O values of the carbonate samples were determined during clumped isotope analysis. Values were standardized to VPDB based on concurrent analyses of NBS-19, or internal standards calibrated to NBS-19, and the analytical precision is <0.1h for both values. δ 18 O values of waters in equilibrium with precipitating calcite cements were calculated from T (47 ) and carbonate δ 18 O values using the calibration of Kim and O’Neil (1997). For the samples with mixed calcite and dolomite, two end-member water δ 18 O values were calculated by assuming the sample is pure dolomite using the dolomite calibration of Vasconcelos et

al. (2005) and by assuming the sample is pure calcite using calcite calibration. The final water δ 18 O value is calculated from the two end-member water δ 18 O values and the relative abundance of calcite and dolomite based on XRD. The uncertainties of the calculated water δ 18 O values were derived from the uncertainties of T (47 ) because the uncertainties of carbonate δ 18 O values were small. 5. Results 5.1. Mineralogical, TIC%, and petrography analysis Results of XRD, carbonate type, and TIC%, are presented in Table 1. Supporting XRD patterns are presented in Fig. A.1. XRD results show that 18 of the 21 studied samples contain only LMC, one sample (RF4660) in western Nebraska contains HMC with an HMC/LMC ratio of 0.25, one sample (CB49) in the central Rockies contains dolomite with a dolomite/LMC ratio of 0.23, and one sample (Lag30-7) in western Nebraska contains dolomite with a dolomite/LMC ratio of 0.45. The weight percentage of TIC varies from 2.5% to 12.0%. TIC% is used to estimate carbonate weight percentage. The estimated carbonate weight percentage of all the samples, assuming the carbonate cements are LMC, varies from 27% to 100%, with the highest occurring in the oolitic limestone (OB1-25) and altered snail sample (EF2-04 snail). This assumption overestimates the carbonate percentages of the samples containing HMC and dolomite by at most 2%. The carbonate percentage and spar percentage data are not correlated to the burial depth of the samples (Fig. 2). Under a polarizing microscope, the carbonate cements are observed to be mostly micritic with blocky, sparry carbonate occurring in isolated vugs less than 1 mm in diameter in most of the studied samples (Fig. 3). In only two samples (SP3-99 and AH 1m+C001), micrite-spar microtexture occurs, and the two samples are classified as Category 1 cement. Siliciclastic grains are dominantly quartz, fresh feldspar, and volcanic glass, and detrital carbonate grains are not observed. Minor grains include undeformed mica and sedimentary and volcanic lithic fragments. Under a cathodoluminescence microscope, the sparry carbonate cement shows bright orange luminescence color, and the micritic carbonate cement shows two different luminescence colors of variable abundance. Based on the luminescence color difference, we classified the rest of the samples into two categories (Fig. 3). Category 2 samples exhibit dull or no luminescence, with minor flecks of bright orange luminescence. Category 3 samples exhibit at least two different intensities of luminescence colors of almost equal proportions that are spatially intermixed, or bright orange luminescence occurring as isolated, vugular pore fillings in the form of blocky spars in the matrix of intermixed luminescences. No carbonate of clear rhombic crystals is observed in our samples. The amount of sparry carbonate is highly variable (Table 1). Although the vug-fills of bright orange luminescence and cement of different intensity of luminescence are the clear result of diagenesis, it was not possible to develop a consistent cement stratigraphy, and there was insufficient material for 47 analysis. 5.2. Isotope analysis All isotope data are presented in Table 2. Supporting data for clumped isotope analyses are presented in Table A.1. 47 values ranged from 0.64h to 0.76h (relative to the ARF), with an average difference between replicate measurements of 0.015h. Temperatures calculated from the 47 values (T (47 )) range from 16 ◦ C to 37 ◦ C (Table 2). Four observations were made from the T (47 ) results: 1) the T (47 ) value of the altered land snail sample (EF2-04) is 4–5 ◦ C higher than the samples about 100 m

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

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Table 2 Clumped isotope analysis results. Sample ID Central Rockies OB1-25 OB1-25 OB1-25 OB1-25 avg OB2-11 OB2-11 OB2-11 avg WS230 WS230 WS230 avg SP2-0 SP2-0 SP2-0 SP2-0 avg SP3-99 SP3-99 SP3-99 avg SP3-104 SP3-104 SP3-104 avg CB146 CB146 CB146 CB146 avg CB49 CB49 CB49 avg MPm40 MPm40 MPm40 avg MPm10+3 MPm10+3 MPm10+3 avg LTG740 LTG740 LTG740 LTG740 avg LTG156 LTG156 LTG156 avg EF10-04 EF10-04 EF10-04 avg EF15-90 EF15-90 EF15-90 EF15-90 avg EF2-04 snail EF2-04 snail EF2-04 snail avg Western Great Plains LaG30-3 LaG30-3 LaG30-3 LaG30-3 avg LaG30+7 LaG30+7 LaG30+7 avgf LaG6-34 LaG6-34 LaG6-34 LaG6-34 avg MC-H-10m-top MC-H-10m-top MC-H-10m-top avg MC-H-top MC-H-top MC-H-top avg H35-4540 H35-4540 H35-4540 avg RF4660 RF4660

δ 13 Ca (h, VPDB)

δ 18 O (h, VPDB)

47 a (h, ARF)

− 0.1 0 .1 0 .0 0 .0 −2.3 −2.2 −2.3 −3.6 −3.6 −3.6 −4.4 −4.4 −4.5 −4.5 −6.0 −6.0 −6.0 −6.0 −5.9 −6.0 −6.2 −6.3 −6.3 −6.3 −6.4 −6.5 −6.5 −6.1 −6.1 −6.1 −6.8 −6.8 −6.8 −6.0 −6.0 −6.1 −6.1 −5.8 −5.8 −5.8 −6.4 −6.4 −6.4 −6.7 −6.6 −6.6 −6.7 −7.2 −7.0 −7.1

−6.6 −6.7 −6.4 −6.6 −14.8 −14.8 −14.8 −15.1 −15.1 −15.1 −15.4 −15.5 −15.5 −15.4 −17.8 −17.8 −17.8 −17.7 −17.7 −17.7 −14.9 −14.9 −15.0 −15.0 −15.1 −15.0 −15.1 −13.6 −13.7 −13.6 −15.2 −15.5 −15.4 −12.6 −12.7 −13.3 −12.9 −12.7 −12.8 −12.8 −12.0 −11.9 −12.0 −10.8 −10.8 −10.6 −10.7 −12.9 −12.5 −12.7

0.685 0.678 0.671 0.678 0.687 0.701 0.694 0.701 0.723 0.712 0.683 0.701 0.699 0.694 0.713 0.717 0.715 0.732 0.724 0.728 0.750 0.728 0.714 0.731 0.734 0.710 0.722 0.759 0.749 0.754 0.729 0.740 0.735 0.717 0.725 0.741 0.727 0.696 0.721 0.708 0.709 0.695 0.704 0.708 0.690 0.716 0.705 0.659 0.682 0.671

−6.6 −6.6 −6.4 −6.5 −6.4 −6.4 −6.4 −7.4 −7.4 −7.4 −7.4 −6.9 −6.9 −6.9 −6.8 −6.8 −6.8 −7.5 −7.4 −7.5 −7.7 −7.6

−13.4 −13.4 −13.3 −13.4 −12.7 −12.8 −12.8 −9.6 −9.7 −9.8 −9.7 −10.3 −10.2 −10.3 −8.9 −8.8 −8.9 −12.0 −12.0 −12.0 −12.6 −12.5

0.644 0.654 0.669 0.656 0.684 0.672 0.678 0.702 0.699 0.670 0.690 0.709 0.710 0.710 0.722 0.723 0.723 0.676 0.709 0.693 0.696 0.699

SEb (h, ARF)

T (47 )c (◦ C)

δ 18 Owarer d (h, VSMOW)

Estimated age (Ma)

Burial depth (km)e

References

Roberts et al., 2005

0.008

32.0 ± 1.8

−2.8 ± 0.3

49 ± 1

2.8

0.009

28.5 ± 1.9

−11.7 ± 0.4

24 ± 1

1.6

0.009

24.6 ± 1.9

−12.9 ± 0.4

28.2 ± 1

1.8 Love, 1961, 1970

0.008

28.5 ± 1.7

−12.4 ± 0.3

25 ± 2

1.6

0.009

24.1 ± 1.8

−15.7 ± 0.4

16 ± 0.3

1.6

0.009

21.5 ± 1.8

−16.0 ± 0.4

16 ± 0.3

1.6

0.010

20.9 ± 2.0

−13.5 ± 0.4

10.5 ± 0.8

1.2

0.012

22.7 ± 2.4

−13.9 ± 0.5

10.5 ± 0.8

1.2

0.009

16.4 ± 1.7

−13.1 ± 0.4

4±1

1.0

0.009

20.1 ± 1.8

−14.0 ± 0.4

5±1

1.0 Anna, 2005

0.008

21.6 ± 1.6

−11.2 ± 0.3

32.5 ± 1

1.5

0.013

25.5 ± 2.7

−10.3 ± 0.5

36 ± 1

1.5

0.009

26.4 ± 1.9

−9.4 ± 0.4

33 ± 0.5

1.5

0.008

26.3 ± 1.7

−8.1 ± 0.3

32.5 ± 0.5

1.5

0.011

33.8 ± 2.5

−8.6 ± 0.5

33 ± 1.5

1.5 Swinehart et al., 1985

0.008

37.2 ± 1.9

−8.6 ± 0.3

34 ± 0.5

1.7

0.009

32.1 ± 2.0

−10.4 ± 0.4

34 ± 0.5

1.7

0.010

29.4 ± 2.2

−6.5 ± 0.4

32 ± 1

1.5

0.009

25.2 ± 1.9

−7.9 ± 0.4

26 ± 3

1.3

0.009

22.6 ± 1.8

−7.0 ± 0.4

26 ± 3

1.3

0.017

28.9 ± 3.6

−8.9 ± 0.7

19 ± 1

1.0

(continued on next page)

40

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

Table 2 (continued) Sample ID

δ 13 Ca (h, VPDB)

δ 18 O (h, VPDB)

47 a (h, ARF)

SEb (h, ARF)

T (47 )c (◦ C)

δ 18 Owarer d (h, VSMOW)

RF4660 avg ACH 004 ACH 004 ACH 004 ACH 004 avg AH1m+C001 AH1m+C001 AH1m+C001 avg AH2m+base AH2m+base AH2m+base avg

−7.7 −6.7 −6.7 −6.7 −6.7 −6.5 −6.6 −6.5 −6.9 −7.0 −7.0

−12.5 −12.9 −12.9 −12.8 −12.9 −12.7 −12.8 −12.7 −12.1 −12.3 −12.2

0.697 0.681 0.715 0.718 0.705 0.713 0.705 0.709 0.749 0.717 0.733

0.009

27.8 ± 1.9

−9.6 ± 0.4

18 ± 0.5

1.0

0.012

26.3 ± 2.5

−10.2 ± 0.5

8±1

0.7

0.009

25.4 ± 1.9

−10.3 ± 0.4

8±1

0.7

0.016

20.4 ± 3.1

−10.8 ± 0.7

5±2

0.6

Estimated age (Ma)

Burial depth (km)e

References

a

ARF: Absolute Reference Frame. SE is the standard error of 47 , SE = SD(47 )/SQRT( N ). When SD of a sample is less than the observed long-term SD of lab standards (0.013h), the long-term value of 0.013h is assigned as the SD of the sample. c T is calculated using Ghosh et al. (2006) calibration, adapted to the ARF scale as reported by Dennis et al. (2011): T = SQRT(63 600/(47 + 0.0047)) − 273.15. d Kim and O’Neil (1997): 103 ln α = 18030/( T + 273.15) − 32.42. e Burial depth is the sum of the thickness of preserved overlying strata and the amount of post-Miocene erosion, see text for the details. f Temperatures for samples with dolomite and low-Mg calcite are based on the calcite calibration of Ghosh et al. (2006), as there is not yet a published, dolomite-specific experimental calibration. Water δ 18 O values are calculated from T (47 ) and δ 18 O (mineral) values using a linear combination of the dolomite calibration of Vasconcelos et al. (2005) and calcite calibration of Kim and O’Neil (1997), and relative abundance of the minerals based on XRD results. b

Fig. 2. Plots showing the relationship between carbonate percentage and burial depth of the samples (A), and the relationship between spar percentage and burial depth of the samples (B). Numbers in B represent carbonate category in Table 1.

stratigraphically below and above (Table 2); 2) T (47 ) values decrease ∼10 ◦ C from ∼35 Ma to ∼5 Ma in both the central Rockies and western Great Plains, similar to the trend of global deepsea carbonate δ 18 O values (Zachos et al., 2001a, 2001b) (Fig. 4); 3) T (47 ) values in the central Rockies do not show a distinct decrease across the Eocene–Oligocene transition (Fig. 4A), that corresponds with the decrease of global deep-sea carbonate δ 18 O values (Zachos et al., 2001a, 2001b). 4) T (47 ) values increase 4–8 ◦ C from the central Rockies to the western Great Plains with gradients of 1.0–2.8 ◦ C per degree of longitude during the Oligocene and Miocene (Fig. 5A). Carbonate δ 13 C and δ 18 O values vary between −7.7h and 0.0h, and between −17.8h and −6.6h, respectively (Fig. 6). The calculated water δ 18 O values generally vary from −16.0h to −6.5h, except for sample OB1-25 (lacustrine oolitic limestone), which has a value of −2.8h. The calculated water δ 18 O values decrease ∼4h from ∼35 Ma to ∼5 Ma in both the central Rockies and the Great Plains (Fig. 4). The calculated water δ 18 O values increase 3–6h from the central Rockies to the western Great Plains with gradients of 0.8–1.8h per degree of longitude during the Oligocene and Miocene (Fig. 5B). The calculated water δ 18 O values are weakly correlated with the T (47 ) values (Fig. 6B).

6. Discussion 6.1. Screening for diagenesis Before the isotope ratios can be used to reconstruct temperature, precipitation δ 18 O values, and paleoelevation, diagenesis of carbonate cement must be evaluated to assess its possible influence on clumped and stable isotope geochemistry. Carbonate cement of siliciclastic rocks can be formed at any time after the deposition of the siliciclastic grains, and are classified into three categories: early diagenetic (eodiagenetic), deep burial diagenetic (mesodiagenetic), and late diagenetic (telodiagenetic) carbonate (Choquette and Pray, 1970). Eodiagenesis takes place at near-surface temperature and pressure, and the stable isotope compositions of eodiagenetic carbonate cement record surface water isotope compositions, and thus may primarily record climate and environmental information. On the other hand, mesodiagenesis takes place at high temperature and pressure in the presence of fluid that may have isotopic compositions different from those of surface water because of potential water–rock interactions deep in sedimentary basins (Choquette and Pray, 1970; Brand and Veizer, 1981), and telodiagenesis occurs near the surface during post-burial uplift and associated erosion, in the presence of

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

41

Fig. 3. Photomicrographs of studied carbonate samples. Scale bars are 200 μm. (A) Category 1 cement under plane-polarized light showing micrite-spar microtexture; (B) Category 2 cement under cathodoluminescent light showing dull luminescence with minor flecks of bright orange luminescence. (C–F) Category 3 cement under cathodoluminescent light showing two different intensities of luminescence that are spatially intermixed, or bright orange luminescence occurring as isolated, vugular pore fillings in the form of blocky spars. O: ooids that are recrystallized to sparry calcite or experienced dissolution, R: root traces. Blue grains are quartz and green grains are feldspar. (G–H) Category 3 cement under cross-polarized light showing isolated, irregular vugs that are filled by blocky spars (outlined in dashed lines). (I) Category 1 cement under cross-polarized light showing that micritic carbonate fills pore spaces, and the presence of undeformed minerals. B: biotite, F: feldspar, L: lithics, Q: quartz, VG: volcanic glass. (For interpretation of the references to color in this figure, the reader is referred to the web version of this article.)

fluid that may have isotopic compositions different from those of the original surface water because of changes of near-surface climate conditions through time (e.g., Morad, 1998). Determination of the diagenetic history of carbonate cement often relies on thin section microscopy and mineralogy studies, and carbonate oxygen and carbon isotope compositions (e.g., Beckner and Mozley, 1998; Garzione et al., 2004). Multiple lines of evidence suggest that the dominant carbonate cement of our studied samples, in the form of micritic calcite, formed primarily during eodiagenesis. First, undeformed ductile grains such as mica, and unaltered feldspar and volcanic glass, are present in the samples, suggesting the samples were not compacted by deep burial (Fig. 3I). Mesogenesis occurring during burial tends to alter mica, volcanic glass, and feldspar to clay minerals (e.g., Morad, 1998). Second, the mean TIC% of the 34–24 Ma old sandstone samples is 5.1%, which is equivalent to 43% calcite. The high amount of cement indicates the intergranular volumes of the samples are similar to the average original porosity of sandstone, which is ∼40% (Sclater and Christie, 1980), suggesting the cements were formed before compaction. Finally, the clumped isotope temperatures of all the samples are not high compared to modern soil temperatures in the region (Hough et al., 2014), suggesting the carbonate was mineralized at near surface temperatures. Although we cannot completely rule out the influence of mesodiagenesis and telodiagenesis on our samples based on luminescence alone, several lines of evidence suggest that the amount of carbonate formed by telodiagenesis and mesodiagenesis is low,

and exerts minimal influence on our isotope data. First, the major period of exhumation in the study area occurred during the last 8–6 myr, which caused ∼1.5 km of erosion in the Rockies (McMillan et al., 2006; Pelletier, 2009). If telodiagenesis induced by this erosion event occurred, we would expect to see that the youngest samples have experienced the highest degree of telodiagenesis because these samples are closest to surface. However, the carbonate category and the percentage of spar of our seven samples younger than 11 Ma are not systematically different from the samples older than 11 Ma (Table 1). This observation suggests exhumation of the rocks in the last 8–6 myr did not result in growth of the sparry calcite or formation of carbonate with differing luminescence. Second, our oldest sample (OB1-25) was formed as oolitic limestone in Lake Gosiute during the middle Eocene and has experienced significant diagenesis as evidenced by large amounts of sparry calcite (∼35%) as well as intermixed luminescence (Fig. 3E). Telodiagenesis should decrease the δ 18 O value of sample OB1-25 because surface water has lower δ 18 O value compared to evaporated lake water. However, the δ 18 O and δ 13 C values of this sample are consistent with the isotope values of other lacustrine micritic calcite formed in evaporative environments in the Green River Formation (Carroll et al., 2008). The calculated δ 18 Owater value is −2.8h, suggesting that the water of the Lake Gosiute was evaporatively enriched in 18 O, consistent with the documented lake expansion–contraction cycles, mudcracks and evaporate minerals (Surdam and Stanley, 1979). Third, spar percentage and the carbonate category of our samples

42

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

Fig. 4. T (47 ) values of carbonate cements (squares) and calculated δ 18 Owater values (circles) in the central Rocky Mountains and western Great Plains. Sample categories are explained in Table 1 and Fig. 3 caption. Values represent averages of 2–3 analyses of each sample. Global deep-sea carbonate δ 18 O values are from Zachos et al. (2001a, 2001b). Dashed and solid lines represent the general trends of T (47 ) values in the Central Rockies, and Western Great Plains, respectively. Q = Quaternary; P = Pliocene.

are not systematically related to the burial depth of the samples (Fig. 2), suggesting that the sparry calcite and carbonate with differing luminescence were not completely formed by mesodiagenesis. Finally, the only sample showing complete calcite recrystallization (EF2-04 Snail) has a T (47 ) value only 4–5 ◦ C higher than the calcite cement of sandstone samples stratigraphically below or above. If the sparry calcite of this sample was completely formed by mesodiagenesis, the T (47 ) value suggests that the burial depth of this sample when the mesodiagenesis occurred was only 200–250 m assuming a geothermal gradient of 25 ◦ C/km. We suggest that the sparry calcite and micritic calcite of differing luminescence were mainly formed during eodiagenesis in vadose and phreatic zones near the surface. Vadose zone cement displays dense micritic fabric and phreatic cement includes blocky spar filling isolated vugs and the remaining pore space surrounded by micrite (Wright and Tucker, 1991; Beckner and Mozley, 1998). Both textures are observed in our samples (Fig. 3). Cathodoluminescence intensity is controlled by minor chemical constituents, primarily Mn2+ and possibly rare earth elements, which activate orange luminescence, and Fe2+ , which quenches luminescence (Habermann et al., 2000; Machel, 2000). The presence of these elements is related to redox and pH conditions of fluid, with the result being that carbonate precipitated under different conditions can be identified from the color of luminescence (Machel and Burton, 1991; Solomon and Walkden, 1985). Vadose meteoric environments are generally well oxygenated and contain low levels of divalent Mn and Fe. Carbonate formed in such an environment exhibits low-intensity luminescence. The phreatic zone is a reducing environment, and thus is likely to contain much higher levels of divalent Mn and display bright luminescence if divalent Fe concentration is low. Patchy variations in luminescence can result from the periodic influx of waters from the phreatic zone into the sediments of the vadose zone (Morad, 1998). Mixed micritic and sparry cement of differing cathodoluminescence colors that were

formed in the vadose and phreatic zones have been documented in the Miocene Zia Formation in New Mexico (Beckner and Mozley, 1998). Therefore, we suggest that the micrite in Category 1 samples, and the carbonate cement of dull or no luminescence in both Categories 2 and 3 samples formed in the vadose zone, and the micritic and sparry cement of orange luminescence formed in the phreatic zone when the sample was only slightly buried or in the vadose zone when the water table fluctuated over time. Phreatic calcite cement is found preferentially in areas that contain early vadose calcite, suggesting cementation was controlled by pre-cementation permeability and presence of nucleating material (Hall et al., 2005). In summary, we argue that the calcite cements of our samples were predominantly formed by eodiagenesis, and mesodiagenesis and telodiagenesis are of limited influence. Therefore, the calcite cement T (47 ) and δ 18 Owater values of our samples primarily reflect near-surface temperature and surface water soon after deposition. 6.2. Formation of dolomite and HMC Sample Lag30-7 from the White River Group in the Great Plains and sample CB49 from the Split Rock Formation in the central Rockies contain microcrystalline dolomite as well as micritic and sparry calcite. Mesodiagenetic dolomite usually exhibits coarse, rhombic crystals with chemical zonation, caused by water– sediment interaction when minerals such as recycled dolomite grains are dissolved in pore water to increase water magnesium concentration (e.g., Solomon and Walkden, 1985; Driese and Mora, 1993). The two samples do not show coarse, rhombic crystals under cathodoluminescence and petrographic microscopy, suggesting the dolomite is not mesodiagenic. Alternatively, the dolomite may have been formed by authigenic processes. In modern terrestrial environments, dolomite occurs

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

43

6.3. Paleotemperature during the middle and late Cenozoic Our cement T (47 ) values provide the first direct temperature record in the western U.S.A. spanning the last ∼50 myr. Our data show that in the central Rockies, T (47 ) values decreased during the middle–late Eocene, remained stable across the Eocene– Oligocene boundary, increased during the late Oligocene, and decreased during the middle–late Miocene (Fig. 4A). Although the T (47 ) record in the Great Plains contains less data and the age gap of the samples is large, the T (47 ) values generally decreased after the Oligocene (Fig. 4B). These general trends are consistent with the global deep-sea temperature (Zachos et al., 2001a, 2001b), mean annual surface air temperature in western U.S.A. based on fossil leaf physiognomy (Wolfe, 1978, 1994), and mean annual temperature in Oregon based on paleosol chemical weathering index (Retallack, 2007). Our data, however, do not clearly show a sharp temperature drop in eastern Wyoming at the Eocene– Oligocene transition (Zanazzi et al., 2007), although the temporal resolution of our data is insufficient to allow for definitive conclusions in this respect. 6.4. Middle and late Cenozoic relief between the Central Rockies and the adjacent Great Plains

Fig. 5. Late Eocene to late Miocene temperature (A) and calculated surface-water δ 18 O values (B) in both the central Rocky Mountains and western Great Plains. Curved black and gray lines represent the mean and maximum/minimum elevation along the transect in Fig. 1A. Numbers in brackets in the legend represent the slopes of regression lines.

in lacustrine environments, saline soils, and soils developed on basalt when the Mg/Ca ratio of lake and soil water is high due to evaporation or bedrock dissolution (e.g., Kohut et al., 1995; Capo et al., 2000; Last et al., 2012). Pedogenic dolomite is also documented in soil nodules formed during the late Paleocene and early Eocene in the Axhandle Basin in Utah (VanDeVelde et al., 2013). Dolomite can be also formed by low-temperature microbial precipitation (Roberts et al., 2004; Sánchez-Román et al., 2008). Sample Lag30-7 was collected from the fluvial deposits, which contain pedogenic features such as branching root traces and leached soil horizons (Retallack, 1983; Terry, 2001). Sample CB46 was also collected from similar rhizolith rich strata. These observations dispute a lacustrine origin, but favor a soil origin of the dolomite. The calculated δ 18 Owater value (−9.8h) is comparable to the calculated δ 18 Owater value of comparable age in the central Rockies, suggesting that the dolomite was not formed in evaporative saline soil conditions. The dolomite was also not formed by low-temperature microbial precipitation because methanogenesis that tends to increase carbonate δ 13 C values (Roberts et al., 2004; Sánchez-Román et al., 2008). The carbonate δ 13 C value of this sample is similar to the values of other calcite cements (Table 2). We therefore suggest this dolomite was formed by increased magnesium concentration by the weathering of mafic volcanic glass to smectite (Evanoff et al., 1992), which is similar to the process forming dolomite soil on basalt (Capo et al., 2000). We suggest the presence of HMC in sample RF4660 formed via a similar process, but with a lower magnesium concentration compared to the other two samples.

Our results show that T (47 ) gradients similar to or greater than today were established between the central Rockies and western Great Plains before the early Oligocene, and lasted into the late Miocene. The gradient of 1.0–2.8 ◦ C per degree of longitude for T (47 ) values during the Oligocene and Miocene is higher than the Quaternary temperature gradient of 0.6 ◦ C per degree of longitude based on T (47 ) of soil carbonates (Hough et al., 2014), and the modern air temperature gradient of ∼0.4 ◦ C per degree of longitude (Fig. 5A). These high temperature gradients were not a result changing atmospheric pCO2 . Atmospheric pCO2 was high during the early Oligocene and decreased to modern values by the end of the Oligocene (Zachos et al., 2001b). High pCO2 should depress the Oligocene temperature gradient if the Oligocene paleorelief was similar to today (Poulsen and Jeffery, 2011). Our results show that the temperature gradient during the Oligocene was the highest (Fig. 5A), opposite of the prediction. Because the eastward increase of Quaternary soil carbonate formation temperatures and modern air temperatures reflect the eastward decreases of mean surface elevation (Fig. 6D) (Hough et al., 2014), these higher T (47 ) gradients suggest the paleorelief between the central Rockies and Great Plains during the middle and late Cenozoic was at least comparable to today (Fig. 5A). By using a temperature lapse rate of −4 ± 1 ◦ C/km, which represents the average and variation of lapse rates of Quaternary soil and lake carbonate T (47 ) values and modern air and lake water temperature in western U.S.A. (Meyer, 1992; Wolfe, 1992; Huntington et al., 2010; Hough et al., 2014), the eastward increase of T (47 ) value of 4–6 ◦ C yield a regional relief of at least 1 km during the Oligocene and Miocene. The variations of the longitudinal gradient of clumped isotope temperature through time may reflect differential vegetation cover (and hence differential heating of the soil surface by solar radiation), and seasonal variation of carbonate formation (e.g., Passey et al., 2010). The high Oligocene T (47 ) gradient may suggest the elevation of the central Rockies was higher during the early Oligocene than today, however, central Wyoming experienced ∼0.6 km of extensional collapse during the latest Miocene–Pliocene (Love, 1970; Steidtmann and Middleton, 1991; Snoke, 1993), and there is no geologic evidence supporting middle and late Oligocene extensional tectonics in central Wyoming. Our results also show that δ 18 Owater gradients similar to or greater than today were established between the central Rockies

44

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

Fig. 6. Plots showing the relationship between δ 18 O and δ 13 C values of carbonate cement (A), calculated water δ 18 O and T (47 ) values (B), calculated water δ 18 O values and sampling elevation (C), T (47 ) values and sampling elevation (D). Grey lines represent the regression of Quaternary soil carbonate values in Hough et al. (2014), and black lines represent regressions of late Eocene–Miocene data presented in this study. OB1-25 is not included in the regression due to the influence of evaporation.

and western Great Plains before the Oligocene. The gradient of 0.8–1.8h per degree of longitude for δ 18 Owater values during the Oligocene and Miocene is slightly greater than the Quaternary precipitation gradient of 0.6h per degree of longitude (Hough et al., 2014), and modern precipitation gradient of ∼1.3h per degree of longitude (Vachon et al., 2010) (Fig. 5B), both suggesting the regional relief was at least 1 km. The variations of the longitudinal gradient of the δ 18 Owater values could be caused by the degree of mixing between local precipitation in the Great Plains and distal groundwater charged by snowmelt from the central Rockies (Dutton et al., 2005) as well as climate change on million-year scale. Our carbonate clumped isotope and oxygen isotope results support a hypothesis that the high elevation gradient between the central Rockies and the western Great Plains has existed as a major topographic feature since at least the early Oligocene with mean elevation contrast comparable with modern topography. This is in good agreement with the smectite stable oxygen isotope paleoaltimetry results (Sjostrom et al., 2006) and volcanic glass δ D values (Fan et al., 2014). Because the Laramide intermontane basins were near sea level during the early Eocene (Fan et al., 2011), data presented here suggest that the basin floors in the central Rockies gained much of the elevation between the early Eocene and early Oligocene. We suggest this surface uplift may have occurred during the late Eocene, coincident with

the widespread depositional hiatus or erosion during the late Uintan–Duchesnean (42–37 Ma) in Wyoming (Lillegraven, 1993; Cather et al., 2012), and the widespread erosional unconformity between the latest Eocene–Oligocene White River Group and the Late Cretaceous Pierre Shale in Nebraska and South Dakota (Retallack, 1983; Terry and LaGarry, 1998). Our data, however, are not compatible with the continued subsidence of Laramide basins until the late Miocene (e.g., Bradley, 1987; McMillan et al., 2002, 2006; Leonard, 2002). Late Eocene uplift of the central Rockies may be attributed to isostatic rebound and/or mantle dynamics induced by the foundering of lower mantle lithosphere (Roberts et al., 2012) or the Farallon slab (Roy et al., 2009; Liu and Gurnis, 2010; Cather et al., 2012), both of which have been interpreted to cause the middle Cenozoic uplift of the Colorado Plateau and surrounding Rocky Mountains. Isostatic rebound induced by significant erosion of the Laramide mountain ranges during the middle and late Eocene may also contribute to the uplift of the basin floors as well mountain ranges. Our results also do not preclude additional small-magnitude modifications to topography after the latest Miocene. Such modifications include latest Miocene–Pliocene dynamic uplift of the central Rockies caused by mantle dynamic processes (McMillan et al., 2006; Karlstrom et al., 2011; Duller et al., 2012), and collapse of the Granite and Owl Creek mountains and the south end of the Wind River Range (Love, 1970; Steidtmann and Middleton, 1991).

M. Fan et al. / Earth and Planetary Science Letters 408 (2014) 35–47

7. Conclusions We use the T (47 ) and δ 18 O values of carbonate cement to constrain the temperature changes and the paleorelief between the central Rockies and the adjacent Great Plains during the middle and late Cenozoic. Our results bear on three important issues pertaining to the application of clumped isotope geothermometry to paleotemperature and paleotopographic reconstructions in the western U.S.A. First, sampled carbonate cements are dominantly micritic low-magnesium calcite, with a variable amount of sparry calcite present in most samples and intermixed dolomite and highmagnesium occurring in only three out of twenty five samples. Based on petrographic observations, these carbonates were placed into two categories representing different degrees of mixing of vadose and phreatic carbonates formed in near-surface conditions. Our careful examination of carbonate texture suggests that the influence of mesogenesis and telogenesis on these carbonate cements is limited, and thus that these carbonates retain original information regarding near-surface environmental conditions in the study area. Second, our clumped isotope temperatures show parallel decreases of ∼10 ◦ C, and calculated water δ 18 O values show parallel decreases of ∼4h from ∼35 Ma to ∼5 Ma in both the central Rockies and western Great Plains, similar to the middle and late Cenozoic global cooling trend and the decreases of mean annual temperature in western U.S.A. reconstructed using leaf physiognomy. Finally, both the measured T (47 ) values and calculated water δ 18 O values increase eastward with similar-to-modern or higher-than-modern longitudinal gradients, a pattern beginning as early as the early Oligocene. This observation suggests the differential uplift of the central Rockies with respect to the Great Plains occurred before the early Oligocene. In combination with other published paleoaltimetry results and the widespread depositional hiatus during the late Eocene, these data suggest that the differential uplift the central Rockies occurred during the late Eocene. Our data do not support the interpretation that the Rockies experienced continuous subsidence until the late Miocene; rather they support the interpretation that the high mean elevation of the central Rockies was established during the late Eocene, in association of mantle lithosphere delamination or Farallon slab removal. Acknowledgements This research was funded by NSF EAR1119005 and a UT Arlington research enhancement grant 14-7489-05. We thank Dr. Robert Diffendal and Sarah Allen for assisting in fieldwork. We thank Gregory Henkes, Shuning Li, Haoyuan Ji, and Gabrielle Stephens for help with the clumped isotope analyses. We also thank Dr. Steven G. Driese, Lauren Michel, and Emily Beverly for the permission and assistance of using CL facility at the Baylor University. We are grateful to three anonymous reviewers for thoughtful reviews, which helped us improve this paper. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2014.09.050. References Anders, M.H., Saltzman, J., Hemming, S.R., 2009. Neogene tephra correlations in eastern Idaho and Wyoming: implications for Yellowstone hotspot-related volcanism and tectonic activity. Geol. Soc. Am. Bull. 121, 837–856. Angevine, C.L., Flanagan, K.M., 1987. Buoyant sub-surface loading of the lithosphere in the Great Plains foreland basin. Nature 327, 137–139. Anna, L.O., 2005. Geologic assessment of undiscovered oil and gas in the Powder River Basin Province, Wyoming and Montana. In: U.S. Geological Survey Powder River Basin Assessment Team (Ed.), Total Petroleum Systems and Geologic Assessment of Oil and Gas Resources in the Powder River Basin Province,

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