Rapid Middle Eocene temperature change in western North America

Earth and Planetary Science Letters 450 (2016) 132–139

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

Rapid Middle Eocene temperature change in western North America Katharina Methner a,∗ , Andreas Mulch a,b , Jens Fiebig b , Ulrike Wacker b , Axel Gerdes b , Stephan A. Graham c , C. Page Chamberlain d a

Senckenberg Biodiversity and Climate Research Centre (BiK-F), Senckenberganlage 25, 60325 Frankfurt am Main, Germany Institut für Geowissenschaften, Goethe-Universität, Altenhöferallee 1, 60438 Frankfurt am Main, Germany c Department of Geological and Environmental Sciences, Stanford University, 450 Serra Mall, Bldg. 320, Stanford, CA, 94305, USA d Department of Environmental Earth System Science, Stanford University, 473 Via Ortega, Rm. 140, Stanford, CA, 94305, USA b

a r t i c l e

i n f o

Article history: Received 19 November 2015 Received in revised form 27 May 2016 Accepted 28 May 2016 Available online xxxx Editor: M. Frank Keywords: paleoclimate Eocene stable isotopes clumped isotope thermometry western North America

a b s t r a c t Eocene hyperthermals are among the most enigmatic phenomena of Cenozoic climate dynamics. These hyperthermals represent temperature extremes superimposed on an already warm Eocene climate and dramatically affected the marine and terrestrial biosphere, yet our knowledge of temperature and rainfall in continental interiors is still rather limited. We present stable isotope (δ 18 O) and clumped isotope temperature (47 ) records from a middle Eocene (41 to 40 Ma) high-elevation mammal fossil locality in the North American continental interior (Montana, USA). 47 paleotemperatures of soil carbonates delineate a rapid +9/−11 ◦ C temperature excursion in the paleosol record. 47 temperatures progressively increase from 23 ◦ C ± 3 ◦ C to peak temperatures of 32 ◦ C ± 3 ◦ C and subsequently drop by 11 ◦ C. This hyperthermal event in the middle Eocene is accompanied by low δ 18 O values and reduced pedogenic carbonate concentrations in paleosols. Based on laser ablation U/Pb geochronology of paleosol carbonates in combination with magnetostratigraphy, biostratigraphy, stable isotope, and 47 evidence, we suggest that this pronounced warming event reflects the Middle Eocene Climatic Optimum (MECO) in western North America. The terrestrial expression of northern hemisphere MECO in western North America appears to be characterized by warmer and wetter (sub-humid) conditions, compared to the post-MECO phase. Large and rapid shifts in δ 18 O values of precipitation and pedogenic CaCO3 contents parallel temperature changes, indicating the profound impact of the MECO on atmospheric circulation and rainfall patterns in the western North American continental interior during this transient warming event. © 2016 Elsevier B.V. All rights reserved.

1. Introduction Short-term hyperthermals such as the Paleocene–Eocene Thermal Maximum (PETM) or the Eocene Thermal Maximum 2 (ETM2) reflect rapid global warming events and are a recurrent phenomenon of early and middle Eocene climate dynamics (e.g., Zachos et al., 2008). Both events punctuated an already warm Paleogene climate, interrupted the long-term mid-to-late Eocene cooling trend and coincided with perturbations in the global carbon cycle (e.g., Bowen et al., 2015; Sluijs et al., 2013; Zachos et al., 2008). In contrast to the short-lived (<100 ka) Early Eocene hyperthermals, the Middle Eocene Climatic Optimum (MECO) about 40 Ma ago is unique in terms of its long duration (500–750 ka) and gradual temperature increase (Bohaty and Zachos, 2003; Bohaty et al., 2009; Edgar et al., 2010). A suite of globally distributed

*

Corresponding author. Tel.: +49 69 7542 1815. E-mail address: [email protected] (K. Methner).

http://dx.doi.org/10.1016/j.epsl.2016.05.053 0012-821X/© 2016 Elsevier B.V. All rights reserved.

deep-sea sediment cores revealed that this warming event was characterized by a transient temperature increase of 3 ◦ C to 6 ◦ C of both surface and intermediate deep-waters, a brief global shoaling of the calcite compensation depth due to acidification of the oceans, and shifts in species distribution and abundance patterns (Bijl et al., 2010; Bohaty and Zachos, 2003; Bohaty et al., 2009; Boscolo Galazzo et al., 2013; Edgar et al., 2010; Witkowski et al., 2012). An increase in atmospheric CO2 concentrations by a factor of 2–3 compared to pre- and post-MECO pCO2 levels further characterizes the MECO (Bijl et al., 2010). The postulated global rise in pCO2 and temperature, accounting for deep and shallow ocean warming, should exhibit substantial influence on continental environments as increasing mean global temperature will affect atmospheric and ocean circulation dynamics and thus alter continental rainfall patterns (e.g., Allen and Ingram, 2002; Chou et al., 2013; Held and Soden, 2006; Marvel and Bonfils, 2013). However, the implications of continental temperature change during the MECO on continental hydrology are still largely elusive: Palynological and

K. Methner et al. / Earth and Planetary Science Letters 450 (2016) 132–139

sedimentological data in Central Asia indicate a major aridification step associated with monsoonal intensification possibly as a consequence of post-MECO cooling (Bosboom et al., 2014). Oxygen (δ 18 O) and carbon (δ 13 C) isotope analyses of lacustrine carbonates (Elko Basin, NV, USA) reveal large shifts in lake water δ 18 O, most likely due to rapid lake water freshening at the climax of the MECO and indicate that temperature seasonality was sufficiently strong to (re-)establish regular lake overturn (Mulch et al., 2015). These two terrestrial MECO records indicate changes in the hydrological cycle of northern hemisphere mid-latitudes in concert with either surface uplift processes of the Tibetan Plateau for Central Asia and a southward encroachment of an Eocene plateau (SWEEP) in western North America (Mix et al., 2011) and/or retreats of epicontinental seas (proto-Paratethys and Mississippi Embayment, respectively). Redistribution of rainfall patterns, and thus δ 18 O values of precipitation, results from changes in atmospheric circulation and hydrological cycling that are both sensitive to global warming (e.g., Marvel and Bonfils, 2013). Despite being a key element in global climate dynamics, continental rainfall patterns in a warmer world are far from being understood and predictions range from wet-getwetter and dry-get-drier patterns (e.g., Allen and Ingram, 2002; Chou et al., 2013; Held and Soden, 2006) to an overall (midlatitude) drying in a warmer future (Sherwood and Fu, 2014). In order to evaluate the impacts of global warm periods on terrestrial temperature and rainfall records we sampled a suite of middle Eocene (∼40 Ma) paleosols in the hinterland of the North American Cordillera. Stable isotope (δ 18 O) and clumped isotope paleo-temperature (47 ) data, together with laser ablation inductive coupled plasma-mass spectrometry (LA-ICP-MS) U–Pb dating of carbonates permits to develop a middle Eocene terrestrial temperature record and investigate changes in the hydrological cycle and consequently in atmospheric circulation patterns over western North America during an important Cenozoic warm period. 2. Geological setting Beginning in the early Eocene, the North American continental interior underwent extension at various levels of the crust with accompanying magmatism, formation of metamorphic core complexes, and basin formation within the North American Cordillera (e.g., Constenius, 1996; Dickinson, 2002; Fields et al., 1985; Sonder and Jones, 1999). Most basins in western Montana (and east– central Idaho) resulted from Paleogene extensional deformation in the hinterland of the Sevier fold-and-thrust belt (e.g., Constenius, 1996; Fields et al., 1985; Rasmussen, 2003) and basins evolved as intermontane half-grabens in a north–south trending rift system (Janecke, 1994; Sears and Ryan, 2003). In an overall extensional context, the Sage Creek Basin (MT) has been described as an extensional intermontane basin (Constenius, 1996), a flexural basin (Janecke et al., 1999), as part of the “Renova braidplain” on the uplifted eastern rift shoulder (Sears and Ryan, 2003) or as part of a continuous depositional system of the DillonRenova Basin, flanked by volcanic highlands (Fritz et al., 2007). Independent of the mode of basin formation sedimentation in the Sage Creek Basin started in the early to middle Eocene and lasted until the mid-Miocene, thus making the Sage Creek Basin one of the temporally most extensive basins in western Montana that also yielded a vast amount of vertebrate fossils spanning the Bridgerian (∼51 Ma) to Barstovian (∼12 Ma) NALMAs (North American Land Mammal Ages) (Fields et al., 1985; Tabrum et al., 1996). Age constraints of the Sage Creek Basin deposits are based on biostratigraphic, paleomagnetic and radiometric dating (Fields et al., 1985; Kent-Corson et al., 2006; M’Gonigle and Dalrymple, 1996; Tabrum et al., 1996). Based on vertebrate fossil assemblages, Paleogene sediments exposed in the Sage Creek Basin can be divided into the Bridgerian Sage Creek Formation (∼51.0–46.3 Ma), the

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Uintan Dell Beds (46.3–40.0 Ma), and the Chadronian–Orellan Cook Ranch member (37–32 Ma) (see compilations of Fields et al. (1985) and Tabrum et al. (1996); ages in parenthesis indicate age limits of NALMA in Woodburne (2004), modified for the Bridgerian after Smith et al. (2008)). Even though southwestern Montana local faunas reflect high levels of provincialism/endemism, which make correlations with other North American mammal assemblages challenging (Tabrum et al., 1996), these age zones were confirmed by paleomagnetic (Tabrum et al., 1996) and geochronologic analyses (Fritz et al., 2007; Kent-Corson et al., 2006; M’Gonigle and Dalrymple, 1996; Rothfuss et al., 2012). Within this depositional context we collected micritic carbonate samples and wherever present pedogenic carbonate nodules from different pedogenic horizons of the well characterized middle Eocene Upper Dell Beds (Sage Creek Basin, SW Montana, USA; Fig. 1; e.g., Kent-Corson et al., 2006, 2010; Retallack, 2007, 2009; Tabrum et al., 1996). The Dell Beds unconformably overlie the early Eocene Sage Creek Formation and represent a sequence of stacked paleosols, which mainly consist of tuffaceous mudstone interbedded with sandstone and minor conglomerates (Tabrum et al., 1996). The sampling locality (corresponding to the “Kay Draw” or “Hough Draw 1” locality of Tabrum et al., 1996; Fig. 1c) reflects the stratigraphically highest part of the Dell Beds and consists of tuffaceous mud- and siltstones with abundant paleosol formation. Pedogenic accumulation of carbonate occurred in B horizons (Hanneman et al., 2003); these range in thickness from 0.1 m to 0.4 m and are typically well indurated. A shallow carbonate formation environment is reported by Retallack (2007, 2009) with carbonate formation depths of the Dell Beds (at Douglass Draw) estimated to range between 27 cm and 120 cm, averaging 54 cm ± 22 cm (Retallack, 2007, 2009, and downloadable Excel files therein). We measured a composite section covering about 50 m (Figs. 1c and 1d, see Appendix for field photographs and detailed sections) and analyzed micritic paleosol carbonate and pedogenic carbonate nodules for δ 18 O and clumped isotope analyses as well as U–Pb geochronology. 3. Methods 3.1. Stable isotope analyses Oxygen (δ 18 O), carbon (δ 13 C) and clumped (47 ) isotope analyses of pedogenic carbonates were performed at the Joint Goethe University – BiK-F Stable Isotope Facility at the Institute of Geosciences, Goethe University Frankfurt, Germany. Carbonate powders were drilled at 1 to 5 different spots within the same hand specimen using a low-speed dental drill. After digestion of 0.1 mg to 0.5 mg carbonate powder with H3 PO4 at 72 ◦ C in a sealed reaction vessel flushed with helium gas, the evolved CO2 was sampled by a Finnigan Gas Bench and isotope ratios were measured on a Finnigan MAT 253 mass spectrometer. Repeated measurements of an in-house standard (Carrara marble) yielded external precision of <0.07h for δ 18 O and <0.03h for δ 13 C. All isotopic results are reported in standard delta notation and corrected to VSMOW (δ 18 O) or VPDB (δ 13 C). Carbonate contents (in %) were calculated by relating the signal size of the samples and the averaged signal size of the daily standards (Carrara marble, n = 12) and assuming 100% CaCO3 for the Carrara marble (Appendix Tabs. A.1 and A.2). The oxygen isotopic composition of soil water (δ 18 Owater ) was calculated from 47 temperature and δ 18 Ocarbonate value pairs using the oxygen isotope fractionation coefficient of Kim and O’Neil (1997) (Appendix Tab. A.9). 3.2. Clumped (47 ) isotope analyses Clumped isotope analyses were performed using a Thermo Finnigan MAT 253 mass spectrometer modified to measure a mass

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Fig. 1. (A) Overview map of the western United States. Star marks the location of Fig. 1B. (B) Geological map of SW Montana, USA, with sampling locality of the Upper Dell Beds, southern Sage Creek Basin. (C) Dell Beds at the Kay Draw locality. The sections are composed of a series of stacked paleosols. The section exposes the disconformable contact between the Dell Beds and the overlying Oligocene Cook Ranch member (Tabrum et al., 1996). (D) Composite section of the Dell Beds. Dots mark carbonate samples (orange for 47 samples), stars denote the samples for U–Pb dating. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

range of 44 to 49. Carbonate digestion, CO2 purification procedure and measurements follow Wacker et al. (2013, 2014). In short, for 47 analyses 8–28 mg of carbonate were reacted at 90 ◦ C in >106% H3 PO4 , and the liberated CO2 was purified by passage through multiple cryogenic traps, including a gas chromatographic column held at −20 ◦ C. “Heated gases” (equilibrated at 1000 ◦ C) were measured to monitor the non-linearity of the mass spectrometer and source scrambling effects. CO2 gases, equilibrated at 25 ◦ C, and “heated gases” were subsequently used to determine the empirical transfer functions (ETF) (Appendix Tabs. A.3–A.5). We applied the theoretical acid fractionation factor of +0.069h (Guo et al., 2009) to the data obtained from carbonates reacted at 90 ◦ C to allow comparison to carbonates digested at 25 ◦ C (see Wacker et al., 2014). 47 values are reported in the absolute reference frame for interlaboratory comparison (Dennis et al., 2011). Temperatures are calculated using the calibration of Wacker et al. (2014):

47 = 0.0327 (±0.0026) ∗ 106 / T 2 + 0.3030 (±0.0308)

(i)

with 47 in h and T in K. This regression line (i) is within analytical error statistically indistinguishable from (a) the empirically determined calibration line of Henkes et al. (2013) and (b) the theoretical calibration line (Guo et al., 2009). These temperature-47 relationships, however, differ from the original calibration (Ghosh et al., 2006) and several other calibration lines (e.g., Tripati et al., 2010; Zaarur et al., 2013). The cause(s) of these discrepancies are still not fully understood. We refer to the calibration of Wacker et al. (2014) that has been determined at Goethe University Frankfurt since this provides that all analyses are performed under similar analytical conditions (digestion at 90 ◦ C; similar gas preparation and data reduction protocols; measurements performed at the same mass spectrometer). In addition, 47 values of standard materials (NBS 19, Carrara marble) and of a calibration sample (Dyscolia wyvillei) obtained during this study match (within standard error) the 47 values determined during the calibration study of Wacker et al. (2014) (Appendix Tab. A.4).

3.3.

238

U–206 Pb dating of pedogenic carbonates

As pedogenic carbonates are common paleoclimate and paleoenvironmental proxy materials direct dating of carbonates has a huge potential for improving the timing of climatic events (e.g., Rasbury et al., 1998). We acquired U–Pb ages in situ on polished thick sections of two pedogenic carbonate nodules (samples 11-KM-124 and 11-KM-122) by laser ablation sector field ICP-MS at Goethe University Frankfurt using a modified method previously described in Gerdes and Zeh (2006, 2009). Uranium, thorium and lead isotopes were analyzed using a Thermo-Scientific Element 2 coupled to a Resolution S-155 (Resonetics) 193 nm ArF Excimer laser (CompexPro 102, Coherent) equipped with a two-volume ablation cell (Laurin Technic, Australia). Samples were ablated in a helium atmosphere (0.6 l/min) and mixed in the ablation funnel with 0.7 l/min argon and 0.03 l/min nitrogen. Static ablation used a spot size of 235 μm and a fluence of <1 J cm−2 at 8 Hz. For the standard material SRM-NIST 614 this setup yielded a depth penetration of about 0.5 μm s−1 and an average sensitivity of 280,000 cps/μg g−1 for 238 U. The detection limit for 206 Pb and 238 U was ∼0.5 and 0.03 ppb, respectively. Each analysis consists of 20 s background acquisition followed by 20 s of sample ablation and 20 s washout. Prior to analysis each spot was pre-ablated for 3 s to remove surface contamination. Soda-lime glass SRM-NIST 614 was used as a reference glass together with 2 carbonate standards to bracket sample analysis. Raw data were corrected offline using an in-house MS Excel© spreadsheet program (Gerdes and Zeh, 2006, 2009). Following background correction, outliers (±2σ ) were rejected based on the time-resolved 207 Pb/206 Pb and 206 Pb/238 U ratios. The 207 Pb/206 Pb ratio was corrected for mass bias (0.6%) and the 206 Pb/238 U ratio for inter-element fraction (∼5%), including drift over the sequence time, using SRM-NIST 614. Due to the carbonate matrix an additional off-set factor of 0.921 has been applied, which was determined using WC-1 carbonate reference material dated by TIMS (251 ± 2 Ma; E.T. Rasbury, pers. comm., 2014). Repeated analyses (n = 24) of a Zechstein dolomite (Gypsum pit, Tettenborn, Germany) used as secondary (in-house) stan-

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Fig. 2. Stratigraphic, stable isotope, and clumped isotope record of the Dell Beds, Montana. (A) Clumped isotope (47 ) temperatures of pedogenic mudstones (squares) and nodules (circles) in [◦ C] with 2σ standard errors. n denotes the number of replicate measurements for each sample. (B) Carbonate content of pedogenic mudstone in [%]. (C) Oxygen [h, VSMOW] isotope record. Colored symbols represent average values from individual measurements (grey). Stars and grey lines mark the position of U–Pb geochronology samples. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

dard yielded a lower intercept age of 257.0 ± 4.7 Ma (MSWD: 1.3), implying an accuracy and repeatability of the method of ∼2%. 4. Results

47 values of pedogenic micritic carbonate in soil nodules and pedogenic mudstone range from 0.654h to 0.685h, translating into temperatures of 32.2 ◦ C to 19.4 ◦ C (Fig. 2a; Appendix Tabs. A.6 and A.7). Each soil carbonate sample was analyzed 4 to 6 times and in general external standard errors for replicate measurements range from 0.002h to 0.006h, which translate into temperature errors of 0.8 ◦ C to 3.2 ◦ C. Reproducibility of measurements was very good and all analyzed samples were included in this study. Only one sample (11-KM-116) yielded a larger standard error (0.013h = 5.1 ◦ C) for the mean 47 value after six replicate measurements. The large(r) spread in this particular sample may result from (1) short and long term variations of mass spectrometric parameters, (2) heterogeneities in the phosphoric acid reaction environment, and (3) inhomogeneity in the clumped and bulk isotopic compositions of sample. In the lower 15 m of section 47 temperatures of disseminated pedogenic carbonate are 20.6 ◦ C ± 5.1 ◦ C (at 4.2 m of section) and 23.2 ± 3.2 ◦ C (12.8 m) and are thus indistinguishable from 47 temperatures of coexisting pedogenic carbonate nodules (22.0 ◦ C ± 1.7 ◦ C and 22.9 ◦ C ± 0.8 ◦ C at 12.9 m). Up-section there is a clear increase in temperature to maximum temperatures of 29.5 ◦ C ± 1.9 ◦ C and 32.2 ◦ C ± 2.7 ◦ C at 17.1 m and 21.0 m, respectively, followed by a drop of about 11 ◦ C to temperatures between 21.2 ◦ C ± 1.9 ◦ C (23.9 m) and 19.4 ◦ C ± 2.2 ◦ C (27.1 m). Since pedogenic carbonate nodules of the two most prominent nodular bands in the studied section (at 12.9 m and 14.6 m–15.7 m, respectively) reveal an internal variability in their oxygen isotopic composition (Appendix Tab. A.2), we obtained two 47 temperatures from the same nodule. Sample 11-KM-122A and sample 11-KM-122B yield temperatures of 22.0 ◦ C ± 1.7 ◦ C (mean δ 18 O of 18.6h, n = 5) and 22.9 ◦ C ± 0.8 ◦ C (mean δ 18 O of 14.8h, n = 4), respectively, and are thus indistinguishable from each other in terms of their temperature record.

δ 18 O and δ 13 C values as well as carbonate contents (%CaCO3 ) (Figs. 2b and 2c; Appendix Tabs. A.1 and A.2) in the paleosols corroborate the overall pattern of low δ 18 O and elevated δ 13 C values in Eocene paleosols of the western United States (e.g., compilation in Chamberlain et al., 2012). At first sight, δ 13 C values appear high for an Eocene C3-dominated ecosystem (e.g., Cerling and Quade, 1993). However, under elevated Eocene soil temperatures and high (1000–3000 ppm) atmospheric pCO2 values, measured δ 13 C values yield soil respiration rates (e.g., Caves et al., 2014; Cerling, 1984) that are not unrealistic for early Cenozoic mid latitude continental interiors. In agreement with regional paleobotanical records indicating overall forest-dominated ecosystems and a long-term transition from wet, warm paratropical to dry, open ecosystems (e.g., DeVore and Pigg, 2010) we therefore attribute the comparatively high δ 13 C values to low soil respiration rates (cf., Kent-Corson et al., 2010). Collectively, δ 18 O and %CaCO3 data show a broadly similar pattern with low %CaCO3 corresponding to low δ 18 O values and vice versa. Within 0–8 m of section mean δ 18 O values of pedogenic carbonate are low (12.0h to 12.9h). A major positive shift occurs at 10.3 m of section with δ 18 O values reaching 18.7h and with highest %CaCO3 values of ∼47% at 12.8 m, followed by a rapid drop in δ 18 O and %CaCO3 values between 12.8 m and 14.4 m with (δ 18 O) = −6.4h and (%CaCO3 ) = −25%, respectively. Relatively low carbonate contents in the paleosols (22 ± 4%) and δ 18 O values of 12.3h to 12.8h characterize the high-47 temperature interval (17–21 m). Up-section (21.0–22.9 m), δ 18 O values and %CaCO3 increase to 19.2h and 42%, respectively, marking a second shift of (δ 18 O) = +6.4h and (%CaCO3 ) = +20% coinciding with a major drop in 47 temperatures. After reaching maximum δ 18 O values (19.4h between 22.9–27.1 m), δ 18 O values remain at 14–16h with %CaCO3 values at about 40%. Calculated soil water δ 18 O values based on the individual 47 temperatures range from −9.7h to −16.9h (Appendix Tab. A.9). Nodules used for U–Pb analysis (n = 2) were sampled slightly below the detected temperature excursion in the Dell section (12.9 and 14.6 m). Using laser spot sizes of 213 μm a combination of matrix carbonate, calcite inclusions in the nodule (sample 11-KM-124; n = 61 U–Pb laser spot analyses) and syngenetic frac-

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Fig. 3. Tera–Wasserburg diagrams of paleosol samples. (A) 238 U/ 206 Pb-207 Pb/206 Pb plot of sample 11-KM-124 (at 14.6 m of section) yields an isochron age of 39.5 ± 1.4 Ma with a MSWD (mean squared weighted deviates) of 0.89 and (B) 238 U/ 206 Pb-207 Pb/ 206 Pb plot of sample 11-KM-122 (at 12.9 m of section) yields an isochron age of 40.1 ± 0.8 Ma with a MSWD of 1.15. All uncertainties are reported at the 2σ level.

ture fills (sample 11-KM-122; n = 55 U–Pb laser spot analyses) provided adequate spread in the parent–daughter isotope ratio to allow calculation of U–Pb ages (Fig. 3; Appendix Tab. A.8). Analyzed fracture fills occur only within the nodule (not cross-cutting it) and are interpreted as syngenetic desiccation cracks (see discussion below). U–Pb ages of 40.1 ± 0.8 Ma (sample 11-KM-122, 12.9 m of section; Fig. 3) and 39.5 ± 1.4 Ma (11-KM-124, 14.6 m) are identical within analytical uncertainty. U–Pb data were plotted in Tera–Wasserburg diagrams and ages were calculated as lower intercepts using Isoplot 3.71 (Ludwig, 2007; uncertainties are reported as 2σ ). 5. Discussion 5.1. Soil carbonate formation and 47 temperatures Soil temperatures in central Montana increase rapidly by ca. 9 ◦ C from 23.2 ◦ C ± 3.2 ◦ C (12.8 m) to maximum temperatures of 32.2 ◦ C ± 2.7 ◦ C (21.0 m) before they drop by ∼11 ◦ C (Fig. 2a). This soil temperature excursion occurs over a brief interval of only 11 m of section (12.8–23.9 m) and is accompanied by low CaCO3 contents in the paleosols and low δ 18 O values in nodular and disseminated pedogenic carbonate (Fig. 2). The near-surface temperatures (ranging from 19.4 ± 2.2 ◦ C to 32.2 ◦ C ± 2.7 ◦ C) recorded in both carbonate nodules and pedogenic mudstone document that the analyzed carbonate component formed within the soil column and did not experience any latediagenetic overprint. In addition, the observed temperature excursion occurs within only ∼11 m of section without any discernable change in lithology, diagenesis or fluid flow and is thus difficult

to explain by means of localized non-pedogenic alteration processes. Further, the U–Pb ages show that the inner parts of the (septarian) nodules formed at ∼40 Ma and thus at the estimated time of sediment deposition based on vertebrate fossils and paleomagnetic data (Tabrum et al., 1996) and not as a consequence of late-diagenetic processes. Therefore, we exclude diagenesis to have played a role in setting the 47 carbonate formation temperatures. Soil carbonate formation typically occurs when soils dewater during warm and dry periods and requires arid to semi-arid (typically <800 mm/a), in some cases sub-humid or monsoonal climates accompanied by precipitation seasonality (e.g., Breecker et al., 2009). Consequently, 47 temperature information in pedogenic carbonates might be biased towards warm month mean temperatures (WMMT; e.g., Hough et al., 2014; Passey et al., 2010; Peters et al., 2013; Quade et al., 2013) or, if soil dewatering occurs in non-summer months may be closer to mean annual air temperatures (MAAT; e.g., Peters et al., 2013). Quantifying how much of the observed variability in soil temperatures (a) is the result of global or local temperature change (e.g., by surface elevation change), (b) may be due to the timing of carbonate formation during marked seasonality, or (c) reflects changes in soil carbonate formation depth remains challenging. Major drainage reorganization, related to topographic adjustments of a high-elevation plateau upstream of the Sage Creek Basin (Chamberlain et al., 2012; Mix et al., 2011), already occurred at about 49 Ma resulting in low δ 18 O of regional precipitation prior to the middle Eocene (Carroll et al., 2008; Kent-Corson et al., 2006). In contrast, the δ 18 O excursion of up to −6h at about 40 Ma in the Dell Beds record is (a) not unidirectional and (b) too rapid to be explained by surface elevation changes and thus requires alternative explanations including modification of local soil temperature and/or atmospheric circulation patterns and associated changes in δ 18 O values of precipitation over southwestern Montana. Paleotopography and paleoclimate modeling studies reveal the presence of an Eocene plateau spanning western North America at times of a monsoonal climate with a summer precipitation season (Chamberlain et al., 2012; Huber and Goldner, 2012; Feng et al., 2013; Sewall et al., 2000). Thus, soil dewatering and carbonate formation were likely to occur during or shortly after the main (summer) rainfall season. 47 carbonate formation temperatures are significantly higher when compared to modeled (7.3 ◦ C, Feng et al., 2013) and calculated (13.4 ◦ C, Mix et al., 2011) MAAT, consistent with warm season soil carbonate formation. We thus consider it unlikely that the major part of the temperature shift is due to changes in carbonate formation seasonality biased towards warmer summer months alone. Differences in seasonality and depth of carbonate formation occur mainly through changes of the precipitation season and precipitation amounts (e.g., Breecker et al., 2009; Retallack, 2005) with a positive correlation between precipitation amount and depth of the carbonate accumulation horizon, Bk (e.g., Retallack, 2005). We exclude reduced soil carbonate formation depths as the main driver for the high 47 temperature interval, because estimates of Bk depth from the Dell Beds lack pronounced perturbations throughout the measured section by Retallack (2007, 2009). More importantly, the contemporaneous low δ 18 O and low %CaCO3 values argue against increasingly evaporative conditions, which we would expect for upper soil horizons at soil temperatures exceeding 30 ◦ C. Even though low δ 18 O values and low CaCO3 contents characterize the high-47 temperature interval (12.8-23.9 m), the former do not correlate with 47 temperatures (Fig. 2) on the scale of the entire section thus excluding soil temperature changes as the sole driver for δ 18 O changes in the pedogenic carbonate. The effect of lowering carbonate δ 18 O values by increasing soil temperature (and thus fractionation) is minor (∼0.25h/◦ C;

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Kim and O’Neil, 1997) and is typically balanced by the effects of temperature on precipitation δ 18 O values (e.g., Rozanski et al., 1993) and soil evaporation, both increasing soil carbonate δ 18 O (e.g., Cerling and Quade, 1993). Thus, major changes in the hydrological cycle and the atmospheric circulation pattern are required to explain the warm, low-δ 18 O/low-%CaCO3 interval. When placed in a longer-term temporal framework depth-to-Bk measurements seem to indicate changes in rainfall intensity and greenhouse gas concentrations in mid-latitude North America at around 40 Ma (Sheldon et al., 2002; Retallack, 2009) coincident with the %CaCO3 , δ 18 O and 47 data presented here; a result that should be investigated further in the future. 5.2. Timing of soil carbonate formation U–Pb ages of 39.5 ± 1.4 Ma and 40.1 ± 0.8 Ma indicate a middle Eocene age for soil carbonate formation. As the dated fracture fills represent syngenetic desiccation cracks, which formed during drying processes in the paleosols, we suggest that these fractures were filled during nodule formation consistent with very early crystallization and growth of the dated calcite nodules. A much earlier formation of the nodules (and hence sedimentation age of the Dell Beds) compared to the U–Pb ages of the (septarian) fracture fills is precluded by the occurrence of late Uintan (∼43 to 40 Ma) mammal fossils (Tabrum et al., 1996) in the sedimentary sequence. The U–Pb ages are further in perfect agreement with paleomagnetic and biostratigraphic (vertebrate fossils) studies and now provide the first direct geochronological evidence supporting a middle Eocene age of the Dell Beds (Tabrum et al., 1996): Based on North American Land Mammal Ages, the upper part of the Dell Beds (Hough Draw local fauna, which includes 19 mammalian taxa) is assigned to be latest Uintan (Tabrum et al., 1996), which is set to 42.8–40.0 Ma (Ui3, late Uintan; after Woodburne, 2004). A paleomagnetic section (∼40 m of continuous section) was measured at the sampling locality by D.R. Prothero in 1987 (cited in Tabrum et al., 1996) and despite two samples the entire section is of reversed polarity. We relate our sampled section to the paleomagnetic data by tying the top of both sections that is marked by a massive paleosol at the contact to the overlying Oligocene Cook Ranch member (Fig. 1C). Our composite section of about 50 m of Dell Beds expands further downward thus covering a slightly larger time interval. Based on the local fossil fauna, the paleomagnetic section has been interpreted to correlate with the later part of Chron C18r (Tabrum et al., 1996; C18r = 41.15 Ma to 40.15 Ma (GTS12; Ogg, 2012)). In summary, U–Pb ages, NALMA assignments and magnetostratigraphy collectively argue for a Middle Eocene age of this section centered around 40 Ma that is overlapping with the time interval of the (marine) MECO warming event (at 40.0 Ma; Bohaty et al., 2009). 5.3. North American MECO temperature conditions Based on U–Pb geochronology as well as magneto- and biostratigraphy we feel confident that the Dell Beds section analyzed here covers the MECO interval. Given the uncertainty of our U–Pb geochronology, however, it is not possible to identify potential leads and lags between marine and terrestrial climate signals (Fig. 4). A comparison of pedogenic carbonate 47 data in Montana with marine MECO temperature records shows that North American middle Eocene 47 temperatures display typical MECO characteristics (Fig. 4): Benthic foraminifera δ 18 O, Uk37 and TEX86 proxy records and continental 47 temperatures exhibit a transient temperature increase towards MECO peak conditions, followed by a large temperature decline at the termination of the MECO event. The magnitude of terrestrial temperature change (9–11 ◦ C), however, is strongly amplified when compared to the deep-sea records

Fig. 4. Comparison of temperature estimates and δ 18 O values of the Middle Eocene Climatic Optimum in marine (A, B; adapted from Sluijs et al., 2013) and terrestrial records (C, D; this study). (A) UK37 (blue) and TEX86 (purple) sea surface temperatures (SST) reflect 3-point running average (Bijl et al., 2010). (B) Benthic foraminifera δ 18 O data and corresponding ice-free temperatures from Southern Ocean ODP sites 738 (orange) and 748 (red), adapted from Bohaty et al. (2009). (C) Clumped isotope (47 ) paleotemperatures of the Dell Beds with 2σ standard error. The terrestrial warming event is characterized by a transient increase in temperatures followed by a profound temperature drop after peak warming, which is similar to the first-order characteristics of the marine δ 18 O record. Our dating results (white dots) place the temperature excursion slightly after the marine MECO, however within error U–Pb geochronology permits the Dell Beds to cover the same time interval as the marine MECO records (see text for details). (D) δ 18 O values of soil water in [h, VSMOW], calculated from carbonate δ 18 O values considering 47 temperatures and calcitewater oxygen isotope fractionation (Kim and O’Neil, 1997). (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

(3–6 ◦ C; Bijl et al., 2010; Bohaty et al., 2009), consistent with a higher continental temperature sensitivity. In the absence of evidence for another middle Eocene climate extreme of similar duration and magnitude in the 41–40 Ma time interval we have therefore chosen to correlate the timing of the marine and terrestrial records. The combined records document that a) western North American soil temperatures experienced a protracted large magnitude (+9 ◦ C) temperature increase followed by a rather rapid drop in temperatures (−11 ◦ C) and b) the MECO had a clear impact on mid-latitude terrestrial climate by strengthening the hydrological cycle (inducing low δ 18 O values of soil water and ultimately precipitation) in continental western North America. Compared to the younger post-MECO interval, low paleosol CaCO3 contents and low carbonate δ 18 O values during the MECO, argue against enhanced evaporation under elevated temperatures, but indicate rather semiarid to sub-humid conditions and ultimately enhanced water vapor transport into the continental interior and a strengthened hydrological cycle (e.g., Allen and Ingram, 2002) during MECO warming. A strengthened hydrological cycle during the MECO is consistent with paleoenvironmental data from marginal marine basin sites (e.g., Boscolo Galazzo et al., 2013; Spofforth et al., 2010; Witkowski et al., 2012). Such continental-scale intensification of the hydrological cycle may hence be a typical characteristic of climate warming during hyperthermals (e.g., PETM; Bowen et al., 2004; Huber and Sloan, 1999) and is consistent with elevated δ 18 O values immediately prior and after the MECO interval in central Montana (Figs. 2 and 4). The rapid post-MECO increase in pedogenic CaCO3 content further points to enhanced levels of aridity, which is in good agreement with (a) regional paleobotanical evidence indicating cooling and drying (climatically and orographi-

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cally induced) during the middle Eocene (e.g., DeVore and Pigg, 2010), and (b) post-MECO mid-latitude cooling and aridification in the northern hemisphere (Bosboom et al., 2014). 6. Conclusion Collectively, our data indicate that we identified the Middle Eocene Climatic Optimum in western North America and hence the first terrestrial isotopic evidence for MECO warming through integration of U–Pb dating of pedogenic carbonate, stable isotope, and clumped isotope records. 47 temperature data of pedogenic carbonates from the middle Eocene terrestrial Dell Beds (Montana, USA) reveal a hyperthermal event with a temperature excursion of +9 ◦ C during the MECO. The MECO in this high-elevation North American site is characterized by low δ 18 O values and reduced pedogenic carbonate concentrations in the terrestrial record. Large and rapid shifts of δ 18 O values and CaCO3 contents in concert with temperature changes require profound changes in the atmospheric circulation pattern in western North America. The MECO interval is marked by warm and semi-arid to subhumid conditions, whereas cooler and drier climatic conditions including aridification in northern hemisphere are likely for the post-MECO interval (e.g., Bosboom et al., 2014). We suggest that transient warming and subsequent global cooling at the end of MECO played a crucial role in modifying mid-latitude rainfall patterns in the hinterland of the North American Cordillera. This analysis highlights the result that western North America appears to get wetter during times of global warming (e.g., Retallack, 2009). Our observation is consistent with studies of the PETM that show an intensification of the hydrologic cycle during this warm period (e.g., Huber and Sloan, 1999; Bowen et al., 2004). It further resembles the warm Pliocene of western United States in which there may have been background El Nino-like conditions that caused western North America to become wetter (Winnick et al., 2013). These observations would have important implications with regard to the amount and distribution of precipitation in the northern hemisphere mid-latitudes during warming, and suggest that it may become wetter, at least in some areas (cf., Chou et al., 2013; Pound et al., 2014), rather than drier in the future (cf., Sherwood and Fu, 2014). Acknowledgements AM and KM acknowledge support through the LOEWE funding program (Landes-Offensive zur Entwicklung wissenschaftlichökonomischer Exzellenz) of Hesse’s Ministry of Higher Education, Research, and the Arts. CPC, SAG and AM acknowledge support through NSF EAR 1019648 and JF through DFG FI-948/4-1. We thank two anonymous journal reviewers for valuable comments helping to improve this manuscript. We further thank J. Caves (Stanford) and T. Schwartz (Stanford) for valuable field support and S. Hofmann (Frankfurt) for technical assistance. Appendix A. Supplementary material Supplementary material related to this article can be found online at http://dx.doi.org/10.1016/j.epsl.2016.05.053. References Allen, M.R., Ingram, W.J., 2002. Constraints on future changes in climate and the hydrologic cycle. Nature 419, 224–232. http://dx.doi.org/10.1038/nature01092. Bijl, P.K., Houben, A.J.P., Schouten, S., Bohaty, S.M., Sluijs, A., Reichart, G.-J., Sinninghe Damsté, J.S., Brinkhuis, H., 2010. Transient middle Eocene atmospheric CO2 and temperature variations. Science 330, 819–821. http://dx.doi.org/10.1126/ science.1193654.

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