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Chemostratigraphic implications of spatial variation in the Paleocene-Eocene Thermal Maximum carbon isotope excursion, SE Bighorn Basin, Wyoming

Corresponding Author

Allison A. Baczynski

Department of Earth and Planetary Sciences, Northwestern University, Technological Institute, 2145 Sheridan Road,, Evanston, Illinois,, 60208 USA

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Francesca A. McInerney,

Francesca A. McInerney

Department of Earth and Planetary Sciences, Northwestern University, Evanston, Illinois, USA

Now at Sprigg Geobiology Centre, Environment Institute and School of Earth and Environmental Sciences, University of Adelaide, Adelaide, SA, Australia

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Scott L. Wing,

Scott L. Wing

Department of Paleobiology, Smithsonian Institution, Washington, D.C., USA

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Mary J. Kraus,

Mary J. Kraus

Department of Geological Sciences, University of Colorado Boulder, Boulder, Colorado, USA

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Jonathan I. Bloch,

Jonathan I. Bloch

Florida Museum of Natural History, University of Florida, Gainesville, Florida, USA

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Doug M. Boyer,

Doug M. Boyer

Department of Evolutionary Anthropology, Duke University, Durham, North Carolina, USA

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Ross Secord,

Ross Secord

Florida Museum of Natural History, University of Florida, Gainesville, Florida, USA

Now at Department of Earth and Atmospheric Sciences, University of Nebraska, Lincoln, Nebraska, USA

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Paul E. Morse,

Paul E. Morse

Florida Museum of Natural History and Department of Anthropology, University of Florida, Gainesville, Florida, USA

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Henry C. Fricke,

Henry C. Fricke

Department of Geology, Colorado College, Colorado Springs, Colorado, USA

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Allison A. Baczynski,

Corresponding Author

Allison A. Baczynski

Department of Earth and Planetary Sciences, Northwestern University, Technological Institute, 2145 Sheridan Road,, Evanston, Illinois,, 60208 USA

([email protected])Search for more papers by this author

Francesca A. McInerney,

Francesca A. McInerney

Department of Earth and Planetary Sciences, Northwestern University, Evanston, Illinois, USA

Now at Sprigg Geobiology Centre, Environment Institute and School of Earth and Environmental Sciences, University of Adelaide, Adelaide, SA, Australia

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Scott L. Wing,

Scott L. Wing

Department of Paleobiology, Smithsonian Institution, Washington, D.C., USA

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Mary J. Kraus,

Mary J. Kraus

Department of Geological Sciences, University of Colorado Boulder, Boulder, Colorado, USA

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Jonathan I. Bloch,

Jonathan I. Bloch

Florida Museum of Natural History, University of Florida, Gainesville, Florida, USA

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Doug M. Boyer,

Doug M. Boyer

Department of Evolutionary Anthropology, Duke University, Durham, North Carolina, USA

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Ross Secord,

Ross Secord

Florida Museum of Natural History, University of Florida, Gainesville, Florida, USA

Now at Department of Earth and Atmospheric Sciences, University of Nebraska, Lincoln, Nebraska, USA

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Paul E. Morse,

Paul E. Morse

Florida Museum of Natural History and Department of Anthropology, University of Florida, Gainesville, Florida, USA

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Henry C. Fricke,

Henry C. Fricke

Department of Geology, Colorado College, Colorado Springs, Colorado, USA

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First published: 06 September 2013

https://doi.org/10.1002/ggge.20265

Citations: 35

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Abstract

[1] The Paleocene-Eocene Thermal Maximum (PETM) is marked by a prominent negative carbon isotope excursion (CIE) of 3–5‰ that has a characteristic rapid onset, stable body, and recovery to near pre-CIE isotopic composition. Although the CIE is the major criterion for global correlation of the Paleocene-Eocene boundary, spatial variations in the position and shape of the CIE have not been systematically evaluated. We measured carbon isotope ratios of bulk organic matter (δ¹³C_org) and pedogenic carbonate (δ¹³C_carb) at six PETM sections across a 16 km transect in the SE Bighorn Basin, Wyoming. Bed tracing and high-resolution floral and faunal biostratigraphy allowed correlation of the sections independent of chemostratigraphy. The onset of the CIE in bulk organic matter at all six sections occurs within a single laterally extensive geosol. The magnitude of the CIE varies from 2.1 to 3.8‰. The absolute and relative stratigraphic thickness of the body of the CIE in bulk organic matter varies significantly across the field area and underrepresents the thickness of the PETM body by 30%–80%. The variations cannot be explained by basinal position and instead suggest that δ¹³C_org values were influenced by local factors such as reworking of older carbon. The stratigraphic thickness and shape of the CIE have been used to correlate sections, estimate timing of biotic and climatic changes relative to the presumed carbon isotope composition of the atmosphere, and calculate rates of environmental and biotic change. Localized controls on δ¹³C_org values place these inferences in question by influencing the apparent shape and duration of the CIE.

Key Points

PETM bulk soil organic matter carbon isotope records from 6 sites across 16 km
Initial carbon isotope shift occurs within same laterally extensive geosol
Shape of CIE highly variable and PETM thickness underestimated by 30-80%

1. Introduction

[2] The Paleocene-Eocene epoch boundary at ∼56 Ma is marked by a significant perturbation to the global carbon cycle. The release of thousands of gigatons of isotopically light carbon to the ocean-atmosphere system resulted in an abrupt, transient episode of extreme global warming known as the Paleocene-Eocene Thermal Maximum (PETM) [Kennett and Stott, 1991; McInerney and Wing, 2011; Sluijs et al., 2007; Zachos et al., 1993, 2001]. Mean annual global temperature increased by ∼5–8°C during the PETM [Kennett and Stott, 1991; Secord et al., 2010; Sluijs et al., 2006; Tripati and Elderfield, 2005; Weijers et al., 2007; Wing et al., 2005; Zachos et al., 2001, 2003, 2005], and in the North American midcontinent the event is concurrent with major changes in faunas [Gingerich, 1989, 2003; Koch et al., 1992; Secord et al., 2012] and floras [Wing and Currano, 2013; Wing et al., 2005, 2009]. In the marine realm, the PETM carbon cycle perturbation and associated warming are marked by the dissolution of deep marine carbonate due to ocean acidification, the extinction of benthic foraminifers, and an increase in the abundance and geographic range of the dinoflagellate, Apectodinium [Crouch et al., 2001, 2003; Sluijs et al., 2005, 2006; Thomas, 1989, 1998; Zachos et al., 2005]. The carbon cycle perturbation is recorded in both terrestrial and marine carbonate and organic carbon, plant and algal lipids, and mammalian tooth enamel as a prominent negative carbon isotope excursion (CIE) [e.g., Bains et al., 2003; Bowen et al., 2001; Domingo et al., 2009; Dupuis et al., 2003; Fricke et al., 1998; Kennett and Stott, 1991; Koch et al., 1995, 2003; Magioncalda et al., 2004; McInerney and Wing, 2011; Pagani et al., 2006; Secord et al., 2010, 2012; Sluijs et al., 2006; Smith et al., 2007; Thomas and Shackleton, 1996; Zachos et al., 2005].

[3] The CIE has a characteristic magnitude (∼3–5‰) and shape as defined by the duration of the onset (≤20 kyr), body (∼100–120 kyr), and recovery (∼70–80 kyr) [Cui et al., 2011; Murphy et al., 2010; Rohl et al., 2007]. The exact magnitude and shape of the CIE in atmospheric CO₂ is unknown, and both the magnitude and shape vary according to the environment and material in which the CIE is recorded, diagenesis, and shifts in depositional rates within local sections [McInerney and Wing, 2011]. It is important to understand the factors that influence the shape and magnitude of the excursion in different depositional environments because the CIE is used to correlate continental, transitional, and marine PETM records, develop age models for local sections, and constrain the mechanism and source(s) of carbon release.

[4] The terrestrial record of the CIE in bulk organic matter is predominantly preserved in ancient floodplain depositional environments [Domingo et al., 2009; Magioncalda et al., 2004; Manners et al., 2013; Schmitz and Pujalte, 2007; Thiry et al., 2006; Wing et al., 2005; Yans et al., 2006]. Observations of contemporary fluvial systems reveal significant spatial heterogeneity in carbon preservation across the floodplain [Cierjacks et al., 2010; Hill and Cardaci, 2004; Hoffmann et al., 2009; Samaritani et al., 2011; Shrestha, 2011] and the mixing of carbon from multiple sources (modern and aged plant-derived organic matter and ancient rock carbon) [Blair et al., 2010; Brackley et al., 2010; Clark et al., 2013]. Floodplain environments in the geologic past may have had similar spatial heterogeneity and mixing of carbon from multiple sources that would have affected the isotopic record of bulk organic matter. Preserving an unparalleled record of PETM sedimentary rocks and fossils from ancient floodplain environments, the Bighorn Basin in north central Wyoming provides a unique opportunity to examine how the floodplain environment shapes the expression of the globally recognized CIE associated with the PETM.

[5] In this study we examine the spatial variability in the expression of the CIE in soil organic matter, a material widely used to identify the PETM, at six sites across a 16 km transect in the SE Bighorn Basin, Wyoming (Figure 1). The six sections (HW 16, CAB 10, CAB 3, Big Red Spit, Pyramid Point, and North Butte) have been correlated by physically tracing the surface exposures of marker beds (Figure 2), and correlations are further supported by floral and faunal biostratigraphy identifying the position of the PETM (Figure 3). We also measured the carbon isotope ratios of soil carbonates (δ¹³C_carb) where present. We compare the carbon isotope ratios of bulk organic matter (δ¹³C_org) from these sections with each other and with those from three other Bighorn Basin sections: Polecat Bench [Magioncalda et al., 2004], Sand Creek Divide [Rose et al., 2012], and Honeycombs [Yans et al., 2006]. We consider three main questions: (1) does the onset of the CIE have the same stratigraphic position, magnitude, and thickness across the field area; (2) does the body of the CIE have a uniform thickness and shape across the field area; and (3) what are the implications of this study for chemostratigraphy and our understanding of ecological, evolutionary, and carbon cycle changes at the PETM? The answers to these questions have implications for chemostratigraphic correlation, characterization of the shape of the CIE in relation to time, and understanding how local factors can influence the interpretation of a global event.

Details are in the caption following the image — **Figure 1**
Open in figure viewer PowerPoint

Map showing exposures of the Willwood Formation (shaded) in the Bighorn Basin. Index map in upper right corner shows location of the Bighorn Basin in northern Wyoming. Stars on map and field area inset map mark the locations of PETM sections at Polecat Bench, Sand Creek Divide, HW 16, CAB 10, CAB 3, Big Red Spit, Pyramid Point, Honeycombs, and North Butte. Base map and location of basin axis and Tensleep Fault modified after *Finn et al*. [2010].

2. Study Area

[6] The Bighorn Basin in northern Wyoming preserves a thick sequence of upper Paleocene-lower Eocene continental strata and has been the site of extensive paleontological, sedimentological, and geochemical research on the PETM [Bowen et al., 2001; Clyde and Gingerich, 1998; Fricke et al., 1998; Gingerich, 2003; Koch et al., 1992, 1995; Kraus and Riggins, 2007; Kraus et al., 2013; Magioncalda et al., 2004; Rose et al., 2012; Secord et al., 2010, 2012; Smith et al., 2007; Strait, 2001; Wing et al., 2005, 2009; Yans et al., 2006]. Abundant exposures and well-understood, high-resolution biostratigraphy make the Bighorn Basin an exceptional setting for examining the influence of local biogeochemical, sedimentological, and environmental factors on the preservation and shape of the CIE.

[7] In the SE Bighorn Basin (Figure 1), the PETM coincides approximately with the transition from the Fort Union Formation (Fm) to the Willwood Fm [Kraus et al., 2013; Rose et al., 2012; Secord et al., 2012; Wing et al., 2009]. Both formations are composed largely of mudstone, fine-grained sandstone, and carbonaceous shale. These lithologies are indicative of low-lying floodplain, fluvial channel or channel margin, and abandoned channel depositional environments [Kraus et al., 2013; Wing, 1998; Wing et al., 2005]. Paleosols that formed on fine-grained floodplain deposits are typically drab gray, green, and yellow in the Fort Union Fm but variegated red, purple, and orange in the Willwood Fm as the result of iron oxidation under well-drained conditions [Kraus and Riggins, 2007; Kraus et al., 2013].

[8] Four informal lithostratigraphic sequences have been recognized in the Paleocene-Eocene rocks of the SE Bighorn Basin [Kraus et al., 2013; Secord et al., 2012; Wing et al., 2009]: (1) a “transitional sequence” comprising the uppermost 5–15 m of the Fort Union Fm with localized, purple paleosols; (2) a “lower red sequence” of carbonate-bearing paleosols in the basal 5–10 m of the Willwood Fm that coincides with the onset and early part of the PETM as determined by biostratigraphy (Wa-0 mammalian biozone) and the marked negative shift in δ¹³C values of bulk organic carbon, n-alkanes [Smith et al., 2007], and mammalian tooth enamel [Secord et al., 2012]; (3) a “yellow-gray (or cut-and-fill) sequence” ∼15 to 20 m thick with weakly developed, yellow-brown paleosols containing large carbonate nodules; and (4) a 15 to 20 m thick “big red sequence” comprised of thick, well-developed, purple and purple-red paleosols that encompasses the uppermost part of the CIE body, all of the recovery phase, and the earliest post-PETM Eocene.

3. Methods

3.1. Bed Tracing

[9] Distinctive and laterally continuous marker beds were tracked on foot across the study area from U.S. Route 16 (HW 16 section) to North Butte (Figure 2). Coordinates were recorded frequently along bed traces with a differentially corrected GPS (DGPS: WAAS enabled Trimble ProXRS 2003 model and Trimble Geo-XH 2005 model). Although exposure is not completely continuous, gaps were small enough that bed identity could be confirmed by visual similarity of bed features. The total data set includes 5109 bed points. DGPS coordinates from the bed traces were plotted in Google Earth, and higher fidelity traces of two prominent and laterally extensive marker beds were drawn on high-resolution color satellite images in Google Earth using the coarsely spaced DGPS coordinates as ground truth (Figure 2). Figure 2 shows only a fraction of the tracing data available for assessing bed identity because many other superjacent and subjacent beds were also traced. Short gaps in beds of interest due to vegetative cover or terrace fill did not necessarily produce uncertainty on bed identity because traceable beds up or down section could be followed until the bed of interest reappeared. DGPS points taken on isolated exposures are not usually represented in Figure 2, as the satellite imagery does not allow extension of those points into linear traces. Additionally, beds in areas of excellent exposure far from measured stratigraphic sections were traced visually, but not always marked by DGPS points, because the exposure was so obvious. DGPS geographic accuracy and precision are submeter. Elevations determined by DGPS are accurate to within 2.5 m relative to the global geoid and also have submeter precision [Secord et al., 2012, Supplemental Figure 1]. Using DGPS data to record the difference in elevation between fossil localities and laterally continuous marker beds throughout the field area, we were able to project the relative stratigraphic position of fossil localities onto the nearest measured stratigraphic section (Figure 3).

3.2. Sample Collection

[10] Stratigraphic sections were measured with a Jacob's staff and sighting level after digging ∼1 m wide trenches down to fresh, unweathered rock (Figures S1–S6¹). GPS coordinates were taken at each trench (Table S1). Paleosol field units were designated, and numerous hand samples from each stratigraphic unit were described as outlined in Kraus et al. [2013]. Fresh rock samples were collected at high stratigraphic resolution (approximately every 0.25–0.5 m) using a rock hammer and placed in cloth bags for transport and storage. In total, 458 samples were collected from the field area (n = 127 at HW 16, n = 102 at CAB 10, n = 24 at CAB 3, n = 76 at Big Red Spit, n = 14 at Pyramid Point, and n = 115 at North Butte). In situ carbonate nodules were extracted from the rock matrix and collected by stratigraphic level when present. Thirty-two total carbonate nodules were collected from the field area (n = 12 at HW 16, n = 11 at CAB 10, n = 6 at Big Red Spit, and n = 3 at North Butte).

3.3. Bulk Organic Matter Sample Preparation and δ¹³C_org Analysis

[11] Samples were ground to a fine powder with solvent-rinsed mortar and pestle. The powdered sample (1.5 g) was placed in a centrifuge tube; 30 mL of 0.5 N hydrochloric acid was added and mixed using a vortex mixer. The acidified samples sat for 30 min to 1 h to remove carbonate, and soluble salts were removed by diluting with deionized water, centrifuging, and decanting until a neutral solution, measured by decanting solution onto pH paper, was achieved. Residues were freeze-dried and powdered again; 5–50 mg of sample (mass of sample calculated based on total organic carbon (TOC) values) was weighed into a tin boat for isotopic analysis. Carbon isotope ratios and weight percent TOC were measured in duplicate using a Costech ECS 4010 combustion elemental analyzer coupled to a Thermo Delta V Plus isotope ratio mass spectrometer (IRMS). Carbon isotope values are reported in delta notation normalized to the international standard Vienna Pee Dee belemnite (VPDB) using acetanilide 1 (Indiana University) and International Atomic Energy Agency (United Nations), Vienna, Austria (IAEA) reference standards 600 and CH-3. Isotope ratios and standard deviations of replicate analyses of standards and samples are reported in Table S1. Replicate measurements of standards indicate a measurement precision of ±0.1‰ (n = 557; 1σ). The mean standard deviation on replicate sample analyses is 0.2‰ (n = 449).

3.4. Pedogenic Carbonate Sample Preparation and δ¹³C_carb Analysis

[12] Microsamples of primary micritic calcite from pedogenic carbonate nodules were obtained for stable isotope ratio analysis by drilling on a clean, polished face using a handheld Dremel tool and avoiding secondary diagenetic spar; 250–500 µg of sample was weighed into glass vials and reacted with 100% orthophosphoric acid. The carbon dioxide released by this reaction was purified on a Thermo GasBench II and analyzed on a Thermo Delta V Plus IRMS. Carbon isotope values were measured in duplicate and are reported in delta notation standardized to the VPDB scale using IAEA reference standards NBS-18 and NBS-19. Isotope ratios and standard deviations of replicate analyses of standards and samples are reported in Table S2. Replicate measurements of standards indicate a measurement precision of ±0.1‰ (n = 16; 1σ). The mean standard deviation on replicate sample analyses is 0.3‰ (n = 32).

4. Results

4.1. Lithostratigraphy

[13] Many marker beds with varying degrees of lateral extent were traced in order to correlate adjacent sections and develop a strong lithostratigraphic framework. Most beds are too laterally variable or discontinuous to be traced and/or recognized at every site. However, two prominent, laterally extensive geosols were identified, traced throughout the field area, and used for lithostratigraphic correlation of the six sections (Figure 2). A geosol is defined as a “… laterally traceable, mappable, geologic weathering profile …” found at a consistent stratigraphic position [North American Commission on Stratigraphic Nomenclature, 2005, p. 1560]. We designate the lower prominent geosol as the onset geosol, named for its association with the onset of the CIE, and the upper one as the Big Red 1 (BR1) geosol, named for its position at the base of the big red sequence. We have not formally defined or named these geosols.

[14] Throughout the field area, the onset geosol is unique among uppermost Fort Union Fm and lowermost Willwood Fm paleosols in bearing abundant black, organic fragments. In many locations, it also contains the lowest pedogenic carbonate nodules. The onset geosol varies in color from red to gray. Near U.S. Route 16 (HW 16 section), the onset geosol is gray and lies immediately below the lowest continuous red paleosol (Figure 3). This red paleosol is overlain by another gray bed, forming a gray-red-gray sequence. The gray-red-gray sequence was traced southeast from HW 16 through ∼9 km of badland exposure to the north side of the Cabin Fork drainage (CAB 10 section; Figure 2). In the vicinity of CAB 10, the gray-red-gray sequence is not apparent in outcrop, most likely because it is condensed. The red paleosol in the gray-red-gray sequence is expressed as a purple mudstone in the Cabin Fork sections, perhaps because of a change to more poorly drained soil conditions in this area. The stratigraphically lowest carbonate nodules occur a few meters above the base of this purple paleosol at CAB 10. The onset geosol was then traced ∼7 km east to the vicinity of North Butte (Figure 2). East of the CAB sections, the onset geosol becomes almost entirely gray in color but still bears black organic fragments and the stratigraphically lowest carbonate nodules. West of CAB 3 and north of the east fork of Nowater Creek, the onset geosol is developed to a moderate degree at the top of a thick sandstone lens, which in turn, is underlain by a ferruginous sandstone containing abundant Paleocene fossils.

[15] The BR1 geosol has also been traced across the study area from HW 16 to North Butte (Figure 2). BR1 is the lowest red bed in a sequence of purple-red beds. It is typically brick red in weathered exposures, is just above the highest pedogenic carbonate nodules, and is usually >1 m thick. Paleosols above the BR1 geosol in the big red sequence are more purple in color and lack carbonate nodules.

4.2. Biostratigraphy

4.2.1. Vertebrate Biostratigraphy

[16] Four mammalian interval biozones are associated with the PETM. From oldest to youngest these zones and their abbreviations (in parentheses) are the Copecion (Cf-3), Meniscotherium priscum (Wa-M), Sifrhippus sandrae (Wa-0), and Cardiolophus radinskyi (Wa-1) interval zones [Gingerich, 2001; Secord, 2008; Secord et al., 2006]. The base of each biozone is recognized by the first occurrence of the taxon for which it is named, and the top is defined by the first appearance of the taxon defining the succeeding zone. The Copecion zone [Secord, 2008; Secord et al., 2006], the third and final biozone of the Clarkforkian land-mammal age, is latest Paleocene to earliest Eocene in age. The other three biozones are subdivisions of the Wasatchian land-mammal age, which is earliest Eocene.

[17] Meniscotherium priscum, the index fossil for zone Wa-M, occurs in the lowest 4–5 m of the CIE in the northern Bighorn Basin [Bowen et al., 2001; Gingerich, 2001, 2003; Gingerich and Smith, 2006] but is only tentatively documented in the SE Bighorn Basin based on a single specimen found in the Honeycombs area [Strait, 2003; Yans et al., 2006]. After nine field seasons and the collection of >9000 specimens, no additional Meniscotherium specimens have been found in the SE Bighorn Basin. The lack of Meniscotherium could be related to sampling or paleoecological biases, or the short interval of time during which it was present in the area might be represented by a hiatus or unfossiliferous rock in some sections.

[18] The Meniscotherium priscum zone (Wa-M) is succeeded by the Sifrhippus sandrae zone (Wa-0). The Wa-0 biozone is traditionally referred to as the Hyracotherium sandrae zone, but “Hyracotherium” does not appear to be a valid genus when applied to North American equids [Froehlich, 2002]. Thus, we follow the most recent revision of the early Equidae [Froehlich, 2002] and refer to the Wa-0 biozone as the Sifrhippus sandrae zone. In the SE Bighorn Basin, the base of the Wa-0 biozone corresponds to the onset of the CIE and ends with the first appearance of Cardiolophus radinskyi, the index fossil for biozone Wa-1 (Figure 3).

4.2.2. Floral Biostratigraphy

[19] Two biostratigraphic zones have been described previously for late Paleocene and early Eocene megafloras in the Bighorn Basin: the late Paleocene Persites-Cornus Assemblage Zone [Hickey, 1980] and the early Eocene Metasequoia-Cnemidaria Concurrent Range Zone [Wing et al., 1991]. The Persites-Cornus Zone corresponds in part to the Clarkforkian land-mammal age, and the Metasequoia-Cnemidaria Zone corresponds to mammalian biozones Wa-1 to Wa-4 [Wing et al., 1991]. Persites-Cornus Zone floras commonly have the conifers Glyptostrobus and Metasequoia; typical angiosperms include Davidia antiqua, Beringiophyllum cupanioides, Browniea serrata, Celtis asperum, Porosia verrucosa, and the sycamores Platanus raynoldsii and Macginitiea gracilis.

[20] No formal biozone has been named for PETM megafloras, but they are almost entirely distinct from those that precede and follow [McInerney and Wing, 2011; Wing and Currano, 2013; Wing et al., 2005]. Only two taxa present in the Persites-Cornus Zone are found in the PETM assemblages, which are dominated in diversity and numbers by species in the families Fabaceae (beans) and Arecaceae (palms) [Wing and Currano, 2013; Wing et al., 2005, 2009]. The floras dominated by beans and palms occur only in faunal biozone Wa-0 (Figure 3). One Eocene index taxon, the floating fern Salvinia preauriculata, has its first appearance in the body of the CIE. Conifers and Ginkgo, which are otherwise common throughout the late Paleocene and early Eocene, are absent from the PETM.

[21] Conifers and Ginkgo reappear in the post-PETM Metasequoia-Cnemidaria Zone and are common throughout. Many of the angiosperms present in the late Paleocene also reappear after, or in the recovery phase, of the CIE. Davidia antiqua, Beringiophyllum cupanioides, Browniea serrata, Celtis asperum, and Porosia verrucosa are, however, absent from the Metasequoia-Cnemidaria Zone and may not have returned to the Bighorn Basin following the PETM. Megafossil species with first appearances low in the Metasequoia-Cnemidaria Zone include Alnus sp., Platycarya americana, and the ferns Cnemidaria magna and Lygodium kaulfussi.

4.3. Integration of Lithostratigraphy and Biostratigraphy

[22] Diverse upper Paleocene floras and a modest assemblage of late Clarkforkian (Cf-3, uppermost Paleocene) mammals have been recovered below the base of the onset geosol (Figure 3) [Secord et al., 2012; Wing et al., 2005, 2009]. The youngest Persites-Cornus Zone floras occur ∼8 m below the base of the onset geosol in the HW 16 section, approximately in the middle of the “transitional sequence,” and the youngest Clarkforkian mammals are found less than 4 m below the base of the onset geosol at Big Red Spit and Pyramid Point (Figure 3) [Wing et al., 2009]. Wa-0 faunas and/or PETM floras with abundant beans and palms appear less than 0.5 m above the base of the onset geosol at CAB 10, CAB 3, Big Red Spit, and Pyramid Point and are present throughout the lower red and yellow-gray (or cut-and-fill) sequences. The BR1 geosol occurs in the upper part of the Wa-0 biozone (Figure 3) [Secord et al., 2012; Wing et al., 2005, 2009]. PETM floras and Wa-0 mammals are found within and up to a meter above the second purple-red paleosol in the big red sequence (BR2; Figure 3). The lowest stratigraphic occurrence of a Metasequoia-Cnemidaria Zone (lower Eocene) flora is in a cut-and-fill deposit between BR2 and the third purple-red paleosol of the big red sequence (BR3) at HW 16 and Big Red Spit (Figure 3). Cardiolophus radinskyi, the indicator of biozone Wa-1, also first occurs between BR2 and BR3 (CAB 10) and within a few meters above a shift back to larger body size in the Sifrhippus sandrae lineage [Secord et al., 2012].

4.4. Carbon Isotope Ratios

[23] The δ¹³C_org records for the six PETM sections display broadly similar stratigraphic trends (Figure 3). In stratigraphic order these are as follows: (1) an ∼20 m thick interval in the uppermost Fort Union Fm with δ¹³C_org values typically greater than or equal to −25‰; (2) a 1 to 2 m thick interval in the onset geosol in the basal Willwood Fm in which there is a 2–4‰ negative shift in δ¹³C_org values; (3) an interval of variable stratigraphic thickness in which δ¹³C_org values are typically less than or equal to −26‰; (4) a return to less negative δ¹³C_org values of about −25‰; and (5) an interval in which δ¹³C_org values are generally greater than or equal to −25‰.

[24] We calculate the magnitude of the CIE onset as the difference between the stratigraphically highest δ¹³C_org value that precedes the onset geosol and the most negative δ¹³C_org value within the onset geosol. The thickness of the CIE onset is the stratigraphic separation of these points. The magnitude of the CIE onset ranges from 2.1‰ to 3.8‰, and the thickness of the CIE onset interval varies from 1.7 to 0.5 m (Table 1). The interval of rock with δ¹³C_org values consistently less than or equal to −26‰ ranges from ∼5 to ∼21 m (Table 1). The median of absolute differences between δ¹³C_org values of adjacent samples within the excursion is greater than 0.5‰ at each site (Table 1).

Table 1. CIE Onset Magnitude and Thickness, Thickness of CIE Body Based on δ¹³C_org Values, Biostratigraphic (Wa-0) PETM Body Thickness, δ¹³C_org CIE Body Incompleteness, Interval Between Onset and BR1 Geosols, and Median of Absolute Differences Between Adjacent δ¹³C_org and δ¹³C_carb Samples Within the CIE Body

Site Locality	Onset Magnitude (‰)	Onset Thickness (m)	δ¹³C_org CIE Body Thickness (m)	Biostratigraphic (Wa-0) PETM Body Thickness (m)	δ¹³C_org CIE Body Incompleteness (%)a	Interval Between Onset and BR1 Geosols (m)	Median Absolute Differences δ¹³C_org Within CIE Body (‰)	Median Absolute Differences δ¹³C_carb Within CIE Body (‰)
HW 16	3.6	1.6	21	33	36	30	0.67	0.46
CAB 10	3.8	1.7	5	28.5	82	23	0.58	0.23
CAB 3	3.1	1.7
Big Red Spit	2.8	1.6	6	27	78	20	0.52	0.32
Pyramid Point	2.1	1
North Butte	3.1	0.5	12	18	33	13	0.66	0.3

^a δ¹³C_org CIE body incompleteness calculated as 1 − δ¹³C_org CIE body thickness/Wa-0 thickness.

[25] New measurements of the δ¹³C values of pedogenic carbonate nodules from HW 16, Big Red Spit, and North Butte are presented in gray in Figure 3, along with previously published δ¹³C_carb data from CAB 10 [Wing et al., 2009]. Small arrows in Figure 3 denote first and last appearances of pedogenic carbonate. Carbonate nodules are present only within the PETM body in the SE Bighorn Basin, and all δ¹³C_carb values are between −12 and −16‰ (Figure 3). Although the δ¹³C_carb records do not capture the onset of the CIE, the δ¹³C_carb values we measured from within the PETM body are comparable to CIE body δ¹³C_carb values at other Bighorn Basin sites that do preserve the onset and body of the PETM [e.g., Bains et al., 2003; Bowen et al., 2001; Koch et al., 1995, 2003]. The Polecat Bench PETM section has the longest and highest resolution record of pedogenic carbonate δ¹³C, and sustained δ¹³C_carb values less than −12‰ are found only within the CIE (Figure 4) [Bowen et al., 2001]. The median of absolute differences between δ¹³C_carb values of adjacent samples within the excursion is less than 0.5‰ at each site (Table 1).

5. Discussion

5.1. Position, Magnitude, and Thickness of the CIE Onset

[26] Although it is often assumed that the onset of the CIE can be used to correlate different PETM sections, no previous study has demonstrated that this chemostratigraphic marker occurs in the same lithostratigraphic position across an ancient landscape. The lithostratigraphic and biostratigraphic framework we have established in the SE Bighorn Basin allows us to show that across 16 km, the onset of the CIE occurs within the laterally traceable bed we term the onset geosol. The onset geosol has been estimated to represent ∼24 kyr [Secord et al., 2012, LIRB], giving a maximum duration for the onset of the CIE that is consistent with other recent estimates (∼15–25 kyr) [Bowen et al., 2006; Charles et al., 2011; Cui et al., 2011].

[27] The magnitude of the CIE at the six sites ranges from 2.1‰ to 3.8‰ (Table 1). These spatial variations cannot be attributed to incomplete sampling of the CIE because sample resolution is high across the CIE onset (every ∼0.1–0.5 m) and includes multiple samples from within the onset geosol at all localities. Variations in CIE magnitude among sections probably reflect the mixing of carbon from different sources (with different carbon isotope values), or different degradational states of organic matter, rather than incomplete sampling.

[28] The thickness of the CIE onset interval changes abruptly across the study area (Table 1 and Figure 3). It averages 1.65 m (±0.06 standard deviation) to the west but is much thinner at Pyramid Point (1.0 m) and North Butte (0.5 m). Unlike the gradual thinning of the entire PETM sequence from NW to SE (Figure 3 and Table 1), which is consistent with slower long-term depositional rates in that direction [Bown, 1980; Clyde et al., 2007; Kraus, 1992], the abrupt thinning of the onset across small distances suggests local causes such as small-scale scouring or other floodplain processes.

5.2. Variability in the Thickness and Shape of the CIE Body

[29] The stratigraphic thickness of the interval with excursion δ¹³C_org values (δ¹³C_org < −26‰) varies widely among our sites and does not thin in parallel with the gradual NW to SE thinning of the interval between the onset and BR1 geosols (Figure 3 and Table 1). Rather than maintaining a consistent relationship to the traced geosols, the return to less negative δ¹³C_org values occurs at a different relative stratigraphic position in each section. The lack of consistency in the relative position of the tops of the excursion intervals in different sections implies that they are not contemporaneous and cannot all record the full duration of the PETM body.

[30] Biostratigraphic and isotopic records suggest that none of the bulk organic matter records presented here record the full duration of the body of the PETM. Fossil and isotopic evidence indicates that the PETM body persists from the base of the onset geosol to at least BR2. Wa-0 fossils, which occur only within the PETM body, are found from the onset geosol to ∼1 m above the top of BR2 (Figure 3). Stable carbon isotope measurements from both tooth enamel and leaf wax lipids [Secord et al., 2012; Smith et al., 2007] also indicate that the PETM body extends from the base of the onset geosol to at least BR2. Pedogenic carbonate nodules record excursion values (δ¹³C_carb < −12‰) to at least BR1, above which carbonate nodules are no longer found (Figure 3). In contrast, δ¹³C_org values return to preexcursion values ∼6–24 m below where tooth enamel and leaf wax lipid δ¹³C values return to preexcursion values and where the transition from Wa-0 to Wa-1 mammals occurs. Therefore, in our field area, the CIE observed in bulk organic matter underrepresents the thickness of the PETM body by 30–80% (Table 1).

[31] We propose that the large variation in the shape and thickness of the body of the CIE measured in bulk organic matter results from the mixing of different pools of carbon: autochthonous carbon fixed during the PETM and allochthonous carbon eroded from pre-PETM soils or rocks. The premature return to less negative δ¹³C_org values could be explained by an increase in the proportion of reworked (¹³C-enriched) allochthonous carbon, either due to an increase in the input of allochthonous carbon or to a decrease in the production and/or preservation of autochthonous carbon. The presence of older, refractory carbon has been observed in palynological preparations of PETM rocks from the study area [Wing et al., 2005, supplement]. Shark teeth and dinoflagellates found in terrestrial PETM sediments throughout the field area and Cretaceous zircons found in the CAB 10 section also suggest the introduction of allochthonous Mesozoic material.

[32] Studies conducted in modern fluvial systems have shown that refractory (rock) carbon is found in both riverine and floodplain sediments. In modern systems, the concentration of rock carbon remains relatively constant from sedimentary source rock to riverine and floodplain sediments at ∼0.25–0.4 weight percent [Blair et al., 2010; Clark et al., 2013] and can constitute up to 80% of the particulate organic carbon pool [Clark et al., 2013]. Paleocurrent studies of Willwood strata in the southeastern part of the basin show paleoflow from southeast to northwest [Neasham and Vondra, 1972; Newbury, 2011; Seeland, 1998] and point to the Southern Bighorn Mountains as the likely sediment source for the PETM deposits. The Southern Bighorn Mountains consist of a crystalline core flanked by sedimentary rocks of Paleozoic and Mesozoic age [e.g., Love and Christiansen, 1985]. Cretaceous marine units, which form a major part of those strata, contain considerable thicknesses of gray to black shales that are locally carbonaceous [Eicher, 1962; Hagen and Surdam, 1984; Keefer, 1972]. These carbonaceous Mesozoic rocks are enriched in ¹³C relative to PETM material. Therefore, an increase in the proportion of allochthonous carbon could explain the truncation of the CIE prior to the end of the PETM. Using the most negative PETM bulk organic matter excursion value (−28.5‰) and a mean Mesozoic δ¹³C_org value (−24.3‰, n = 10 formations; Table S3) in a simple binary mixing equation, we calculate that ∼40–80% allochthonous carbon would be required to produce the observed premature return to less negative δ¹³C_org values (see supporting information Table S4)¹. Similarly high proportions of rock carbon have been found in modern river systems, where rock carbon has been shown to constitute from 30 to 80% of the total particulate organic carbon pool [Blair et al., 2010; Clark et al., 2013]. Therefore, mixing of allochthonous rock carbon provides a realistic explanation for the truncation of the CIE. However, this mixing calculation represents a conservative lower-bound estimate of the proportion of allochthonous carbon because the true value of PETM bulk organic matter is unknown, and the estimated value of −28.5‰ likely includes some proportion of rock carbon. Additional constraints on the δ¹³C_org value of PETM bulk organic matter and the provenance of allochthonous carbon would improve this estimate.

[33] Pedogenic processes can also influence organic matter δ¹³C_org values. Depth profiles in modern soils show a logarithmic relationship between the fraction of soil organic carbon remaining (f_SOC) and δ¹³C_org value: as f_SOC decreases, δ¹³C_org values increase, i.e., become more enriched in ¹³C due to Rayleigh distillation [Wynn, 2007]. A similar logarithmic relationship between weight percent TOC and bulk organic matter δ¹³C values has been observed previously in Paleogene sediments from this field area [Wing et al., 2005] and is also found here (Figure S7). Although decomposition may have played an important role in altering the carbon isotope signature of organic matter in paleosols, attributing the logarithmic relationship in the Paleogene to the same mechanism identified in modern soils relies on an implicit assumption that weight percent TOC is a reliable proxy for f_SOC. For this to be true, the initial carbon content of all the sediments must have been the same, which is unlikely. Therefore, although soil organic matter degradation likely influenced δ¹³C_org values, these effects are difficult to identify in our samples.

[34] A similar pattern of ¹³C enrichment with decreasing weight percent TOC (Figure S7) could also be caused by the mixing of ¹³C-enriched allochthonous rock carbon with autochthonous carbon, if the proportion of allochthonous carbon increased with decreasing weight percent TOC. The allochthonous/autochthonous mixing ratio could vary systematically with TOC due to (1) sedimentary processes (e.g., avulsion deposits in the yellow-gray interval with low TOC and likely high allochthonous inputs versus well-developed paleosols with high autochthonous inputs and higher TOC), (2) preservational bias (e.g., preferential preservation of refractory rock carbon relative to PETM carbon in sediments with low TOC), or (3) varying production of organic carbon (e.g., higher relative input of PETM organic matter in higher TOC sediments such as those that preserve fossil leaves). If the allochthonous/autochthonous mixing ratio underlies the relationship between δ¹³C_org values and TOC, we would predict that PETM sediments should exhibit a larger ¹³C enrichment with decreasing TOC than pre- and post-PETM sediments. This is because pre- and post-PETM δ¹³C_org values are similar to Mesozoic δ¹³C_org values and would therefore be less affected by mixing than PETM δ¹³C_org values. The data in fact show greater ¹³C enrichment with decreasing weight percent TOC in PETM samples than in pre- and post-PETM samples (Figure S7), suggesting that mixing of allochthonous and autochthonous carbon could play an important role in shaping these carbon isotope records.

5.3. Implications for Chemostratigraphy and Our Understanding of the PETM

5.3.1. Proximal Correlations: Variability Across the Field Area

[35] This study clearly documents variation in the expression of the CIE in bulk organic matter across 16 km of outcrop. Differences in δ¹³C_org values within beds traced laterally between sites are on the order of 1–2‰, and the median of absolute differences between successive samples within the body of the CIE at each section is >0.5‰ (Table 1). Variability in δ¹³C_org values over short stratigraphic and lateral distances is commonly observed in Bighorn Basin Paleogene sections [Diefendorf et al., 2007; Smith et al., 2007; Wing et al., 2005, 2009] and can be attributed to the diversity of organic input to soils, differential preservation/degradation of organic components, and isotopic effects that occur during pedogenesis and microbial processing [Benner et al., 1987; Bowen and Beerling, 2004; Natelhoffer and Fry, 1988; Schmidt et al., 2011; Schweizer et al., 1999; Wynn, 2007].

[36] Local variation in input and preservation/degradation of organic material may obscure the identification and accurate correlation of even large isotopic shifts such as the recovery of the CIE. The observed truncation of the body of the CIE in bulk organic matter could potentially lead to false correlations, misidentification of the PETM recovery phase, and significant (30–80%) underestimations of the thickness of PETM strata. Smaller magnitude shifts in δ¹³C_org values are not likely to be reliable for correlation, even for stratigraphic sections that are relatively close together like ours. Only the onset of the CIE appears to be a shift large enough to overwhelm local heterogeneity, and without dense sampling, even the onset of the CIE can be difficult to identify.

5.3.2. Proximal Correlations: The Adjacent Honeycombs Section

[37] A vertebrate microsite near North Butte in the Honeycombs badland area of the SE Bighorn Basin (Figures 1 and 2) contains mammal species indicative of the Wa-0 faunal zone [Strait, 2001, 2003]. Yans et al. [2006] conducted a chemostratigraphic study in the vicinity of the microsite, producing a δ¹³C_org profile through the PETM using protocols similar to those we have used (Figure 5). They identified the onset of the CIE as a −3.7‰ excursion between the 7.8 and 8.5 m levels in their stratigraphic section and offered a possible correlation of four negative isotope landmarks between the Honeycombs and Polecat Bench δ¹³C_org records by aligning the base of the CIEs. The Honeycombs CIE identified by Yans et al. [2006] approximately matches the Polecat Bench δ¹³C_org excursion in both magnitude and stratigraphic thickness, and a Wa-0 fauna is preserved within the Honeycombs CIE [Strait, 2001, 2003].

[38] Comparison of the Honeycombs section to our sites in the SE Bighorn Basin, however, indicates a revised position for the base of the CIE in the Honeycombs section (Figure 5). A DGPS trace of marker geosols from the Honeycombs section to the sections we have measured shows that the “gray tracer” bed of Yans et al. [2006] is the same as our onset geosol, and the “thick red” of Yans et al. [2006] is the same as our BR1 geosol (Figure 2 and Table 2). These equivalences are most apparent in traces between the Honeycombs section and the North Butte section 1.4 km to the southeast (Figure 5). The base of the CIE proposed by Yans et al. [2006] is 13.6 m below the onset geosol that records the onset of the CIE in our six sections in the SE Bighorn Basin.

Table 2. Summary of Papers Referring to Beds That Are Stratigraphically Equivalent to the Onset and BR1 Geosols

Source	Stratigraphic Equivalent	Meter Level (Base)	Comments on Equivalence
Onset geosol
Wing et al. [2005]	(Not named but identified in stratigraphic column)	∼0 m	Earliest identification of CIE in the field area at this level
Wing et al. [2009]	First laterally persistent red paleosol containing CaCO₃ nodules	Multiple sectionsa	Base of first laterally persistent red is base of onset geosol within field area
Chester et al. [2010]	(Not named but identified in stratigraphic column)	∼0 m	Same stratigraphic column as Wing et al. [2005]
Secord et al. [2012]	LIRB (Lowest Intermittent Red Bed)	∼14 m	Base of LIRB is base of onset geosol in composite section tied to CAB 10
Kraus et al. [2013]	(Not named but identified in stratigraphic column)	∼20 m	Fine scale stratigraphic description of onset geosol at HW 16
Yans et al. [2006]	Gray tracer	∼22 m	Traced laterally between Honeycombs section and our field area using DGPS
Kraus and Riggins [2007]	1st Persistent Red	0 m	Traced through intermittent outcrops to onset geosol at HW 16
Rose et al. [2012]	Red 1	0 m	Traced through intermittent outcrops to onset geosol at HW 16; supported by biostratigraphy
BR1 geosol
Wing et al. [2005]	(Not named but identified in stratigraphic column)	∼35 m	Earliest identification of big red sequence in field area
Wing et al. [2009]	Lowest red paleosol in the big red sequence	Multiple sectionsa	Earliest identification of BR1 within field area
Chester et al. [2010]	(Not named but identified in stratigraphic column)	∼35 m	Same stratigraphic column as Wing et al. [2005]
Secord et al. [2012]	BR1 (Big Red 1)	∼42 m	BR1 within composite section tied to CAB 10
Kraus et al. [2013]	(Not named but identified in stratigraphic column)	∼50 m	Fine scale stratigraphic description of big red sequence at HW 16
Yans et al. [2006]	Thick red	∼37 m	Traced laterally between Honeycombs section and our field area using DGPS
Kraus and Riggins [2007]	Big Red Paleosol	∼18 m	Traced through intermittent outcrops to BR1 geosol at HW 16
Rose et al. [2012]	Big Red	∼23 m (S); ∼17 m (N)	Traced through intermittent outcrops to BR1 geosol at HW 16; supported by biostratigraphy

^a Candystripe Ridge and Picnic Area Hill [Wing et al., 2009] together comprise the HW 16 section presented in Figure 3. Big Red Spit [Wing et al., 2009] is the Pyramid Point section in Figure 3 of this study.

[39] This revised position for the base of the CIE in the Honeycombs section corresponds to an isotopic shift of only −1.2‰ (Figure 5), but it is not discordant with any of the biostratigraphic data presented by Yans et al. [2006]. The lowest Wa-0 mammal fossils are from the Castle Gardens and Halfway Hill S localities, ∼1–2 m above the base of gray tracer [Strait, 2001; Yans et al., 2006]. The stratigraphically highest Wa-0 fossils in the Honeycombs section are within the thick red paleosol, which is our BR1 geosol (Table 2), a paleosol that also contains Wa-0 mammals in our sections (Figures 3 and 5) [Secord et al., 2012; Wing et al., 2005]. Yans et al. [2006] reported a single specimen of Meniscotherium, the index fossil for biozone Wa-M, ∼2–5 m below gray tracer. At Polecat Bench in the northern Bighorn Basin, Meniscotherium coincides with minimum CIE values, below the first Wa-0 mammals but above the onset of the CIE [Bains et al., 2003; Bowen et al., 2001; Gingerich, 2003; Gingerich and Smith, 2006; Magioncalda et al., 2004]. The stratigraphic position of the Meniscotherium specimen from the Honeycombs section, ∼2–5 m below where we place the base of the CIE, can be explained by a small amount of downslope movement of the specimen after it was eroded from the outcrop [Wing et al., 2009; Wood et al., 2008]. The onset geosol in this area is immediately underlain by a laterally extensive sandstone unit ∼2–3 m thick, which forms a steeply weathered face. We have found many places where moderately abundant vertebrate fossils weathering out of the onset geosol have been washed down this steep face and come to rest on nearly horizontal surfaces just below [Chester et al., 2010; Wing et al., 2009]. The revised position of the onset of the CIE in the Honeycombs section casts doubt on the correlations of the smaller isotopic landmarks proposed between Honeycombs and Polecat Bench [Yans et al., 2006].

5.3.3. Basinwide Comparisons: Sand Creek Divide and Polecat Bench

[40] The sequence of paleosols at the Sand Creek Divide PETM locality [Rose et al., 2012], ∼19 km northwest of HW 16 (Figures 1 and 2), is similar to the paleosol sequence in the SE Bighorn Basin (Figure 4). We tentatively identify the onset geosol as the gray paleosol immediately underlying the lowest persistent red paleosol [Rose et al., 2012, “Red 1”], which marks the base of the Willwood Fm and contains the lowest pedogenic carbonate nodules (Figure 4 and Table 2) [Rose et al., 2012]. Red 1 is overlain by a prominent gray mudstone [Rose et al., 2012, Figures 2A and B], forming a gray-red-gray sequence similar to that in the HW 16 section (Figure 4). We also tentatively identify the BR1 geosol as the “Big Red” paleosol of Rose et al. [2012] (Table 2). Although the sequence of paleosols at the more distant Polecat Bench [Bowen et al., 2001; Gingerich, 1989, 2001; Magioncalda et al., 2004] in the northern Bighorn Basin (Figure 1) is different enough to impede lithostratigraphic correlation, biostratigraphic and chemostratigraphic evidence for the PETM enables us to compare and contrast the carbon isotope record from the Polecat Bench section with those from the SE Bighorn Basin (Figure 4).

[41] Both the Polecat Bench and Sand Creek Divide δ¹³C_org profiles have thicker CIE onset intervals relative to sites in the SE Bighorn Basin (Figure 4). In the northern Bighorn Basin, the δ¹³C_org record at Polecat Bench shifts −3.7‰ over 3.9 m [Magioncalda et al., 2004, Figure 2 and Table DR1]. At Sand Creek Divide, the coarse sampling resolution prevents a precise estimate of the thickness over which the onset occurs, but the shift spans at least several meters of section and is more gradual [Rose et al., 2012] than at sites to the southeast where the onset occurs over <2 m [this study; Yans et al., 2006]. Therefore, in general, the CIE onset interval recorded in bulk organic matter thins southeastward across the entire Bighorn Basin, from Polecat Bench in the north to the sections in this study and the Honeycombs site [Yans et al., 2006] at the southeastern margin of the basin (Figure 1). Likewise, the body of the PETM, defined by lithostratigraphy and biostratigraphy, also thins to the SE margin of the Bighorn Basin; from ∼40 m at Polecat Bench to ∼32 m at Sand Creek to <32 m in our field area in the SE Bighorn Basin (Figure 4).

[42] Comparison of the PETM sections provides insight into spatial differences in sedimentation rates across the Bighorn Basin. However, because of the landscape scale variations observed in the δ¹³C_org curves in this study, we caution against correlations among distant sites based on shifts in isotope ratios from bulk organic matter within the PETM.

5.3.4. Ecological, Evolutionary, and Carbon Cycle Changes at the PETM

[43] Marine carbonate records of the CIE often display truncation due to dissolution caused by the shoaling of the lysocline [McCarren et al., 2008; Zachos et al., 2005]. In contrast, terrestrial records are often considered to more completely preserve the CIE [McInerney and Wing, 2011]. The data presented here demonstrate that bulk organic matter δ¹³C profiles recorded in soils and floodplain sediments can also suffer from incomplete preservation of the CIE and can underrepresent both the thickness of the PETM body and the magnitude of the CIE. This influences mass balance approaches to modeling the amount of carbon that fueled the global warming at the Paleocene-Eocene boundary, which rely on the magnitude of the CIE [See McInerney and Wing, 2011 for review] and hypotheses regarding source(s), mechanism(s), and rate of carbon release. Variation in the magnitude and shape of the CIE across the field area demonstrates the importance of local processes and suggests that bulk organic matter cannot be reliably used to infer the shape or magnitude of the global CIE in areas where allochthonous organic matter could have been a significant contributor to the sedimentary organic carbon pool.

[44] Truncation of the CIE in terrestrial bulk organic matter also has implications for the reconstruction of ecological and evolutionary responses to the PETM. Studies that rely solely on bulk organic matter CIEs risk assuming that a truncated CIE body represents the full ∼100–120 kyr of the CIE body. The results presented here suggest that the CIE in bulk organic matter could underestimate the PETM body by 30–80%. In a worst case scenario, the observed CIE could represent only 20% of the actual PETM body, leading to the calculation of rates of migration, evolution, and sedimentation that are underestimated by a factor of 5. The estimation of rates of migration and evolution during the PETM has implications for anticipating future biotic responses to global warming.

6. Summary and Conclusions

[45] Six high-resolution bulk soil organic matter δ¹³C records from a 16 km transect across a suite of ancient floodplains in the SE Bighorn Basin, Wyoming, have been correlated using floral and faunal biostratigraphy and by tracing distinctive lithologic marker beds that characterize the PETM. This multiproxy approach has established a comprehensive PETM stratigraphic framework with which we can carefully examine how local processes influence and overprint the global PETM signal. The onset of the CIE occurs within the prominent and laterally extensive onset geosol at all six sites, confirming for the first time that this important chemostratigraphic marker occurs in the same lithostratigraphic position across an ancient floodplain.

[46] Carbon isotope ratios in bulk organic matter underestimate the thickness of the body of the PETM (as indicated by floral and faunal biostratigraphy and stable carbon isotope measurements from pedogenic carbonate nodules, tooth enamel, and leaf wax lipids) by ∼30%–80%. The premature recovery to less negative δ¹³C_org values in the SE Bighorn Basin does not occur at the same level in relation to biostratigraphic and lithostratigraphic markers at each site, nor does it agree with δ¹³C records from tooth enamel, leaf waxes, and carbonate nodules, which record a thicker CIE. As a result, δ¹³C_org records could lead to false correlations of the CIE recovery, underestimations of the thickness of PETM strata, and the miscalculation of rates of sedimentation, evolution, and migration.

[47] We suggest that the discrepancy between the δ¹³C_org records and the biostratigraphic and isotopic records from other proxies is caused by erosion of older rocks and redeposition of allochthonous carbon in this PETM floodplain environment, similar to that observed in modern fluvial systems. According to this hypothesis, the contribution of older allochthonous organic matter reduced the overall magnitude of the CIE in the δ¹³C_org record. Stratigraphic fluctuations in the ratio of allochthonous to autochthonous carbon are responsible for oscillations in the bulk organic matter δ¹³C record for each local section, and an increase in the ratio caused the body of the CIE to appear truncated. The minimum proportion of allochthonous carbon required to explain the truncation of the CIE ranges from 40 to 80%, which is in agreement with observations in modern floodplain settings. The variability in δ¹³C_org ratios within a single record and the differences in CIE shape across our field area demonstrate the importance of local environmental, biogeochemical, and/or sedimentological factors in determining δ¹³C_org values in terrestrial floodplain environments, particularly where abundant allochthonous organic matter has been eroded from older rocks and redeposited during the CIE. These strong local controls mean that a single δ¹³C_org curve may not be representative of even a relatively restricted field area, let alone an entire basin or continent. Multiple, high-resolution records across a field area with biostratigraphic and lithostratigraphic constraints can provide more robust identification of the PETM. Within such a stratigraphic framework, the nature of the CIE preserved in bulk organic matter can be examined to answer questions about carbon cycle dynamics among the atmosphere, terrestrial ecosystems, and soils during the PETM.

Acknowledgments

[48] We thank Alexa Socianu for help with carbonate isotope laboratory work, Stephen Chester for assisting with biostratigraphic and lithostratigraphic studies, Brady Foreman and Elizabeth Denis for sample collection, and the many students and volunteers who helped with fieldwork and laboratory work. We also thank two anonymous reviewers for their constructive comments. Funding was provided by NSF awards EAR-0720268 (F.A.M.), EAR-0717892 (S.L.W.), EAR-0718740 (M.J.K.), EAR-0719941 (J.I.B.), EAR-0640076 (J.I.B., R.S., and John Krigbaum), Initiative for Sustainability and Energy at Northwestern (F.A.M.), and Australian Research Council FT110100793 (F.A.M.). Vertebrate fossils were collected under Bureau of Land Management permits to J.I.B. (PA04-WY-113, PA10-WY-185). A portion of this manuscript was written when J.I.B. was supported as an Edward P. Bass Distinguished Visiting Environmental Scholar in the Yale Institute for Biospheric Studies (YIBS).

Supporting Information

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Citing Literature

Volume14, Issue10

October 2013

Pages 4133-4152

Filename	Description
ggge20265-sup-0001-suppinfo01.pdf659.2 KB	Supporting Information
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ggge20265-sup-0003-suppinfo03.xls232 KB	Supporting Information

Chemostratigraphic implications of spatial variation in the Paleocene-Eocene Thermal Maximum carbon isotope excursion, SE Bighorn Basin, Wyoming

Abstract

Key Points

1. Introduction

2. Study Area