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Geological Society of America Bulletin

doi: 10.1130/B30401.1 2012;124, no. 3-4;444-462Geological Society of America Bulletin

David L. Fox, James G. Honey, Robert A. Martin and Pablo Peláez-Campomanes

-dominated grasslands4isotopes and the evolution of Cthe Neogene in the southern Great Plains, southwest Kansas, USA: Carbon Pedogenic carbonate stable isotope record of environmental change during

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Pedogenic carbonate stable isotope record of environmental change during the Neogene in the southern Great Plains, southwest Kansas, USA:

Carbon isotopes and the evolution of C4-dominated grasslands

David L. Fox1,†, James G. Honey2, Robert A. Martin3, and Pablo Peláez-Campomanes4

1Department of Geology and Geophysics, University of Minnesota, Minneapolis, Minnesota 55455, USA2Department of Geological Sciences, University of Colorado, Boulder, Colorado 80309, USA3Department of Biological Sciences, Murray State University, Murray, Kentucky 42071, USA4Department of Paleobiology, National Museum of Natural History, Consejo Superior de Investigaciones Científi cas, Jose Guttierez Abascal 2, Madrid 28006, Spain

GSA Bulletin; March/April 2012; v. 124; no. 3/4; p. 444–462; doi: 10.1130/B30401.1; 6 fi gures; 2 tables; Data Repository item 2011262.

†Correspondence: [email protected], 612-618-3092.

ABSTRACT

Fossiliferous strata in the Meade Basin (southwest Kansas) preserve numerous superposed mammalian faunas and calcar-eous paleosols that range in age from the Clarendonian North American Land Mam-mal Age (NALMA; 12.0–9.0 Ma, early late Miocene) to the early Irvingtonian NALMA (ca. 2.5–ca. 1.0 Ma, early Pleistocene). Fau-nas from these sections document the evolu-tion of the small mammal community of the modern grassland ecosystem of the region, and the stable isotope composition of paleo-sol carbonates provides a means by which the environmental context of the evolution of the modern ecosystem may be documented. We used the stable carbon isotope composition (δ13C relative to Vienna Peedee belemnite [VPDB]) of 194 pedogenic carbonates from 19 measured sections to reconstruct the history of C4 grass abundance in the Meade Basin. Paleosol carbonate δ13C values refl ect the proportion of C3 (trees, shrubs, cool- climate grasses) and C4 (warm-climate grasses) plants that grew in an ancient soil and pro-vide a means with which to reconstruct past mammalian habitats. Paleosol carbonate δ13C values record a three-phase increase in the abundance of C4 biomass during the Neogene in the Meade Basin. Late Miocene sections have mean δ13C values of −7.6‰ ± 0.90‰ (Clarendonian) and −6.5‰ ± 0.31‰ (Hemphillian NALMA, 9.0–4.9 Ma), consis-tent with 17% and 26% C4 biomass, respec-tively. Miocene δ13C values from Meade are statistically identical to published δ13C values

for Miocene paleosol carbonates elsewhere in the southern Great Plains, supporting the widespread presence of ~20% C4 biomass on average in the region throughout the Mio-cene. The abundance of C4 biomass increased between the end of the Hemphillian sec-tion and the beginning of the early Blancan NALMA (5.0–3.0 Ma). Early and middle Blancan (3.0–2.5 Ma) carbonates have sta-tistically identical δ13C values (−4.9‰ ± 0.90‰ and −5.0‰ ± 1.10‰, respectively), suggesting a stable ecosystem during the early Pliocene, although high δ13C variability in densely sampled intervals suggests a high degree of landscape-scale variation in C4 abundance. The fi nal phase, geochronologi-cally controlled by two well- characterized ashes (Huckleberry Ridge, 2.10 Ma; Cerro Toledo B, 1.47–1.23 Ma) and magneto-stratigraphy, is a trend to higher δ13C values from the late Blancan to early Irvingtonian (ca. 2.5–ca 1.0 Ma) from −4‰ at the base of the section to ~1‰ at the top, correspond-ing to an increase from almost 50% to 65% C4 biomass. The abundance of C4 biomass fi rst reaches modern levels for the region (78% ± 10.9%) around the level of the Cerro Toledo B ash, indicating that a modern-like grassland ecosystem fi rst appeared in the region ca. 1.3 Ma, although δ13C values do not remain consistently high through the rest of the section.

INTRODUCTION

The origin of grassland ecosystems is a topic that has received considerable attention over the

last several decades, largely due to the evolution-ary, ecological, and biogeochemical importance of grassland ecosystems today and over their geological record, and the economic importance of grasslands to human societies today and in the past. Modern grasslands can be broadly divided into tropical, subtropical, and warm temperate grasslands that are dominated by grasses using the C

4 photosynthetic pathway and temperate

and high-altitude grasslands that are dominated by grasses using the C

3 photosynthetic pathway

(Edwards and Smith, 2010). Two conceptions about the origin of grasslands dominated by C

4

grasses have been persistent in the literature. First, many studies accept a simple chrono-logical scenario in which C

4 grasses begin to

become ecologically dominant globally around 8 Ma (latest Miocene), and the increase in abun-dance is relatively rapid and monotonic. Second, many studies accept a simple climatic scenario in which C

4 grasses become ecologically domi-

nant in response to increased aridity from the late Miocene to the present, based on the assumption that the C

4 photosynthetic pathway is at least in

part an adaptation to water stress.In this paper, we present a large data set of

carbon isotope values of paleosol carbonates from well-constrained sections in the Meade Basin, in the area of the town of Meade in south-western Kansas, USA, to document the timing of the evolution of the modern, C

4-dominated

grassland in the region over the last 9–12 m.y. We begin with a brief review of research into the origin of grass-dominated ecosystems that is organized around four research questions. This review is intended as an update to the comprehensive review of Jacobs et al. (1999)



Paleoclimate and C4 biomass in the Neogene of the Great Plains

Geological Society of America Bulletin, March/April 2012 445

and refl ects the increasingly interdisciplinary nature of research into the long-term history of grasslands. Then, before discussing the data and their implications, we give an overview of the stratigraphy of the Meade Basin and lay out the assumptions behind our interpretations of the carbon isotope data. In a companion paper in this issue (Fox et al., 2011), we pres-ent the associated oxygen isotope data from the paleosol carbonates to document the paleocli-matic context of the evolution of the modern, C

4-dominated grassland ecosystem in the

region. The data and discussions in these two papers address both of the common conceptions about the evolution of C

4 grasslands, at least

with regard to the southern Great Plains region of the central United States.

A BRIEF REVIEW OF THE ORIGIN OF GRASS-DOMINATED ECOSYSTEMS

Given the diversity of scientists across a broad range of disciplines that study the evolution of grassland ecosystems, it is important to rec-ognize that the topic can be broken down into several separate, but interrelated, research ques-tions, each of which implies auxiliary questions regarding causes and mechanisms of evolution-ary and ecological changes. Taken in roughly geochronological order from oldest to most recent, these questions are as follows. First, when did the ecologically and taxonomically dominant clades of grassland plants evolve, in particular the grasses of the clade Poaceae? This question has been addressed using molecular phylogenetic studies of Poaceae (Grass Phylog-eny Working Group [GPWG], 2001), pollen and plant macrofossil records (Linder, 1987; Muller et al., 1987; Crepet and Feldman, 1991), fossil phytoliths (Prasad et al., 2005), and molecular clock studies based on molecular phylogenetic hypotheses of various clades and calibration points from the fossil record (Bremer, 2002; Janssen and Bremer, 2004). Although fossils will generally underestimate the origination of a clade, and molecular clock studies are never ide-ally calibrated, these methods and data suggest that Poaceae had evolved by the Late Cretaceous. Poaceae could only have had a single evolution-ary origin, so the answer to this fi rst question is global in scope, although both phylogeny and the fossil record can be used to address the fi rst appearance of grasses in a given continent or region. The remaining questions regarding the evolution of grassland ecosystems are most appropriately addressed on continental or, better yet, regional spatial scales, given that they con-cern the histories of specifi c ecosystems.

The second question is: When did open, grass-dominated ecosystems fi rst appear on

each continent? Grass-dominated ecosystems vary in the density of the woody and shrubby component and range from tropical savan-nah woodlands with a relatively high density of woody and shrubby vegetation distributed uniformly or in patches to temperate treeless grasslands and steppes (Pratt et al., 1966; Cou-pland, 1992; Jacobs et al., 1999). In the Great Plains, the appearance of open, grass-dominated ecosystems has been examined with fossil phy-toliths (Strömberg, 2004), paleosols (Retallack, 1997, 2001; Terry, 2001), the evolution of feed-ing and locomotor specializations among terres-trial animals (Webb, 1977; Janis et al., 2004), and plant macrofossils (Thomasson, 1990). Different data sets suggest different answers to this question for the Great Plains, ranging from the late Eocene–early Oligocene (paleosols) to sometime during the late Oligocene to early Miocene (phytoliths) to the middle Miocene (terrestrial mammals, plant macrofossils).

The third question is: When did the C4 photo-

synthetic pathway evolve among grasses? The majority of plants, including trees, shrubs, and cool-growing-season grasses, use the C

3 photo-

synthetic pathway, in which atmospheric CO2 is

fi xed by carboxylation of ribulose biphosphate by the enzyme Rubisco. Most plants using the C

4 pathway are warm-growing-season grasses.

Although sedges and some dicots have also evolved C

4 photosynthesis, grassland ecosys-

tems in tropical and warm temperate regions today are dominated by C

4 grasses (Sage, 2004).

Plants using the C4 pathway have anatomi-

cal and biochemical adaptations that allow for the concentration of CO

2 within cells around

Rubisco prior to carbon fi xation. Plants using the crassulacean acid metabolism or CAM path-way, primarily cacti and other succulents, have a combination of C

3 and C

4 characteristics. CAM

plants are not thought to have ever formed a major component of grassland ecosystems, and we do not consider them further here.

C4 photosynthesis is fundamentally a set of

adaptations used to reduce or eliminate loss of CO

2 during photorespiration (oxidation of

ribulose biphosphate by Rubisco rather than carboxylation; Sage, 2004). Environmental conditions generally associated with greater effi ciency of C

4 photosynthesis, such as warm,

arid climates (Hattersley, 1983) or low atmo-spheric pCO

2 (Ehleringer et al., 1991, 1997;

Cerling et al., 1997), act by increasing photo-respiration rate but are probably not the direct selective factors for C

4 photosynthesis (Sage,

2004). The C4 pathway actually subsumes a

variety of biochemically distinct pathways that evolved independently numerous times among angiosperms (Sage, 2004), including at least 20 independent originations of different C

4 sub-

types among grasses (GPWG, 2001; Giussani et al., 2001; Hilu and Alice, 2001; Edwards and Smith, 2010; Edwards et al., 2010). Thus, the question of C

4 origin has different answers for

different clades, and the timing will likely vary on different continents depending on the histori-cal biogeography of specifi c clades of grasses.

The timing of evolutionary origins of C4

photosynthesis (i.e., fi rst appearance of species using C

4 photosynthesis) in different regions,

which presumably happened at different times in different clades in different regions, has not been determined defi nitively, but it can be constrained using a variety of approaches. The origin of C

4 photosynthesis can be somewhat

constrained by the fi rst appearance of plant macrofossils with the characteristic Kranz anatomy of C

4 plants (Tidwell and Nambudiri,

1989), although such fossils will necessarily underestimate the origin of C

4 grasses given the

low preservation potential of grasses. Molecu-lar clock studies (Christin et al., 2008; Edwards and Smith, 2010) have been used to estimate divergence times, although the poor grass fossil record limits calibrations, and molecular clocks have not yet been used in an explicitly biogeo-graphic context, which would be necessary to understand the origins of C

4 lineages in a given

region or continent, as opposed to globally. Cen-ters of origin have been informative for evaluat-ing both the locations in which and the environ-mental conditions under which various clades of C

4 dicots originated, but this approach is less

useful for the C4 clades of grasses and sedges,

for which centers of origin are still unclear due to greater species richness and apparently greater antiquity compared to C

4 dicots (Sage,

2004). Moreover, centers of origin cannot be calibrated to the time scale easily.

One powerful technique for recognizing the presence and estimating the abundance of C

4

plants in ancient ecosystems, at least above very low levels, is the stable carbon isotope composition of materials that inherit the C

3:C

4

ratio signature of the local environment, such as pedogenic carbonate (e.g., Quade et al., 1989; Kingston et al., 1994; Latorre et al., 1997; Kleinert and Strecker, 2001; Fox and Koch, 2003, 2004; Wang and Deng, 2005; Behrens-meyer et al., 2007), sedimentary organic com-pounds and biomarkers of terrestrial plants (e.g., Freeman and Colarusso, 2001; Huang et al., 2007), single pollen grains (Nelson et al., 2008), and the mineralized tissues of primary consum-ers (i.e., herbivorous animals; e.g., Morgan et al., 1994; Cerling et al., 1997; Latorre et al., 1997; Passey et al., 2002; Fox and Fisher, 2004; Wang and Deng, 2005), and, potentially, plants (i.e., phytoliths; Smith and White, 2004). This approach takes advantage of the difference in the



Fox et al.

446 Geological Society of America Bulletin, March/April 2012

fractionation of carbon isotopes during photo-synthetic fi xation of atmospheric CO

2 by C

3 and

C4 plants; it can constrain the appearance of C

4

plants in a region, but it will necessarily under-estimate the origin of C

4 plants in the region due

to the low preservation probability of suitable materials and the diffi culty of discriminating low levels of C

4 biomass from environmental

variation in the composition of C3 plants.

The primary control on the carbon isotope composition of C

3 plants is fractionation dur-

ing carboxylation of ribulose biphosphate by Rubisco, which discriminates strongly against 13CO

2 (O’Leary, 1981, 1988; Farquhar et al.,

1989). Considering smaller fractionations asso-ciated with steps prior to carboxylation, such as diffusion of CO

2 through stomata and disso-

lution of CO2 into cells and diffusion to chlo-

roplasts, the overall apparent carbon isotope enrichment during fi xation of CO

2 by C

3 pho-

tosynthesis is approximately −19.6‰ (Passey et al., 2002), and modern C

3 plants have a mean

δ13C value of about −27‰ (all carbon isotope values are expressed relative to Vienna Peedee belemnite [VPDB]). Because of the adaptations for CO

2 concentration in C

4 plants, the critical

carboxylation step (initially of phosphoenol-pyrovate in C

4 plants rather than the secondary

carboxylation of ribulose biphosphate) is not limiting; instead, stomatal diffusion, which does not discriminate against 13CO

2 as strongly as fi x-

ation by Rubisco, exerts the primary control on net isotope enrichment during fi xation of CO

2

by C4 photosynthesis. Apparent isotope enrich-

ment during C4 photosynthesis is about −4.7‰

(Passey et al., 2002), and modern C4 plants have

a mean δ13C value of about −13‰. Various envi-ronmental factors infl uence net fractionation during photosynthesis, so that both C

3 and C

4

plants exhibit considerable variability about their mean values, although the distributions of δ13C values do not overlap (O’Leary, 1988; Far-quhar et al., 1989). In grassland environments, the most important environmental infl uences are water and light stress, both of which decrease net fractionation and shift plants to higher δ13C values (Farquhar et al., 1989). Aridity in par-ticular primarily affects C

3 plants (Cerling and

Harris, 1999; Passey et al., 2002), and δ13C values of C

4 grasses in arid environments do

not exhibit any relationship with mean annual rainfall (Namibia—Schulze et al., 1996; South Africa—Swap et al., 2004).

Because of the dependence of the isotopic composition of C

3 plants on aridity and light

stress, mass balance interpretations of carbon isotope measurements cannot unambiguously distinguish the presence of C

4 plants at low

percentages of biomass from C3 ecosystems in

water- or light-stressed habitats; thus, the car-

bon isotopic approach can detect the presence of C

4 biomass using some threshold δ13C value

(see following) but will underestimate the time of origin in a region. In this regard, the carbon isotopic approach is not unlike the constraints provided by the plant macrofossil record.

The weight of evidence from all of these approaches indicates that C

4 grasses were pres-

ent on at least some continents prior to the Mio-cene and that C

4 photosynthesis likely evolved

initially, at least in grasses, during the Oligo-cene (Sage, 2004; Tipple and Pagani, 2007; Christin et al., 2008). This timing suggests a link between the evolution of C

4 photosynthesis

and the drawdown of atmospheric pCO2 from

the high levels of the Eocene greenhouse to the characteristically low levels of the Neogene and present icehouse (Pagani et al., 2005; Christin et al., 2008; Liu et al., 2009). Thus, the mechanism proposed by Cerling et al. (1997) for a late Mio-cene expansion of C

4 grasses may be correct in

principle but not in timing.The fi nal question, addressed in this paper

for the southern Great Plains, is: When did C4

grasses become ecologically dominant in tropi-cal and warm temperate grasslands? Essen-tially, this question concerns the origins of modern, C

4-dominated grassland ecosystems

of the tropical and warm temperate regions, which account for almost 25% of gross primary productivity on Earth today (Still et al., 2003; Edwards et al., 2010). The primary methods with which we can assess the abundance of C

4

grasses in an ancient ecosystem are the same carbon isotope approaches that can be used to constrain the appearance of C

4 grasses in an

ecosystem (i.e., pedogenic carbonate, sedi-mentary organic matter, mineralized tissues of herbivorous animals and plants). Pedogenic carbonates provide an integrated measure of the composition of ancient ecosystems that can record the C

3:C

4 ratio on time scales that

range from 103 to 105 yr, depending on the stage of carbonate development (Gile et al., 1966; Machette, 1985). Use of δ13C values of sedimentary organic matter can be limited by adequate abundance and the state of preserva-tion, and in many regions, such as the Great Plains, highly oxidized soils preserve little or no organic matter. Moreover, degradation of soil organic matter appears to follow a Ray-leigh distillation model in which refractory organic compounds have systematically higher δ13C values than original bulk soil organic mat-ter (Wynn, 2007). Reliance on the composition of fossil herbivores necessarily fi lters the eco-system signal through the feeding behavior and home range dynamics of individual animals, the behavioral ecology of individual species, and the ecological structure of animal communities.

Moreover, the reliance primarily on samples of the teeth of large-bodied mammals, a conse-quence of the ease of mechanical sampling and the rarity of laser ablation–isotope ratio–mass spectrometry systems (Passey and Cerling, 2006), introduces a potential bias in sampling that hides the critical ecological information that could be gained from sampling small-bodied herbivorous mammals (i.e., rodents and lagomorphs), which have much smaller home ranges on average (Harestad and Bunnel, 1979) and therefore sample the landscape at a fi ner spatial scale than do large-bodied mammals. Analysis of herbivorous mammals undoubt-edly is required to understand the structure of whole ecosystems, but it does not necessarily provide an integrated view of the composition of ancient biomes in the same manner as pedo-genic carbonates. Each of these techniques is suitable to answer specifi c questions, and all are valuable tools in reconstructing the origins and evolution of C

4-dominated ecosystems. How-

ever, given that pedogenic carbonate integrates the composition of soil biomass on geologically short time scales, in this study, we used the δ13C of pedogenic carbonates to examine the origin of the modern C

4 grassland of the southern

Great Plains.The appearance of grassland ecosystems

dominated by C4 grasses (i.e., >50% of bio-

mass) has been documented in greatest detail for South Asia based on the paleosol carbonate iso-tope record from the Himalayan foreland basin, which indicates that C

4 grasses reached modern

abundance in northern India and Pakistan during the late Miocene to Pliocene from ca. 8 to 3 Ma (Quade et al., 1989; Behrensmeyer et al., 2007). The initial recognition of an increase in C

4

biomass during the latest Miocene in Pakistan (Quade et al., 1989) and subsequent recognition of more or less coeval increases in consump-tion of C

4 grasses by large-bodied mammalian

herbivores in South Asia, North America, and South America (Cerling et al., 1997) have led to the somewhat entrenched view in the lit-erature of a synchronous, global increase in C

4

biomass during the late Miocene, commonly associated with the interval 8–6 Ma. However, contrasts in the timing and patterns of paleosol records from major grassland biomes of South America (Latorre et al., 1997; Kleinert and Strecker, 2001), North America (Fox and Koch, 2003, 2004), and Asia north and northeast of the Tibetan Plateau (Fan et al., 2007; Passey et al., 2009) do not support such a scenario. Rather, these records indicate that C

4 biomass in differ-

ent regions increased at different times over the last 10 m.y., at different rates, and not in strictly monotonic patterns. From the perspective of the entire Phanerozoic, it is clear that C

4-dominated





ecosystems are a late Neogene phenomenon, but from the perspective of the Neogene, the timing and patterns in different regions clearly vary.

In this paper, we present a detailed record of the evolution of the C

4-dominated grasslands of

the southern Great Plains based on the carbon isotope composition of multiple measured sec-tions in the Meade Basin of southwestern Kan-sas (Fig. 1). This is a higher resolution study of the transition to the modern, C

4-dominated

grassland in the Great Plains than prior stud-ies in the region based on either mammal teeth (Passey et al., 2002) or paleosol carbonates (Fox and Koch, 2003, 2004).

MEADE BASIN

The Meade Basin is an ideal setting for this study for several reasons. First, well-documented

late Miocene to late Pleistocene strata outcrop across ~2500 km2 along the Cimarron River and its tributaries, including Crooked Creek and its tributaries, and north and east from the Cimar-ron to the area around the town of Meade, Kan-sas (Fig. 1). The strata are dominantly fl uvial silts and sands that contain abundant pedogenic carbonate nodules and caliches or calcretes, as well as numerous superposed mammal faunas, particularly micromammal faunas (Martin and Fairbanks, 1999; Martin et al., 2000, 2002, 2003, 2008; Peláez-Campomanes and Mar-tin, 2005). The Miocene Ogallala Formation is exposed in several places to the south and southeast of Meade, Kansas. Lithostratigraphic nomenclature for the Pliocene–Pleistocene strata in the Meade Basin has varied somewhat among past workers. The name Rexroad For-mation has been used for early Pliocene rocks

N 37° 15'

W 100° 30'

1 3 5

kilometers

N

Crooked Creek

CimarronRiver

Lake MeadeState Park

54

160160

23

Plains

Meade

Mea

de C

ount

y

Sew

ard

Cou

nty Borchers N

Ogallala

HighBanks

Saw RockCanyon

XIT B

XIT A

KeefeCanyon

AlienCanyon

FoxCanyon

Wheelbarrow /Mustang

Borchers S

Aries

Kansas

Meade Basin

Figure 1. Map of the Meade Basin fi eld area. Symbols indicate locations of measured sec-tions discussed in the text. On the north side of the Cimarron River, symbols for some sec-tions completely overlap and cannot be distinguished.

exposed southwest of Meade, Kansas, and along the Cimarron River and its tributaries (Hibbard and Riggs, 1949; Hibbard, 1950), but our work has identifi ed several problems with the older stratigraphic models (Honey et al., 2005). To avoid confusion in the literature until we com-plete a full revision of the stratigraphy, we do not use a formal name for the early Pliocene strata. The late Pliocene and early Pleistocene strata in the area of Borchers Badlands (between the Cimarron River and Meade, Kansas) also have been referred to by various names and have been subdivided in various ways. Here, we treat these rocks as the Stump Arroyo and overlying Atwater Members of the Crooked Creek Forma-tion (Martin et al., 2008).

Second, the mammalian biostratigraphy of the Neogene strata is well understood and highly resolved (Fig. 2). Most of the Neogene strata exposed in the Meade Basin are Pliocene and Pleistocene, but at least two biostratigraphically distinct Miocene faunal levels are known in the fi eld area. The Pliocene–Pleistocene units in the Borchers Badlands unconformably overlie out-crops of the late Miocene Ogallala Formation that have produced a limited mammalian fauna assignable to the Clarendonian North American Land Mammal Age (NALMA; 12.5–9.0 Ma; Zakrzewski, 1988; Tedford et al., 2004). Out-crops of the Ogallala Formation to the south of Meade, Kansas (the High Banks locality), have a small, poorly sampled, and unpublished mam-malian fauna that is assignable to the Hemp-hillian NALMA (9.0–4.9 Ma; Tedford et al., 2004; Bell et al., 2004) based on the presence of a rhinoceros. This fauna could be latest Hemp-hillian (i.e., 6.0–5.0 Ma) based on the (apparent) absence of arvicolid and eomyid rodents and the presence of a primitive geomyoid with small, brachydont, rooted molars. Without better sam-pling, we treat this fauna only as Hemphillian (9.0–4.9 Ma) and note that it might correspond to the younger end of this interval after the appearance of horses with end-member C

4 diets

in the Coffee Ranch fauna of northern Texas (Passey et al., 2002). The presence of defi ni-tively Miocene strata beneath younger strata in the Meade Basin allows for direct comparison of both carbon and oxygen isotope composi-tions for Miocene and younger samples from the same fi eld area.

Most of the data in this study come from sec-tions that include mammal faunas of the Blan-can and early Irvingtonian NALMAs (4.9–2.0 and 2.0–ca. 1.0 Ma, respectively; Bell et al., 2004; Martin et al., 2008). At least 26 distinct faunal levels are known from numerous locali-ties from the early Pliocene to the level of the Lava Creek B ash (0.64 Ma; Lanphere et al., 2002) in the Meade Basin (Fig. 2). Age control



Fox et al.


on the mammalian biostratigraphy comes from comparisons to well-dated faunas elsewhere in North America, magnetostratigraphy for some sections (Lindsay et al., 1975; Martin et al., 2008), and the ages of the Huckleberry Ridge, Cerro Toledo B, and Lava Creek B ashes (2.1, 1.47–1.23, and 0.64 Ma, respectively; Izett et al., 1981; Izett and Honey, 1995; Lanphere et al., 2002), which outcrop in the Borch-ers Badlands section, Aries section, and the Cudahy and Subrite ash mines near Meade. We have not sampled carbonates in the inter-val between the top of the Aries section and the ash mines, and the sections at the mines do not have pedogenic carbonate, so our youngest samples occur a few meters above the Cerro Toledo B ash in the Borchers Badlands and Aries sections.

The Pliocene and early Pleistocene fau-nas can be divided into three stratigraphically and temporally distinct groups, following the informal subdivisions of the Blancan in Martin (2007). The sections in the southwestern part of the fi eld area from the canyons that run north and south to the Cimarron River (Fig. 1) can be correlated using the Bishop and Wolf Grav-els and the CC1 and CC2 caliches and have faunas considered here to be early Blancan (4.9–3.0 Ma). Some of the sections and asso-ciated faunas around the Cimarron River have limited magnetostratigraphic control (e.g., Fox Canyon; Lindsay et al., 1975), but those sec-tions do not contain ashes and cannot be corre-lated directly to the geomagnetic polarity time scale. The Wheelbarrow and Mustang sections to the northwest of Meade have taxonomi-cally limited, poorly sampled Blancan faunas that are clearly younger than the early Blan-can faunas along the Cimarron River, and they generally correlate to a series of better-sampled mid-Blancan faunas for which we do not have samples of paleosol carbonates (e.g., Sanders and Paloma; Fig. 2). We consider the Wheel-barrow/Mustang sections to be middle Blan-can (3.0–2.5 Ma). The boundary between the early and middle Blancan used here is younger than in Martin (2007) because of the inferred age of the Wolf Gravel at the top of the early Blancan sections (Fig. 2). Finally, sections in the Borchers Badlands/Aries area have a series of well-sampled, classic (Borchers, Nash 72) and new (Aries A and B) faunas with excellent geochronological control from magnetostratig-raphy (Martin et al., 2008) and the ages of the interbedded Huckleberry Ridge (2.06 Ma) and Cerro Toledo B (1.47–1.23) ashes (Izett et al., 1981; Izett and Honey, 1995). Based on the geochronological evidence, we consider these sections, including the poorly constrained por-tion below the Huckleberry Ridge ash and the

Ep

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6

78

10

11

12

13

14

15

9 ba

RZ

C3a

nC

3nC

2an

C2r

C1r

C2n

C3r

C2a

rC

1n

1

2r

n

nr

1

2

34n

n

n

n

r

3n

n

r

1

2

2r

2r

n

r

1 rn

1 rn

r

r

MPTS

6.57

9.00

12.50

6.276.145.89

5.234.984.894.80

4.624.48

4.294.18

3.583.33

3.223.11

3.04

2.58

2.152.141.951.77

1.070.990.78

Ma

High Banks

Ogallala Fm Hwy 23(age in NALMA notconstrained)

Saw Rock C. (?R)ArgonautFallen Angel (?R)

Fox Canyon (R)Bishop (R)

Ripley (R), ?XIT1 B-DKeefe C., Rapt. 1C (R)XIT1E, XIT2B (R)Wiens, Vasquez

Hornet

Bender 1B (N)Rexroad Loc. 3A-C (N)Deer Park, Rex 3D (R)Rexroad Loc. 2A (R)Rexroad Loc. 2

PalomaSanders (N)

Margaret

Borchers (R)

Aries NE (N)Nash 72 (R)

Aries B (R)

Arkalon, Cudahy (N)

A. A (R)

Faunas

Bishop Gravel

CC1

CC2

CCN2

CCN1

Wolf Gravel

Seger Gravel

Huckleberry Ridge (2.06 Ma)

Cerro Toledo B(1.47–1.23 Ma)

Lava Creek B(0.64 Ma)

Markerbeds

Irvi

ngto

nian

L H

emph

illia

nB

lanc

anM

EL

NA

LM

AC

lare

ndon

ian

Figure 2. Stratigraphic summary and chronology of Meade Basin mammal fau-nas. NALMA—North American Land Mammal Age; MPTS—magnetic polarity time scale; RZ—Rodent Zone of Martin (2003).





Borchers faunal level, to be late Blancan to early Irvingtonian (ca. 2.5–ca. 1.0 Ma).

The sections sampled within each of the three Blancan intervals can be correlated and contain biostratigraphically distinct faunas that are demonstrably superposed within each interval. However, the strata of the three time intervals are not physically superposed in any outcrop in the fi eld area, so we treat each as a separate interval and assume that the compos-ite sections for each do not overlap. Because not all of our measured sections have suffi -cient geochronological control and because sedimentation rates are probably quite variable both spatially and temporally, we do not con-vert meter levels in the measured sections into estimates of absolute age.

Finally, pedogenic carbonate is common throughout the entire Meade Basin sequence from the Ogallala Formation to the youngest levels sampled in this study, which are associ-ated with the Cerro Toledo B ash. Pedogenic carbonate morphologies include stages I (dis-persed carbonate fi laments and grain coatings), II (discrete nodules up to several centimeters in diameter or Bk horizons), and III (coalesced carbonate nodules or well-developed caliches or K horizons; Gile et al., 1966). Caliches or K horizons range in thickness from a few centi-meters to more than 2 m, but none exhibits the brecciation, lamination, or pisolite formation characteristic of stage IV and higher pedogenic carbonates (Gile et al., 1966; Machette, 1985). These carbonate morphologies precipitate from soil water on time scales of 102–103 yr (stage I) to 103–105 yr (Machette, 1985), although, in general, it seems that massive caliches (stage III and higher) must have formed on shorter time scales in the Great Plains than in the southwestern United States (Machette, 1985; Fox and Koch, 2003, 2004). Bulk samples of pedogenic carbonates integrate the C

3:C

4 ratio

of the landscape on the time scale of carbon-ate accumulation, which in this study is on the order of 102 yr to probably no more than 104 yr for the more massive stage III carbonates. A previous study of the carbon isotope com-position of paleosol carbonates in the Great Plains indicated that paleosol carbonates in the Meade Basin record the increase in C

4 biomass

to modern abundance (Fox and Koch, 2003, 2004), although sampling in that study was not adequate to resolve fi ne details of either the temporal or spatial patterns of increasing C

4

biomass in the region. The greater density of sampling both temporally and spatially in this study reveals the broad dynamics of ecosystem change in the southern Great Plains during the evolution to the modern C

4-dominated grass-

land over the late Neogene.

CARBON ISOTOPES IN PEDOGENIC CARBONATES

Authigenic soil carbonate forms in carbon isotopic equilibrium with CO

2 dissolved in soil

water (Cerling, 1984; Cerling et al., 1989). At shallow depths within the soil profi le, soil CO

2

is a mixture of atmospheric CO2 and biogenic

CO2 derived from root respiration and micro-

bial oxidation of soil organic matter. At typical soil respiration rates and modern atmospheric composition, soil pCO

2 below ~30 cm in a soil

is suffi ciently high to prevent deeper penetra-tion of atmospheric CO

2. Consequently, soil

CO2 at depth is primarily biogenic and has a

δ13C value related to that of overlying biomass (Cerling et al., 1991). Overall, pedogenic car-bonate formed at depth refl ects the contribu-tions of isotopically distinct C

3 and C

4 biomass

to soil CO2 but is enriched in 13C by 4.4‰ due

to diffusion of 12CO2 from the soil profi le to the

atmosphere and by an additional 9‰–12‰ due to temperature-dependent carbon isotope frac-tionation during carbonate precipitation (Cer-ling et al., 1989, 1991). Thus, if soil respiration is moderate to high, the δ13C value of soil car-bonate is controlled by the relative abundance of C

3 and C

4 biomass growing in the soil during

carbonate accumulation.The relative proportion of C

4 biomass that

contributed to soil-derived CO2 during carbon-

ate accumulation can be estimated from the δ13C value of paleosol carbonate by simple linear mixing between end members with the mean δ13C value of C

3 and C

4 plants. However,

these end-member values will vary with the δ13C value of atmospheric CO

2, which is known to

have decreased over the past 150 yr from a pre-industrial value of ~−6.5‰ to the modern value of ~−8‰ by the addition of isotopically light anthropogenic CO

2 (Friedli et al., 1986). On lon-

ger time scales, the composition of atmospheric CO

2 has varied with the evolution of and pertur-

bations to the long-term carbon cycle (Leuen-berger et al., 1992; Zachos et al., 2001). Thus, to estimate the proportion of C

4 biomass from the

δ13C value of Neogene paleosol carbonates, we must constrain the δ13C of Neogene atmospheric CO

2. One approach (i.e., that used by Fox and

Koch, 2003, 2004) is to use the observed prein-dustrial value of −6.5‰ (Friedli et al., 1986) and assume that plant biomass prior to the industrial era was depleted in 13C by 1.5‰ relative to mod-ern biomass. However, this approach risks small but systematic errors in estimates of the propor-tion of C

4 biomass if atmospheric CO

2 varied in

composition systematically during the past.To estimate the δ13C values of atmospheric

CO2 and C

3 and C

4 plants during the time rep-

resented by our samples (Clarendonian to early

Irvingtonian NALMAs or late Miocene to early Pleistocene) and the Miocene data set of Fox and Koch (2003, 2004; late Arikareean to Hemphillian or early Miocene to late Miocene), we used the method and parameters of Passey et al. (2002), who based their approach on several earlier studies (Koch et al., 1995; Ekart et al., 1999; Pagani et al., 1999a, 1999b). The δ13C of atmospheric CO

2 through time was estimated

based on the δ13C values of planktonic foramin-ifera from Deep Sea Drilling Project (DSDP) Sites 586 (Whitman and Berger, 1993) and 588 (Pagani et al., 1999b) and the apparent isotope enrichment factor for planktonic foraminifera relative to atmospheric CO

2 (+7.9‰ ± 1.1‰;

Passey et al., 2002) for modern, preindustrial, and late Pleistocene and Holocene (<30 ka) spe-cies of planktonic foraminifera. Estimated δ13C values of atmospheric CO

2 (Fig. 3A) clearly

exhibit a long-term temporal decrease from the high during the Barstovian, and all but three estimated values are greater than the preindus-trial value of −6.5‰ measured in Antarctic ice (Friedli et al., 1986). However, the intercept of a reduced major axis (RMA) regression for the Barstovian and younger samples with age as the independent variable and estimated δ13C value of atmospheric CO

2 as the dependent variable

(−6.47‰, calculated with RMA for Java 1.21; Bohonak and van der Linde, 2004) is essentially identical to the measured preindustrial value.

If we had precise geochronological control throughout the study interval and estimates of local sedimentation rates for all sections, we could use either an age model or smaller tempo-ral bins to assign estimated δ13C values of atmo-spheric CO

2 throughout our measured sections

to account for the trend. Instead, we binned the mean estimated δ13C values by the NALMAs for the Miocene and subdivisions of the NAL-MAs for the Pliocene through early Pleistocene (as discussed in the previous section) and used the mean δ13C value in each NALMA or sub-division (Fig. 3B). This approach captures the long-term trend in Figure 3A and, with the exception of the prolonged excursion to higher values during the Hemphillian, appears to accu-rately represent the δ13C values in each interval. For the Borchers Badlands sections, we used the estimated δ13C of atmospheric CO

2 for the

late Blancan for samples below the Huckleberry Ridge ash and the estimate for the Irvingtonian for the samples above the ash.

We convert measured δ13C values of paleosol carbonate into estimates of percent C

4 biomass

by linear mixing assuming a fractionation fac-tor between CO

2 dissolved in soil water and

carbonate of +15‰, which is the midpoint of the observed range given diffusion and temperature-dependent fractionation (Cerling



Fox et al.


and Quade, 1993). After adjusting the C3 and C

4

end members (δ13CC3

and δ13CC4

, respectively) based on the estimated δ13C of atmospheric CO

2

(Fig. 3B) and the apparent isotope enrichment factors for plant biomass relative to atmospheric CO

2 used by Passey et al. (2002; −19.6‰ and

−4.7‰, respectively; Figs. 3C and 3D), the per-centage of C

4 biomass can be estimated from

measured δ13C values (δ13Cmeas

) by solving the

following mass balance equation for X (percent C

4 abundance):

δ13Cmeas

= δ13CC4

X + δ13CC3

(1 – X). (1)

This approach assumes negligible contribution to soil CO

2 from CAM plants and soil respira-

tion rates that are consistently high enough so that the partial pressure of soil-respired CO

2

excludes penetration of atmospheric CO2 from

the site of carbonate formation. If the con-centration of soil CO

2 were consistently low

throughout the study interval, as suggested by Breecker et al. (2009) for stage I Holocene soil carbonates developed on gravel clasts, carbon-ates would record a mixture of soil-respired CO

2

and isotopically heavy atmospheric CO2, which

could suggest a spuriously early appearance of

Figure 3. Modeled δ13C values of atmospheric CO2 and C3 and C4 biomass for the late Miocene to early Pleistocene. Time scale is also divided into North American Land Mammal Ages (NALMAs). Abbreviations: Pl—Pleistocene; Clarendon—Clarendonian NALMA; Irv —Irvingtonian NALMA. Ages for boundaries are based on Gradstein et al. (2004), Tedford et al. (2004), and Martin et al. (2008). (A) Estimated δ13C value of atmospheric CO2 based on δ13C values of planktonic foraminifera in Whitman and Berger (1993) and Pagani et al. (1999b) and estimation method and parameters of Passey et al. (2002). Gray bars are ±1 s.d. (1.1‰) based on estimate of the apparent enrichment between planktonic foraminifera and atmospheric CO2 from Passey et al. (2002). Solid line is reduced major axis (RMA) linear regression with age as the independent variable and estimated δ13C value of atmospheric CO2 as the dependent variable (δ13C = 0.099 Ma – 6.47‰, R2 = 0.58). (B) Modeled δ13C values (solid line) used to calculate δ13C values of C3 and C4 biomass in each NALMA for mass balance estimation of percent C4 biomass from measured paleosol carbon-ate δ13C values. Open circles are the same data as in A. In each pair of numbers, the upper number is the mean δ13C value for the NALMA, and the lower number is the standard deviation of the estimated δ13C values in the same NALMA. (C) Modeled δ13C values of C3 biomass in each NALMA. (D) Modeled δ13C values of C4 biomass in each NALMA. VPDB—Vienna Peedee belemnite.





C4 biomass. However, in this case, carbonates

would still record any changes in the relative abundance of C

4 biomass over time. If the con-

centration of soil CO2 varied substantially dur-

ing the study interval, then carbonates would suggest parallel temporal variability in the abun-dance of C

4 biomass.

Water stress due to aridity is a potential con-founding factor in the interpretation of paleosol carbonate δ13C values because it substantially decreases net fractionation of atmospheric CO

2

during photosynthesis in C3 plants (Farquhar et

al., 1989; Cerling and Harris, 1999; Passey et al., 2002). Consequently, paleosol carbonates formed from C

3 soil biomass under conditions of

light and/or water stress would have positively shifted δ13C values suggestive of the presence of a small fraction of C

4 biomass. We did not adjust

our estimates of the abundance of C4 biomass

for the possible impact of aridity on C3 biomass

δ13C values but instead used two approaches in this paper to consider the possible impact of aridity on our interpretations of paleosol carbon-ate δ13C values. In the companion paper (Fox et al., 2011), we also used the relationship between paleosol carbonate δ13C and δ18O values in each interval as a direct check on the control of local climate and hydrological balance on the compo-sition of paleosol carbonate.

First, following Fox and Koch (2003), we used the δ13C values of 20 Holocene paleosol carbonates from arid-climate C

3 ecosystems

in North America, Russia, and the Eastern Mediterranean (Quade et al., 1989; Cerling and Quade, 1993; Wang and Anderson, 1998; Mon-ger et al., 1998; Khokhlova et al., 2001) to estab-lish threshold δ13C values for the presence of C

4

biomass in each NALMA interval. The mean δ13C value for these carbonates is −8.7‰ ± 1.06‰ (all means reported ±1 standard devia-tion), and mean annual precipitation for 15 of these sites is 567 ± 186 mm (fi ve sites did not include mean annual precipitation). For the period 1949–2006 (excluding 1973, which did not have complete data), mean annual rainfall in Meade, Kansas, was 545 ± 132 mm, mean annual snowfall was 471 ± 239 mm, and mean annual total precipitation was 1016 ± 289 mm (Meade station National Climatic Data Center [NCDC] Summary of the Day, www.ncdc.noaa.gov/oa/ncdc.html). Snowfall generally converts to a lower equivalent rainfall, but the historical precipitation data for Meade indicate that the Holocene paleosol carbonates provide a con-servative estimate of the infl uence of aridity on the composition of C

3 biomass in the region in

the past. Conventionally, a long-term increase in aridity is presumed to have been an important factor in the evolution of the modern treeless prairie of the Great Plains over the Neogene

(e.g., Axelrod, 1985); thus, the modern condi-tions conceivably represent the most arid cli-mate during our study interval.

We assumed that the Holocene carbon-ates formed under atmospheric CO

2 with a

δ13C value close to the preindustrial value of −6.5‰ (i.e., lower than any interval in the study period); therefore, to establish a threshold that accounts for the secular variation in the δ13C value of atmospheric CO

2 (Figs. 3A and 3B), we

increased the mean value of the carbonate δ13C threshold in each NALMA or subdivision by the difference between the preindustrial δ13C value of atmospheric CO

2 and the modeled δ13C value

for the interval (Fig. 3B). Within each interval, the distribution of the adjusted δ13C values for the Holocene carbonates represents the arid extreme for ecosystems without C

4 biomass.

Any carbonate δ13C values outside that distribu-tion unambiguously require the presence of C

4

biomass independent of the effects of aridity on C

3 biomass. The scales for percent C

4 biomass in

the summaries of our data herein could be reset to the mean of the threshold value under the assumption that conditions in the region in the past were more arid than today, although such aridity is not well supported by any evidence.

The second approach is to model expected paleosol carbonate δ13C values for arid condi-tions using simple linear mixing given the esti-mated δ13C value of atmospheric CO

2 in each

interval and the observed change in apparent enrichment for C

3 biomass in arid environments

today (−16.7‰ for arid vs. −19.6‰ for nonarid conditions; based on values in Passey et al., 2002). In this approach, we did not assume any change in apparent enrichment for C

4 biomass,

based on the lack of a relationship between mean annual precipitation and δ13C values of modern C

4 grasses from arid environments (Namibia—

Schulze et al., 1996; South Africa—Swap et al., 2004). This approach yielded a single value for each interval (as opposed to a distribution as in the fi rst approach) that was close to, but slightly more negative than, the most positive value of the approach based on Holocene carbonates for each interval.

Under closed-canopy vegetation, isotopically light respired CO

2 mixes with ambient CO

2 so

that subsequent photosynthesis utilizes CO2

with lower δ13C values relative to well-mixed atmospheric CO

2, yielding plant biomass that

is negatively shifted relative to the average δ13C value for the photosynthetic pathway. The can-opy effect (van der Merwe and Medina, 1991) would be recognized as paleosol carbonate with unusually low δ13C values. As is clear in the fol-lowing discussion, we do not see any indications of the canopy effect in our data and do not con-sider it as a factor in our interpretations.

To compare our estimates of the percentage C

4 biomass in the past to modern conditions, we

used 29 published δ13C values of modern and Holocene soil organic matter and standing plant biomass and four published values of percent C

4

biomass from Kansas and northwestern Okla-homa (Table 1; data and sources in Table DR11) to estimate the percentage of C

4 biomass in the

region today. Published δ13C values were con-verted to percentage C

4 biomass using the mass

balance as described previously. For δ13C values from modern samples, we assumed a δ13C value for atmospheric CO

2 of −8.0‰; for Holocene

values, we assumed the preindustrial value of −6.5‰ (Friedli et al., 1986); for values in sam-ples with an age range from modern to a date in the Holocene, we assumed the midpoint of the approximate modern and preindustrial values (−7.25‰). For all data (n = 33), the mean abun-dance of C

4 biomass in the region today is 75%

± 12.8%. However, given that the abundance of C

4 biomass might have changed in the past sev-

eral thousand years (Johnson et al., 2007), we compared our data only to the mean for the 20 strictly modern values (78% ± 10.9%).

For both carbon and oxygen isotopes, an assumption of most previous studies has been that paleosol carbonates record mean annual environmental and soil conditions. However, a recent study of carbon and oxygen isotope variation in Holocene stage I paleosol carbon-ates developed on fl oodplain gravels and mod-ern soil CO

2 and water in the Chihuahuan Des-

ert, New Mexico, indicated that soil carbonate isotopic records are seasonally biased to warm and dry conditions when net evaporation of soil water drives carbonate precipitation (Breecker et al., 2009). The climate regime in the Great Plains is broadly similar to that in New Mexico where Breecker et al. (2009) studied Holocene soil carbonates. For Meade, on average, 66% of annual precipitation falls during the warm months of May–September (Meade Station NCDC Summary of the Day, www.ncdc.noaa.gov/oa/ncdc.html). Although most precipita-tion occurs during the warm period, given the intense, episodic, and localized nature of most storms in the region during the warm season, it is likely that soil water content is actually low-est during the warm months, as is generally the case for grassland soils in northeastern Colo-rado (Ferretti et al., 2003). Thus, it is likely that at least modern carbonates in the Meade area

1GSA Data Repository item 2011262, Stable isotope data from pedogenic carbonates in mea-sured sections, is available at http://www.geosociety.org/pubs/ft2011.htm or by request to editing@ geosociety.org.



Fox et al.


may also be biased to the warm months, if not also the Neogene carbonates we sampled. Since the seasonal abundance of C

4 biomass should

be highest during the warm (and dry) season, a warm-season bias in the formation of pedogenic carbonate would lead to an overestimate of the proportion of C

4 biomass from the mass balance

if the estimates are considered annual propor-tions. If our samples are warm season biased, then, given our results here, the comparisons to the mean modern abundance of C

4 biomass

in the region would indicate an even later rise to ecological dominance of C

4 grasses than we

suggest. In the companion paper (Fox et al., 2011), we consider in detail the consequences of different assumptions about the timing of carbonate precipitation (i.e., mean annual con-ditions, integrated monthly conditions, or warm season bias) in the paleoclimatic interpretation of the δ18O values.

METHODS

Field Methods

Hand samples of carbonates were collected from 19 measured sections in the Meade Basin fi eld area (Fig. 1) that range in length from only ~3 m to 38 m. Sections were measured to a pre-cision of 10 cm. Comparisons of sections mea-sured during carbonate sampling by one author (Fox) with sections of the same outcrops (but not necessarily the same location) measured in earlier fi eld seasons by another author (Honey) agree to within 0.5 m (Honey et al., 2005). Car-bonates were identifi ed in the fi eld as pedogenic based on external morphology, inclusions of

clastic sediment indicating engulfi ng growth of carbonates, and presence of rhizoconcretions in the associated clastic sediments or in super- and subjacent beds. Only stage II (nodular) and III (coalesced nodules to caliche) carbonates (Gile et al., 1966) were collected for stable isotope analysis. Nodules were collected in two modes: >30 cm below an obvious stratigraphic bound-ary and either at regular intervals of 50–100 cm or opportunistically (depending on carbonate abundance) from intervals without obvious soil horizonation. Nodules were collected from trenched or otherwise freshly exposed surfaces to avoid weathered material. Single samples were collected from the lower surface of thin stage III carbonates. Samples were also col-lected from the upper surface of thicker stage III carbonates and occasionally also from the mid-dle of particularly thick caliches. Large blocks of caliche were removed with a rock hammer or mattock and then reduced with a rock ham-mer to remove the weathered surface. Although we collected some stage I carbonates (dispersed fi laments and grain coatings) in bulk clastic sediment samples, these were not measured for stable isotope composition.

Laboratory Methods

Field samples were cut with a lapidary saw, rinsed in water, and dried to expose a clean internal surface for sampling. Samples were drilled under stereoscopic observation using a 0.5 mm diamond bur. Sample holes were small relative to the size of most fi eld samples, but they still yielded an excess of powder relative to the amount needed for mass spectrometry.

TABLE 1. SUMMARY STATISTICS OF δ13C VALUES AND ESTIMATED PERCENT C4 BIOMASS FOR EACH BIOSTRATIGRAPHIC INTERVAL

δ13C(‰, VPDB)

Mean SD N Max Min RangeLate Blancan–early Irvingtonian All data –2.6 1.18 39 0.9 –4.4 5.3

Without outliers –2.6 1.21 36 0.9 –4.4 5.3Middle Blancan –5.0 1.10 27 –3.0 –6.2 3.2Early Blancan –4.9 0.90 106 –1.8 –7.5 5.7Hemphillian –6.5 0.31 10 –5.9 –6.8 0.9Clarendonian –7.6 0.90 12 –5.9 –8.7 2.8Miocene combined –7.1 0.91 22 –5.9 –8.7 2.8All –4.7 1.59 194 0.9 –8.7 9.6

Percent C4

Modern 78.1 10.86 20 94.0 60.7 34.0Modern and Holocene 75.0 12.80 33 96.6 48.3 48.3Late Blancan–early Irvingtonian All data 54.5 8.27 39 78.5 43.0 35.6

Without outliers 55.0 8.44 36 78.5 43.0 35.6Middle Blancan 38.3 7.40 27 51.8 30.1 21.7Early Blancan 37.6 6.06 106 58.5 20.1 38.4Hemphillian 25.8 2.10 10 29.8 23.5 6.3Clarendonian 17.3 6.06 12 29.0 10.3 18.7Miocene combined 21.2 6.3 22 29.8 10.3 19.5All 39.2 11.46 194 78.5 10.3 68.3

Note: Data for the late Blancan–early Irvingtonian are reported both with and without three outliers that have δ18O values >29‰. VPDB—Vienna Peedee belemnite.

For a subset of fi eld samples (n = 108) col-lected from 13 measured sections that ranged in age from early Blancan (Saw Rock Canyon) to early Irvingtonian (Borchers Badlands, Aries Quarry), two separate samples were drilled from different locations on the same cut face of individual fi eld samples to examine internal isotopic heterogeneity at a relatively coarse spatial scale (i.e., these fi eld samples were not microsampled in detail). Because these samples generally exhibited low internal variability (see Results), they are reported as the midpoint value both here and in Fox et al. (2011); only one sample was drilled from the remainder of the fi eld samples (n = 86). In total, we measured the carbon and oxygen isotope composition of 194 fi eld samples in 302 unique stable isotope analyses. Sample powders were roasted in vacuo at 400 °C for at least 1 h to combust any organic matter and to remove water. The major-ity of samples were analyzed at the University of Minnesota Stable Isotope Laboratory, where samples were reacted with phosphoric acid at 70 °C in a Kiel automatic carbonate extraction device, and the isotopic composition of the resulting CO

2 was measured using a Finnigan

MAT 252 gas-source isotope-ratio mass spec-trometer. Precision was maintained by repeated measurements of NBS-18 and NBS-19 and Carrara marble laboratory standard during each run of samples, and the samples were normal-ized to the composition of NBS-19. Samples from three sections (Aries, Borchers Badlands north of Kansas Hwy 23, Alien Canyon) were analyzed at the University of California–Santa Cruz, where samples were reacted with phos-phoric acid at 90 °C in a Micromass Isocarb automatic carbonate extraction system, and the isotopic composition of the resulting CO

2

was measured on either a Micromass Prism or Optima isotope-ratio mass spectrometer. These samples were normalized to the composition of Carrara marble laboratory standard, and the normalization was checked against the cor-rected value of NBS-19 run in sequence with the samples (Fox and Koch, 2003, 2004). Data are expressed in standard δ notation as the per mil difference between the ratios of heavy to light isotope abundances (13C/12C or 18O/16O) in a sample and a standard material (e.g., δ13C = [{13C/12C sample – 13C/12C standard}/13C/12C standard] × 1000). Carbon isotope data are expressed relative to Vienna Peedee belemnite (VPDB), and oxygen isotope data are expressed relative to Vienna standard mean ocean water (VSMOW). Precision for all analyses is 0.1‰ or better based on replicate analyses of labora-tory and NBS standards at both laboratories.

Five samples from the Borchers section were analyzed in both laboratories. For four





of these samples, the difference in δ13C values between the mean of two replicate analyses at the University of California–Santa Cruz and single analyses of the same field sample at the University of Minnesota was either 0.0‰ or 0.1‰. For these four samples, the differences in δ18O between the mean for the replicate analyses at the University of Cali-fornia–Santa Cruz and the single analysis at the University of Minnesota were 0.0‰, 0.1‰, 0.2‰, and 0.5‰. For the fifth sample, the difference in δ13C values was 3.4‰ (the maximum intrasample difference in the data set; see results for intrasample variability in the following section), but the difference in δ18O values was only 0.5‰. We conclude that reproducibility between the laboratories is good and that the variability in the fifth sample reflects carbon isotopic heterogene-ity in that particular sample, reflecting local habitat change on the time scale of carbon-ate precipitation.

To analyze stratigraphic trends in the data, we used RMA linear regressions rather than ordinary least squares (OLS) linear regressions because in all cases the independent variable was measured with error (in addition to the dependent variable), which violates an impor-tant assumption of OLS regression. In general, regression analyses of stable isotope values on meter level or age and of δ13C values on δ18O values (or vice versa) should use RMA and not OLS regression because of this violation of the assumptions of OLS, although RMA has rarely been used in the geochemical literature. RMA regressions and 99% confi dence inter-vals (CIs) for slopes were calculated based on the linear method of Sokal and Rohlf (1994) using RMA for Java 1.21 (Bohonak and van der Linde, 2004); slopes for which the 99% confi dence interval does not include 0 are con-sidered statistically distinct from 0, implying a statistically signifi cant stratigraphic trend in the data. In general, RMA regressions yield higher slopes than OLS for the same data (Sokal and Rohlf, 1994). Given this, and the greater familiarity of most workers with OLS regression, we also report the results of OLS for all regression analyses in Table DR2 (see footnote 1), but we do not discuss these results in detail here. In most cases, the results of the two regression types for the same data are concordant. Those cases in which only one method yielded a signifi cant result do not change our overall interpretations. All other statistical tests (Spearman’s rank correlation, t-test, Mann-Whitney U-test, analysis of vari-ance [ANOVA] with post hoc Scheffé test for multiple comparisons) were done using SPSS 13.0 for Macintosh.

RESULTS

Intrasample Variability

Replicate samples from different parts of 108 fi eld samples exhibit low variability for both carbon and oxygen isotopes, and intrasample variability does not exhibit systematic patterns within or between sections (Fig. 4; all paleosol carbonate data are reported in Table DR3 [see footnote 1]). The mean difference in δ13C value within samples is 0.19‰ ± 0.38‰. The distri-bution of intrasample differences in δ13C value is right skewed, with a median of 0.11‰ and a mode ≥0.0‰ and <0.10‰ (Fig. 4A). Only fi ve fi eld samples have intrasample differences greater than 0.5‰, and the maximum of these is 3.4‰, which is more than twice the next high-est intrasample difference (1.55‰). Without the sample with the highest intrasample differ-ence, the mean for the intrasample differences in δ13C values is 0.16‰ ± 0.22‰. Thus, almost all samples have low internal variability relative to the range of values associated with C

3 and C

4

end members, at least at the spatial scale utilized here to sample fi eld samples. Internal variability for oxygen is generally similar to that for car-bon. The mean difference in δ18O value within samples is 0.20‰ ± 0.19‰. The distribution of intrasample differences in δ18O value is also right skewed, with a median of 0.15‰ and a mode that is also ≥0.0‰ and <0.10‰ (Fig. 4B). Eight samples have intrasample differences greater than 0.5‰, and the maximum of these is 1.16‰. Although the intrasample differences in δ13C and δ18O values of the same samples are correlated to a statistically signifi cant degree, the correlation is weak (Spearman’s rank corre-lation coeffi cient ρ = 0.22, p = 0.026; Fig. 4C).

Given the limited mean internal variability within samples, the replicate analyses from these fi eld samples will be represented by the midpoint value in the rest of this paper, and the values from those samples for which we ana-lyzed only single samples can be considered representative. However, given that some sam-ples have higher variability, it is possible that at least some pedogenic carbonates in the Meade Basin have complex diagenetic histories that could be examined with detailed microsampling to resolve variations in isotopic compositions on time scales short relative to the duration of car-bonate deposition.

Late Miocene: Clarendonian (12–9 Ma) and Hemphillian (9–4.9 Ma) Sections

Based on the faunal data available, the two Miocene sections in the Ogallala Formation are distinguishable by age and are discussed sepa-

Figure 4. Histograms of intrasample vari-ability in (A) δ13C values and (B) δ18O val-ues. (C) Correlation between intrasample variability in δ13C and δ18O values for indi-vidual samples and Spearman’s rank corre-lation coeffi cient (ρ). N = 108 in all.

rately here; in a later discussion, we treat these sections both separately and as a grouped Mio-cene data set for the Meade area. The Clarendo-nian Ogallala section below the Crooked Creek Formation in the Borchers Badlands (Fig. 5; summary statistics for all sections are presented in Table 1; stratigraphic columns for each



Fox et al.


Figure 5. Composite stratigraphic section for pedogenic carbonate samples around Meade, Kansas, with measured δ13C values (left) and estimated percent C4 biomass (right) in each biostratigraphic interval. For δ13C values in each interval, long dashed line and light-gray box indicate mean δ13C value ± 1 standard deviation for 20 Holocene paleosol carbonates from arid-climate C3 ecosystems in North America, Russia, and the Eastern Mediterranean (sources in text), adjusted for estimated δ13C of atmospheric CO2 for the interval in Figure 3A; short dashed line indicates arid C3 end member based on enrichment of Passey et al. (2002) and δ13C of C3 biomass for each interval in Figure 3C. Lines through δ13C values in each interval are reduced major axis (RMA) regressions (see Table 2 for parameters). Box and whiskers plot indicates the mean (thin vertical line), median (thick vertical line), 25th and 75th percentiles (left and right edges of box, respectively), and 10th and 90th percentiles (tips of left and right whiskers, respectively) of δ13C values of Miocene carbonates from across the Great Plains from Fox and Koch (2003, 2004). Percent C4 biomass in each interval was calculated by mass balance based on measured δ13C values and estimated end-member values in Figure 3; heavy dashed line and light-gray box indicate mean modern abundance of C4 biomass in the region ± 1 standard deviation (see Table DR1 [see text footnote 1]). VPDB—Vienna Peedee belemnite.





biostratigraphic unit are presented individually in the supplemental materials as Figs. DR1–DR5 [see footnote 1]) has a mean δ13C value of −7.6‰ ± 0.90‰, and δ13C values range from −8.7‰ to −5.9‰ (n = 12). RMA regression with meter level of samples as the independent variable indicates that δ13C values increase up section to a statistically signifi cant degree by 0.27‰ per meter (Table 2). The section is not long (~12 m), but it has carbonates through-out, including several relatively thick, well-developed caliche horizons. Although we do not have internal geochronological control in the section, the abundance and development of the carbonates suggest that this section could repre-sent hundreds of thousands to possibly several million years. Based on mass balance and the estimated δ13C of atmospheric CO

2 for the Clar-

endonian, these values correspond to an average of 17% C

4 biomass and a range from 9% to 29%

(Table 1). However, only two samples have high enough δ13C values that they unambiguously require the presence of C

4 biomass; even under

the assumption of arid conditions, these two samples still require ~12% C

4 biomass.

The Hemphillian High Banks locality in the Ogallala Formation (Fig. 5) exhibits some slight differences in both carbon and oxygen compared to the Ogallala Formation below Borchers. The mean δ13C value at High Banks is −6.5‰ ± 0.31‰ (n = 10). The δ13C values range from −6.8‰ to −5.9‰ and do not change sig-nifi cantly up section (Table 2). This section also contains abundant carbonate throughout and several well-developed caliches and, although it is somewhat shorter than the Clarendonian section, presumably represents sedimentation and carbonate accumulation over a similar time scale. The δ13C values correspond to an average of ~25% C

4 biomass and a range in abundance

during carbonate formation at High Banks of only 23% to ~30% C

4 biomass (Table 1). All

measured δ13C values at High Banks fall out-side of the range for arid condition δ13C values for the Hemphillian. Thus, the younger local-ity refl ects an increase in the abundance of C

4

biomass and less variation relative to the Clar-endonian section. However, no estimates of the

percent C4 biomass for either section fall within

the modern range.

Early Pliocene: Early Blancan (4.9–3.0 Ma) Sections

The early Blancan sections to the south and north of the Cimarron River are treated as a group because they are well correlated based on the repeated occurrences of the Bishop and Wolf Gravels and the CC1 and CC2 caliche lay-ers. Honey et al. (2005) discussed correlations among the sections north of the Cimarron River. Here, we add the localities to the south of the river and use a correlation scheme that differs in detail from that of Honey et al. (2005) due to both slight differences in the locations where sections common to both studies were measured and the broader geographic coverage of the sec-tions measured in this study. Given the isotope compositions from these sections, the differ-ences in correlations do not change our results in a meaningful way. The lowest parts of the XIT-B section are stratigraphically below the Saw Rock Canyon fauna, but the age is not con-strained. However, the sediments are not similar to the Ogallala Formation, and thus we interpret all of the XIT-B section as early Blancan. The Fox Canyon samples are from a series of short, correlated measured sections along ~0.5 km of the east side of Fox Canyon (these sections are not fi gured in the supplemental materials). Based on the faunas, the early Blancan sections represent accumulation from ca. 5.0 to 3.0 Ma, which is consistent with the presence of abun-dant carbonate nodules (except in the gravels) and a few well-developed caliches.

The mean δ13C value for the early Pliocene composite section is −4.9‰ ± 0.90‰ (Table 1). Through the section, δ13C values exhibit consid-erable variability in well-sampled intervals, and the overall range is from −7.5‰ to −1.8‰. Vari-ability within three of the six early Pliocene sec-tions is greater than for all of the early Pliocene data taken together, an indication of the strong central tendency in the data. At the base of the composite early Blancan section, the XIT-B section exhibits an increasing trend that could

be suggestive of an increase over the values of the Hemphillian High Banks section, and at the top of the composite section, the Alien Canyon and XIT-A sections also suggest a stratigraphic trend to even higher δ13C values (Fig. 5). It is possible that both of these intervals record per-manent changes in the regional abundance of C

4 biomass. However, stratigraphic variability

in the early Blancan data is highest where sam-pling density is highest, such as around 40–45 m in the composite section (i.e., around the CC1 carbonate level). Based on RMA regression, the δ13C values decrease signifi cantly through the section, but the line does not fi t the bottom and top of the section well, and the variance explained is extremely low (R2 < 0.01; Table 2). Without additional sections that sample the top and bottom of the early Blancan composite sec-tion, it is diffi cult to determine whether or not the evident trends represent only very local-ized variability in the abundance of C

4 biomass.

Early Blancan carbonates indicate an average of 38% C

4 biomass, with a range of estimates from

20% to 58% C4 biomass for individual samples

(Table 1). All data are more positive than the arid end members for the early Blancan; if arid-ity in the early Blancan were extreme, then the estimated C

4 biomass would range from 2% to

50%. No estimates of the percent C4 biomass

fall in the modern range.

Mid-Pliocene: Middle Blancan (3.0–2.5 Ma) Sections

The middle Blancan samples come from a series of short measured sections along ~1 km of a single shallow arroyo. The mean δ13C value of these samples (−5.0% ± 1.10‰; Fig. 5; Table 1) is similar to that for the early Blancan samples. The range of values in the middle Blancan is about half that of the early Blancan, but the sam-ple size in middle Blancan is smaller (n = 27), and the standard deviation is actually slightly higher, so a larger sample from this interval would likely exhibit a range in values compa-rable to that of the early Blancan. Like the early Blancan samples, the δ13C values decrease up section (by 0.7‰ per meter). Based on the RMA

TABLE 2. REGRESSION STATISTICS FOR REDUCED MAJOR AXIS (RMA) REGRESSIONS WITH METER LEVEL OF SAMPLES IN COMPOSITE SECTIONS FOR EACH BIOSTRATIGRAPHIC INTERVAL AS X

VARIABLE AND δ13C VALUE OF PALEOSOL CARBONATES AS Y VARIABLE

Slope S.e. 99% CI Intercept R 2

Late Blancan–early Irvingtonian All data 0.18 0.023 0.11, 0.24 –5.1 0.36Without δ18O outliers 0.18 0.025 0.11, 0.25 –5.2 0.35

Middle Blancan –0.71 0.138 –1.10, –0.33 –1.95 0.07Early Blancan –0.06 0.006 –0.07, –0.04 –2.6 <0.01Hemphillian –0.13 0.046 –0.29, 0.02 –6.0 <<0.01Clarendonian 0.27 0.054 0.10, 0.44 –9.5 0.59

Note: Regressions and confi dence intervals (CI) were calculated using the standard linear approximation of Sokal and Rohlf (1994) in RMA for JAVA v. 1.21 (Bohonak and van Linde, 2004). Abbreviations: S.e.—standard error. Bold indicates regressions for which the 99% confi dence interval for the slope does not include zero.



Fox et al.


regression, the trend is signifi cant (Table 2), but, again, it explains a small portion of the variance (R2 = 0.07). We do not have internal age control on this section other than its biostratigraphic correspondence to other mid-Blancan faunas in the area, but given that these short sections lack well-developed caliches, the mid-Blancan sec-tions probably represent the shortest duration of sedimentation. The average proportion of C

4

biomass in the middle Blancan section is identi-cal to that for the early Blancan (38%), and the range is narrower (30%–52%; Table 1). All δ13C values are high enough that the estimates of C

4

biomass are well outside the end member for arid-condition C

3 biomass. Under extreme arid-

ity, the range for percent C4 biomass would shift

to 13%–40%. No estimates of the percent C4 bio-

mass fall in the modern range.

Late Pliocene–Early Pleistocene: Late Blancan–Early Irvingtonian (2.5–1.0 Ma) Sections

The late Blancan–early Irvingtonian samples (Fig. 5) have good geochronological control from magnetostratigraphy and the ages on the interbedded Huckleberry Ridge and Cerro Toledo B ashes (Martin et al., 2008). The mean δ13C value in this interval is −2.6% ± 1.18‰ (Table 1). The minimum δ13C value in the late Blancan–early Irvingtonian section is greater than the mean for the middle Blancan section, suggesting a persistent increase in the abun-dance of C

4 biomass. Nevertheless, variation

in δ13C is considerable in the portion of the composite section in which all three measured sections overlap. The δ13C values have a statisti-cally signifi cant trend toward higher values up section (0.18‰ per meter; Table 2). The trend in δ13C values corresponds to an increase up sec-tion in the estimated abundance of C

4 biomass

from ~50% at the base of the Borchers section (i.e., below the Huckleberry Ridge ash and the Borchers fauna) to 60%–80% at the top of the section (Fig. 5). The range in measured δ13C values corresponds to a range in percent C

4 for

average conditions of 43%–78% (Table 1). All measured values are outside the range of δ13C values for arid-condition C

3 biomass; under

extreme aridity, the measured values would cor-respond to a range in C

4 biomass of 28%–72%.

Several samples in the upper part of the com-posite section reach abundances of C

4 biomass

in the modern range.The stratigraphic offset of the Cerro Toledo

B ash in the Borchers section north of Kansas State Highway 23 and the Aries section is based on magnetostratigraphic correlation and rep-resents paleotopography (Martin et al., 2008). If these two sections are correlated simply by

meter level at the base of the Cerro Toledo B ash (and the two Borchers sections are still cor-related on the base of the Huckleberry Ridge ash), RMA linear regression with meter level as the independent variable and δ13C value as the dependent variable still has a slope that is statis-tically signifi cant (all data, δ13C = 0.18 × meters – 5.1, R2 = 0.36, 99% CI for slope: 0.13–0.24; compare to Table 2). Consequently, the strati-graphic trend in the late Blancan–early Irvingto-nian δ13C values is insensitive to the correlation of the Aries section with the Borchers sections.

DISCUSSION

In the following discussion, we fi rst compare the two Miocene sections in the Meade area to the previously published data from Miocene carbonates elsewhere in the Great Plains (Fox and Koch, 2003, 2004) to evaluate the presence of C

4 biomass in the Great Plains during the

Miocene. After that, we consider the possibility of diagenesis or recrystallization of pedogenic carbonate as an explanation for the homogeneity of Miocene carbonate δ13C values. Finally, we examine the long-term trend in the abundance of C

4 biomass in southwest Kansas through sta-

tistical comparisons of the data presented here.

Meade Basin Miocene Sections

The interpretation of Fox and Koch (2003, 2004) that Miocene carbonates in the Great Plains indicate 10%–30% C

4 biomass in the

region throughout the Miocene remains some-what controversial because of the implication of earlier appearance of C

4 grasses in the region

compared to other studies (e.g., Cerling et al., 1997; Passey et al., 2002) and the interpreta-tion of δ13C values of organic matter occluded in carbonate (e.g., Wang and Deng, 2005). The presence of multiple Miocene sections in the Meade area allows us to consider again the evidence for the abundance of C

4 biomass dur-

ing the Miocene. We used pair-wise Student’s t-tests and Mann Whitney U-tests to compare the data from the two Miocene sections in the Meade area, both separately and combined, to the δ13C values of Miocene paleosol carbonates and estimated abundance of C

4 biomass for the

Great Plains from Fox and Koch (2003, 2004; Fig. 6). Results of these statistical tests are dis-cussed in the text and presented in Table DR4 (see footnote 1).

Estimates of percent C4 for the data from

Fox and Koch (2003, 2004) are based on the end members in Figure 3, not the original esti-mates calculated by Fox and Koch (2003, 2004). Mean Clarendonian δ13C values and percent C

4

biomass in Meade are not statistically different

from those for the only Clarendonian section in Kansas from Fox and Koch (2003, 2004), Minium Quarry (mean δ13C value = −7.0‰ ± 0.43‰, mean percent C

4 = 21.5% ± 2.89%,

n = 7). If the data from Minium Quarry are combined with those from the other Clarendon-ian sections in Fox and Koch (2003, 2004), the Corman Ranch section in Nebraska and Port of Entry Pit in Oklahoma, the averages are slightly higher (−6.9‰ ± 0.46‰ and 22.1% ± 3.07%, respectively; n = 29 for the combined sec-tions), and the slightly larger differences with the Meade Clarendonian section are statistically signifi cant. The differences between the Meade Clarendonian data using the means for the com-bined Hemphillian and Clarendonian data from the Great Plains (−6.8‰ ± 0.76‰ and 23.6% ± 5.09%, n = 153) and the total Miocene data from the Great Plains (−6.8‰ ± 0.83‰ and 23.5% ± 3.0%, n = 230) in Fox and Koch (2003, 2004) are also statistically signifi cant. The somewhat lower mean δ13C value and low abun-dance of C

4 biomass during the Clarendonian

in the Meade area, and perhaps western Kansas in general, appear to represent local variation in comparison to the Great Plains as a whole, although the differences are small (less than 1‰ and 7% C

4 biomass on average). However, the

δ13C values in the Meade Clarendonian section do increase up section toward more typical val-ues for the Clarendonian in the Great Plains, and the maximum δ13C value in the Clarendonian section (−5.9‰) is identical to that for the larger Clarendonian data set for the Great Plains.

Results of statistical comparisons of the Hemphillian data from Meade somewhat vary depending on the test used. Based on t-tests, the mean δ13C value and mean abundance of C

4 bio-

mass for the Hemphillian section in Meade are not statistically distinct from the mean for the 12 Hemphillian sections elsewhere in the Great Plains (−6.7‰ ± 0.81‰, n = 124) and the means of the combined Hemphillian and Clarendon-ian data from the Great Plains in Fox and Koch (2003, 2004). For these same comparisons, Mann-Whitney U-tests indicate statistically sig-nifi cant differences in mean value for all four comparisons. Both tests indicate that the means for the Hemphillian in Meade are statistically different from those for the entire Miocene data set of Fox and Koch (2003, 2004). In all three comparisons, the differences in mean values are even smaller than those for the Clarendonian comparisons and also in the opposite direction (i.e., Meade Hemphillian values suggest more rather than less C

4 biomass locally than else-

where in the Great Plains on average).If the data from the two Miocene sections

in the Meade area are combined and treated as one group, the mean δ13C value and the mean





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NE, OK, TX, NM

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Figure 6. Histograms of δ13C values (A–E) and percent C4 biomass for each biostratigraphic interval (F–J) and for the modern (K). (A, F) δ13C values and percent C4 for Miocene sections elsewhere in the Great Plains, including a Clarendonian locality in Kansas, from Fox and Koch (2003, 2004). (B, G) Miocene sections from the Meade Basin. (C, H) Early Blancan sections. (D, I) Middle Blancan sections. (E, J) Late Blancan–early Irvingtonian sections. (K) Modern abundance of C4 biomass in Kansas and northern Oklahoma (Table DR1 [see text footnote 1]). Bin labels indicate the lowest value included in each bin. Bins are 0.5‰ wide for δ13C values and 5% wide for percent C4 biomass. VPDB—Vienna Peedee belemnite. KS—Kansas; NE—Nebraska; OK—Oklahoma; TX—Texas; NM—New Mexico.



Fox et al.


abundance of C4 biomass during the Miocene in

the Meade area are not statistically distinguish-able from either the combined Hemphillian and Clarendonian data or the entire Miocene data set of Fox and Koch (2003, 2004). The Mio-cene data from the Meade Basin might support a local increase in the abundance of C

4 biomass

during the Miocene (statistical signifi cance of the local changes in δ13C value and percent C

4

are assessed later herein), but comparison of the Meade Miocene data to the larger data set from elsewhere in the Great Plains does not suggest that ecosystem composition around Meade dur-ing the late Miocene was substantially different from elsewhere in the Great Plains. Thus, the Miocene data from Meade are consistent with the conclusion of Fox and Koch (2003, 2004) that Miocene δ13C values of paleosol carbon-ates across the Great Plains are suffi ciently high to imply ~20% C

4 biomass on average and are

more or less spatially homogeneous.Diagenesis is not a reasonable explanation for

δ13C values of Miocene pedogenic carbonates in the Great Plains (including those described here from the Meade Basin), which exhibit low spa-tial variability and are suffi ciently high to imply an ecologically signifi cant abundance of C

4 bio-

mass throughout the region. Petrographic thin sections of 19 fi eld samples (not shown) from two early Blancan measured sections in the Meade Basin section have no secondary (e.g., crack-fi lling) sparry calcite, consistent with observations of Fox and Koch (2003, 2004) on thin sections of other Neogene pedogenic carbonates in the Great Plains. Moreover, we did not observe macroscopic sparry calcite in any fi eld samples during sampling for isotope analyses. Nevertheless, these observations do not rule out cryptic recrystallization of micritic carbonate through interaction with groundwa-ter in the unconfi ned High Plains aquifer of the Great Plains. However, Fox and Koch (2003, 2004) pointed out that δ18O values of Miocene pedogenic carbonates in the region preserve geographic and temporal variability that would have been erased by recrystallization, and as discussed in Fox et al. (2011), the δ18O record from Meade also preserves sensible temporal variation. For δ13C values with low variability across the Great Plains to result from recrys-tallization, groundwaters would be required to have been consistently out of equilibrium with calcite and had isotopically homogeneous dis-solved inorganic carbon (DIC). However, Ogal-lala waters are generally in equilibrium with calcite locally, and so they should not be dis-solving or precipitating calcite, and DIC in the Ogallala has both regionally and vertically vari-able δ13C values (Fryar et al., 2001; McMahon et al., 2004, 2007). Along fl ow paths between

wells in the Great Plains aquifer in southwestern Kansas, elemental and mineralogical mass bal-ance models call for precipitation of calcite as cements with δ13C values between −6.5‰ and −4.1‰, which are in good agreement with mea-sured values of cements from bore hole cuttings (McMahon et al., 2004). A physical mixture of this diagenetic calcite formed below the water table and preexisting Miocene pedogenic car-bonate with end-member C

3 values, or partial

replacement of the latter by the former on small spatial scales, could yield measured values for Miocene carbonates now exposed at the sur-face that imply a low proportion of C

4 biomass

(i.e., ~20% C4 as estimated here and in Fox and

Koch, 2003, 2004). However, the geochemistry of the High Plains aquifer in both northern Texas (Fryar et al., 2001) and Nebraska (McMahon et al., 2007) calls for dissolution, not precipitation, of calcite to balance groundwater chemistry along fl ow paths. The carbonates discussed here and reported by Fox and Koch (2003, 2004) do not generally have evidence of late cements. Additionally, the temporal variability in both δ13C and δ18O values reported here and in Fox et al. (2010) argues strongly against diagenesis as an explanation for the similarity of δ13C values across the Great Plains during the Miocene.

Long-Term Trend in Abundance of C4 Biomass in the Meade Basin Record

The regressions between meter level and δ13C values for each composite section in the Meade Basin (Fig. 5; Table 2) indicate that the individ-ual sections do not record a consistent pattern of short-term changes in ecosystem composition and climate. The Clarendonian section trends up section to higher δ13C values, the Hemphillian section has no statistically signifi cant trend, the early and middle Blancan sections have decreas-ing trends in δ13C values, and the late Blan-can–early Irvingtonian sections have a trend to higher values. However, one-way ANOVAs for both δ13C values (F = 81.9, p < 0.001) and esti-mates of the percentage of C

4 biomass within

intervals for the Meade data and for the modern data (F = 180.2, p < 0.001) indicate statistically signifi cant differences in mean values among intervals (Fig. 6). The results are the same if the Hemphillian and Clarendonian data are grouped together (δ13C: F = 103.0, p < 0.001; percent C

4: F = 216.1, p < 0.001). To examine

the long-term trend in δ13C values and the abun-dance of C

4 biomass, we used both the post hoc

pair-wise Scheffé test for multiple comparisons (Table DR5 [see footnote 1]) and pair-wise Mann-Whitney U-tests (Table DR6 for δ13C val-ues and Table DR7 for percent C

4 [see footnote

1]) to identify the intervals in the Meade Basin

that have statistically signifi cant differences in mean values and the intervals that are statisti-cally different from the modern abundance of C

4

biomass in Kansas and northwestern Oklahoma (Table DR1 [see footnote 1]). The Scheffé test is appropriate given the variability in sample sizes; the use of both parametric and nonparametric tests is appropriate because the data are not nec-essarily distributed normally in every interval.

Comparison of the Clarendonian and Hemp-hillian data is ambiguous. The Scheffé test is not statistically signifi cant for either δ13C values or percent C

4, but the Mann-Whitney U-test is sig-

nifi cant for both; this result refl ects the difference in statistical power between the two tests given the sample sizes and data. It is possible that the Meade data record an increase in the abundance of C

4 grasses from some time during the Claren-

donian to sometime during the Hemphillian, but the increase is relatively small. However, given the contrasts between the Clarendonian data and various partitions of the data of Fox and Koch (2003, 2004), an equally plausible explanation is that the Clarendonian in western Kansas had, on average, an unusually low abundance of C

4

biomass compared to elsewhere in the region during the Miocene. Additionally, the Clarendo-nian data exhibit a statistically signifi cant strati-graphic trend in δ13C values (Fig. 5; Table 2) that reaches values higher than the mean δ13C value of the Hemphillian section, so it is possible that the Clarendonian section, which has no internal geochronological control, only captures a short-term, local fl uctuation in the abundance of C

4

biomass at the landscape scale. Regardless, the absolute differences between the Clarendonian and Hemphillian values are smaller than those for the other intervals.

The Scheffé tests and the Mann-Whitney U-tests indicate that mean δ13C value and per-cent C

4 biomass increase signifi cantly from the

Hemphillian to the early Blancan. The result is the same if the combined Clarendonian- Hemphillian data for Meade are compared to the early Blancan data. Without more precise age or stratigraphic control on the top of the Hemp-hillian section, it is impossible to say whether this increase was abrupt or gradual. Given that the early Blancan section is not physically superposed on the Hemphillian, the change appears to be abrupt, but incomplete strati-graphic preservation might be masking a more gradual transition. However, the Hemphillian does not have a signifi cant stratigraphic trend, and the statistically signifi cant (though weak) trend in the entire early Blancan data is toward lower δ13C values, even if some individual early Blancan sections suggest the possibility of short-term stratigraphic trends toward higher values. Indeed, the base of the composite early





Blancan section (in the XIT-B section) actually has higher δ13C values than the Saw Rock Can-yon section (Fig. DR3 [see footnote 1]), which is ~20 m above the base of the composite sec-tion and includes the oldest Pliocene fauna in the area and possibly in North America (Bell et al., 2004). Regardless of the age of the top of the Hemphillian section, the initial increase in the abundance of C

4 grasses above levels typical for

the Miocene in the Great Plains must have taken place during the latest Hemphillian or earliest Blancan, prior to either the Saw Rock Canyon fauna or deposition of the lowermost sediments in the XIT sections.

Neither test indicates statistically signifi cant differences in mean δ13C values or percent C

4

biomass between the early and middle Blancan, which is not surprising given the nearly identical mean values for these two intervals (Figs. 5 and 6; Table 1). Although both sections have strati-graphic trends toward lower δ13C values, nei-ther regression explains much variance in δ13C values, and overall the data suggest that, after an initial increase in the abundance of C

4 bio-

mass during the latest Hemphillian or earliest Blancan, the average abundance of C

4 grasses

stayed more or less constant in the area for as long as 2.5 m.y. Notably, both the early and middle Blancan composite sections exhibit con-siderable variation in δ13C values where mul-tiple measured sections overlap and sampling is densest (e.g., at ~40 m in the early Blancan in Fig. 5 and below 5 m in the middle Blancan). It is not clear from these data if this variabil-ity results simply from short-term and small- spatial-scale variations in the C

3:C

4 ratio in a

more or less homogeneous savanna woodland or treeless prairie (the nature of the C

3 compo-

nent is not well constrained), or if the variations are more patterned temporally and/or spatially. The difference between these two explanations of the variations has important implications for the nature of ecosystems during the transi-tion to the modern C

4-dominated grassland of

the region and for the evolution of the fauna in the region during the Blancan. We are currently addressing the temporal and spatial pattern of this variability with a detailed study of multiple short stratigraphic sections in early Blancan out-crops in Keefe Canyon.

Both sets of statistical tests indicate that both mean δ13C value and percent C

4 biomass

increase signifi cantly in the late Blancan–early Irvingtonian composite section in comparison to the middle Blancan composite section. Although the base of the late Blancan–early Irvingtonian section is close to the mean value for the mid-dle Blancan, the former has a statistically sig-nifi cant stratigraphic trend toward higher values (Fig. 5; Table 2), despite considerable variabil-

ity in the middle part of the section. Thus, the composite late Blancan–early Irvingtonian sec-tion records a second increase in the abundance of C

4 biomass following the protracted interval

of little net change during the early and middle Blancan. The strong stratigraphic trend in δ13C values during the late Blancan–early Irvingto-nian indicates that the abundance of C

4 biomass

increased during the latest Pliocene and early Pleistocene in the region. The abundance of C

4

biomass fi rst reaches modern levels typical for Kansas and northern Oklahoma (78% ± 10.9%; Fig. 6; Table 2) around the level of the Cerro Toledo B ash in the northern Borchers (22 m) and Aries (16 m) sections (see Fig. DR5 for detailed stratigraphy in these sections [see foot-note 1]). Thus, the fi rst appearance of a modern-like ecosystem in the region, at least in terms of the abundance of C

4 grasses, occurs around

the interval 1.47–1.23 Ma. However, both the Scheffé and Mann Whitney U-tests indicate that the mean abundance of C

4 biomass during the

late Blancan–early Irvingtonian interval was signifi cantly less than the modern abundance in the region, despite the presence of a few samples that have δ13C values consistent with the modern abundance of C

4 biomass. The result is the same

for the more conservative comparisons of the late Blancan–early Irvingtonian data using the mean for the combined Holocene and modern C

4 data (75% ± 12.8%; one-way ANOVA: F =

172.9, p < 0.001; Scheffé test for late Blancan–early Irvingtonian and combined Holocene and modern C

4 data, p < 0.001). This could imply that

the fi nal appearance of a permanently sustained grassland with modern abundance of C

4 biomass

happened after the end of Borchers/Aries time. However, given the variance in the modern data and the temporal variations in C

3:C

4 ratios in

some well-dated late Pleistocene records of soil organic matter δ13C values from the region (e.g., Johnson et al., 2007), an alternative explanation is that the modern domination of the region by C

4 grasses represents one end of a continuum of

habitat composition. In this case, the trend in the late Blancan–early Irvingtonian, and potentially the great variability in the well-sampled interval of the early Blancan, indicates a more general pattern of small- or even regional-scale spatial and temporal variability in the abundance of C

4 biomass. The drivers of such variability could

be the dynamics and interactions of climate, fi re, herbivores, and seed predators (Janzen, 1971; Edwards and Crawley, 1999; Bond and Keeley, 2005; Bond, 2008; Osborne, 2008; Edwards et al., 2010).

The statistically signifi cant temporal changes in δ13C values also have implications for the possibility of diagenesis or recrystallization as explanations for the temporal and spatial homo-

geneity of δ13C values during the Miocene in the Great Plains. Given that the Pliocene sediments in the Meade Basin would also be part of the unconfi ned High Plains aquifer (along with the Ogallala Formation), particularly prior to inci-sion of modern drainages, the statistically signif-icant stratigraphic trends and changes in mean δ13C values cannot be explained by interaction with the same groundwaters that putatively would have caused diagenesis and homogeniza-tion in the Miocene sections, both locally and regionally. Thus, temporal variations reinforce the conclusion reached previously that recrys-tallization or precipitation of diagenetic calcite from High Plains aquifer groundwater is not a plausible explanation for the regional homoge-neity of Miocene pedogenic carbonate in the Great Plains. Instead, the presence of an average of ~20% C

4 biomass throughout the region dur-

ing the Miocene, only increasing substantially over the Pliocene and into the early Pleistocene, remains the best explanation.

CONCLUSION

Based on the carbon isotope composition of Neogene carbonates in the Meade Basin, we have reached three conclusions about the history of C

4 biomass in the southern Great Plains. First,

the trends and changes in mean δ13C values and percent C

4 abundances are much larger across

the entire study period than they are within individual biostratigraphic intervals. Second, the δ13C values of paleosol carbonates record a three-stage increase in the abundance of C

4 bio-

mass in the area from the Clarendonian (early late Miocene, 12–9 Ma) to the early Irvingtonian NALMA (early Pleistocene, 2–ca. 1 Ma). Mio-cene δ13C values in the Meade Basin are consis-tent with ~20% C

4 grasses in the ecosystem and

are statistically similar to Miocene values else-where in the Great Plains (Fox and Koch, 2003, 2004), although it is possible that C

4 grasses

were somewhat less abundant during the Clar-endonian in Kansas than further north or south. Between the end of the Hemphillian record (late late Miocene, 9–5 Ma) and the beginning of the Blancan record (Pliocene, 5–2 Ma) in the Meade Basin, the abundance of C

4 grasses

increased to ~40% of biomass and remained relatively constant for up to 2.5 m.y. during the early and middle Blancan. During the last stage, the abundance of C

4 biomass increased steadily

during the late Blancan and early Irvingtonian NALMAs, fi rst reaching modern levels for the region (78% ± 10.9%) around 1.3 Ma. In addi-tion, during the protracted interval of relatively stable abundance of C

4 biomass during the early

and middle Blancan, the most densely sampled intervals exhibit the greatest variability in C

4



Fox et al.


biomass, suggesting that landscape-scale varia-tion in ecosystem composition was high. The potential of high variability in the abundance of grasses, assuming that the C

3 component in

the late Neogene was trees and/or shrubs and not grasses, has important implications for the role of habitat variability in the evolution of the modern mammal fauna in the region. However, the nature of the C

3 component of the ecosystem

(i.e., whether it was woody or grassy vegetation) is not yet well constrained.

In this study, we have answered the question of when C

4 grasses came to dominate the grass-

land ecosystem of the southern Great Plains. Two of the 106 samples from the early Blancan (4.9–3.0 Ma) provide the fi rst evidence for C

4

grasses constituting more than 50% of biomass, but only under the assumption that water and light stress did not shift C

3 biomass to higher

δ13C values. Several samples in the middle Blan-can (3.0–2.5 Ma) also surpass 50% C

4 biomass,

again under the assumption that aridity and light stress were not major factors. C

4 biomass fi rst

reaches modern abundance for the region (78% ± 10.9%) around 1.3 Ma, although C

4 grasses

may not have remained at that level persistently until the present, based on high variability in δ13C values of organic matter in late Pleistocene paleosols from Kansas (Johnson et al., 2007). The rise to ecological dominance of C

4 grasses

in the region did happen during the late Neo-gene, but it was not restricted to a narrow inter-val in the latest Miocene. Instead, the increase was protracted and not monotonic.

Our study does not resolve the contrast between the record of C

4 abundance in the Great

Plains from dietary reconstructions of fossil horses (Passey et al., 2002) and that from paleo-sol carbonates (Fox and Koch, 2003, 2004). However, some clades of large herbivorous mammals maintained end-member C

3 diets

during the Miocene (e.g., proboscideans; Fox and Fisher, 2004) despite evidence from coeval paleosol carbonates for ~20% C

4 biomass in the

region (Fox and Koch, 2003, 2004; this study). Moreover, Koch et al. (2004) have shown, based on dietary reconstructions of late Pleisto-cene large herbivorous mammals in Texas and regional-scale climate-vegetation modeling, that horses may not be as sensitive indicators of ecosystem C

3:C

4 ratio as has been assumed,

at least in the late Pleistocene of the southern Great Plains.

In the companion paper in this issue (Fox et al., 2011), we discuss the climatic context for the long-term increase in the abundance of C

4

grasses based on the δ18O values of the pedo-genic carbonates discussed here. We also com-ment on the possible cause or causes of the increase in the abundance of C

4 grasses in the

southern Great Plains, and by extension else-where, during the Neogene.

ACKNOWLEDGMENTS

This research would not have been possible with-out access to private land granted by numerous land-owners in the Meade area. Jenn Campbell, Mark Clementz, Neil Kelly, Sam Matson, and Seth New-some provided excellent help with fi eld work for this project over several seasons. We thank Sam Matson for comment on a draft of this paper and four thought-ful reviews by Bill Gilhooly, Andrea Lini, Nate Shel-don, and one anonymous reviewer. This research was funded by grants from the National Science Foun-dation to Fox (EAR-0207383) and to Martin (EAR-0207582) and from the National Geographic Society to Martin (5963-07, 6547-99) with matching funds to Martin from Murray State University.

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